ARCHAEAN GEOLOGY SECOND INTERNATIONAL SYMPOSIUM PERTH 1980
Editors: J. E. Glover and D. I. Groves
1981 THE GEOLOGICAL SOCIETY OF AUSTRALIA INCORPORATED
Office Bearers 1981-1982 (Elected 20 August 1981) President: R. D. GEE Vice-Presidents: C. D. BRANCH, M. J. RICKARD Hon. Secretary: R. E. SMITH Hon. Treasurer. K. K. SAPPAL Hon. Editor. R. H. VERNON
Front cover: Landsat satellite image of part of the Pilbara Block, showing Archaean greenstones between the Shaw Batholith (southwest), Corunna Downs Batholith (centre), Mt Edgar Batholith (north) and a strip of Mosquito Creek Beds (east centre). The image was processed by CSIRO Division of Mineral Physics and supplied by K. G. McCracken.
ARCHAEAN GEOLOG Second International Symposium Perth 1980
ARCHAEAN GEOLOGY Second International Symposium Perth 1980 »
Editors J. E. GLOVER and D. I. GROVES
I Special Publication No. 7 GEOLOGICAL SOCIETY OF AUSTRALIA INCORPORATED
December 1981
1981 THE GEOLOGICAL SOCIETY OF AUSTRALIA INCORPORATED
This volume is published in the Special Publications series of the Geological Society of Australia REFERENCES General reference: GLOVER, J. E., & GROVES, D. I. (Eds), 1981: Archaean Geology: Second International Symposium, Perth, 1980. Spec. Pubis geol. Soc. Aust., 7. Two forms of reference to specific papers are possible, as follows: MCCALL, G. J. H., 1981: Progress in research into the early history of the Earth: a review, 1970-1980. Spec. Pubis geol. Soc. Aust., 7\ 3-18.. MCCALL, G. J. H., 1981: Progress in research into the early history of the Earth: a review, 1970-1980; in Glover, J. E., & Groves, D. I. (Eds) Archaean Geology: Second International Symposium, Perth, 1980. Spec. Pubis geol. Soc. Aust., 7, 3-18.
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National Library of Australia Cataloguing-in-Publication
data ,
Archaean geology: second International Symposium, Perth, 1980. Includes bibliographical references.
ISBN 0 909869 19 7.
1. Geology, Stratigraphic—Archaean-Congresses. I. Glover, J. E. (John Edmund), 1924- . II. Groves, D. I. (David Ian, 1942-. III. Geological Society of Australia. (Series: Special publication (Geological Society of Australia); no. 7).
551.7*12
PREFACE The concept of a Second International Archaean Symposium to review progress in understanding the early evolution of the Earth over the decade 1970-1980 was formulated by J. H. Lord, then Director of the Geological Survey of Western Australia. The Symposium coincided with the tenth anniversary of the first Archaean Symposium The Archaean Rocks" held in Perth in 1970. The Symposium was organized under the auspices of the Geological Society of Australia and International Geological Correlation Programme Project "Archaean Geochemistry'' by a committee comprising P. G. Harris (chairman), R. D. Gee, D. I. Groves, J. A. Hallberg, J. H. Lord, R. J. Marston, J. E. Martin and D. L. Rowston. Over 450 earth scientists attended the Symposium which was held between May 12 and May 16, 1980, at the University of Western Australia in Perth. The success of the Symposium, and of the field excursions, owed much to the help of the following organizations and companies, listed alphabetically: Central Norseman Gold Corporation N.L., Commonwealth Scientific and Industrial Research Organisation, Geological Survey of Western Australia, Greenbushes Tin N.L., University of Western Australia, and Western Mining Corporation Ltd. Forty-two of the calO papers presented appear in this volume, and the editors acknowledge the work of referees in ensuring a high standard of published papers. Extended abstracts of all papers presented were previously published in an Extended Abstracts Volume. During the decade that separated the two Archaean symposia in Perth, research on the early history of the Earth proliferated and diversified, and major advances were made in geochronology, sedimentology, geochemistry, metallogeny and paleobiology. The advances in these fields are reflected in the papers in this volume and were discussed by G. J. H. McCall in the introductory address to the Symposium. Much of our understanding of the early history of the Earth can be attributed to advances in isotope geology and geochronology, which provide petrogenetic constraints and a temporal framework for other studies. The importance of integration of structural/metamorphic studies with isotopic data in understanding crustal evolution is demonstrated throughout the volume. The use of the multiple isotopic systems Rb-Sr, Pb-Pb, UPb in zircon, and Sm-Nd to develop the total history of complex terrains is rapidly becoming standard practice. These integrated isotopic studies firmly established the existence of crust 3.5 b.y. old and up to 3.8. b.y. old in several high-grade gneiss belts, and the existence of greenstone sequences 3.5 b.y. old in several cratons. World-wide thermal events resulting in greenstone formation and widespread granitoid plutonism are evident at 2.82.6 b.y., and much of the mineral wealth relates to these events. The isotopic data presented in this volume suggest that the early continental crust evolved both by addition of juvenile sial, produced by melting of mantle-derived source rocks, and by re-working of pre-existing sial. The regional geology of many of the major Archaean cratons of the world is now becoming well established, as shown in a number of reviews in the volume. Petrogenetic models for the evolution of Archaean volcanic suites, particularly the komatiite suite, now have a considerable geochemical data-base and the relationships between volcanics ascribed to komatiitic, tholeiitic and calc-alkaline lineages are understood in some areas where sound field-based studies have been carried out. Integrated field, geochronological and geochemical data illuminate the history of the extensive granitoid-gneiss terrains, and geochemical studies of sedimentary rocks are used to estimate the change in composition of the crust with time. Data presented in this volume indicate broad similarities and important differences between cratons, and warn against over-generalized or all-embracing models for evolution of Archaean terrains. Controversy continues over the kind of crust, if any, on which greenstone sequences were deposited. Detailed studies of relationships between granitoid gneisses and greenstones described in this volume reveal tectonic contacts between them, and structural and isotopic data are commonly equivocal. Many studies suggest that most gneisses represent newly 4
vi
PREFACE
generated, rapidly reworked sial formed contemporaneously or shortly after the greenstone belt sequence although more restricted areas of older gneisses are commonly known from adjacent zones. Structural and metamorphic studies have been used in conjunction in some regions to explain the post-formational, tectonic and thermal history of greenstone and gneiss terrains. In some areas, gneisses and greenstones share a common deformational history. Despite these studies, no general model is currently accepted to explain the relationships between granitoids, gneisses and greenstones, and the tectonic setting of the greenstone belts remains equivocal. Direct comparisons both with modern tectonic settings and models appealing to evolving tectonic processes with progressive evolution of the crust are made in this volume. Models involving back-arc basins or intracratonic shallow rift basins remain most favoured. The importance of Archaean sedimentology in understanding the tectonic setting of greenstone belts is emphasized. At least some greenstone belts appear to have been shallow basins throughout their formation, with addition of mainly volcanic products matching subsidence. The greenstone basins were quite unlike those of modern deep-ocean plains. Fluvial environments transitional to troughs with narrow shelf environments and deeper water suggest the presence of continental margins during the final stages of some greenstone basins. Cratonic areas incorporating greenstones appear to have been present at that stage. The sedimentary rocks also provide information on the habitat of the earliest life-forms which apparently developed by ca3.5 b.y. However, debate continues about oxygen partial pressures in the Archaean atmosphere and the oxidation potential and temperature of the early hydrosphere. These aspects have important implications for Archaean metallogeny. This volume indicates our state of knowledge of Archaean terrains as at 1980 and highlights some of the major problems that persist. Archaean terrains not only remain a major frontier for field geologists, but a challenge to geophysicists, petrologists, geochemists, palaeontologists and others: integrated studies are clearly essential. We hope that the proceedings of the Third International Archaean Conference planned for Perth in 1990 will resolve many of the problems raised here, and that this volume may be used as a yardstick to measure progress in understanding early Earth evolution in the next decade.
CONTENTS Page
INTRODUCTION
G . J . H . MCCALL:
Progress in Research into the Early History of the Earth: a Review, 1970-1980
3
REGIONAL REVIEWS OF ARCHAEAN TERRAINS J . M . BARTON JR:
The Pattern of Archaean Crustal Evolution in Southern Africa as Deduced from the Evolution of the Limpopo Mobile Belt and the Barberton Granite-Greenstone Terrain
21
B. J . DRUMMOND, R . E . SMITH, & R . C . HORWITZ:
Crustal Structure in the Pilbara and Northern Yilgarn Blocks from Deep Seismic Sounding
R . D . GEE, J . L . BAXTER, S. A . WILDE, & I. R . WILLIAMS:
Crustal Development in the Archaean Yilgarn Block, Western Australia .
A . H . HICKMAN:
Crustal Evolution of the Pilbara Block, Western Australia
33 43
....
57
A Review of the Structural Evolution and Geochronology of the Archaean Napier Complex of Enderby Land, Australian Antarctic Territory . . . . . ... . .... ... . . . . ... . . ; . . . . . .... ...
71
P R . JAMES & L . P . BLACK:
P . K . SIMS & Z . E . PETERMAN:
Archaean Rocks in the Southern Part of the Canadian Shield—a Review.
W . A . PADGHAM:
Archaean Crustal Evolution—a Glimpse from the Slave Province. . . . . . .
V . I. KAZANSKY & V. M . MORALEV:
Archaean Geology and Metallogeny of the Aldan Shield, USSR . . . . . . . .
85 99 Ill
SUN DAZHONG & W u CHANGHUA:
The Principal Geological and Geochemical Characteristics of the Archaean Greenstone-Gneiss Sequences in North China
E . WERNICK:
The Atlantic Granulite Belt, Brazil
121
....
133
The Older Precambrian Geochronology of Western Australia. . . . . . . . . .
145
GEOCHRONOLOGY
J . R . DE LAETER, W . G . LIBBY, & A . F . TRENDALL: D . A . NIEUWLAND & W . COMPSTON:
Crustal Evolution in the Yilgarn Block near Perth, Western Australia . . .
159
H . J . CHAPMAN, M . J . BICKLE, J . R . DE LAETER, L . F . BETTENAY, D . I. GROVES, L . S. ANDERSEN, R . A . BINNS, & M . GORTON:
Rb-Sr Geochronology of Granitic Rocks from the Diemals Area, Central Yiigarn Block, Western A u s t r a l i a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
173
CONTENTS
viii
P . J . HAMILTON, N . M . EVENSEN, R . K. O'NIONS, A . Y. GLIKSON, & A.H.HICKMAN:
Sm-Nd Dating of the North Star Basalt, Warrawoona Group, Pilbara Block, Western Australia •
187
J . F . LOVERING, G . A . TRAVIS, D . J . COMAFORD, & P . R . KELLY:
Evolution of the Gondwana Archaean Shield: Zircon Dating by Ion Microprobe, and Relationships between Australia and Wilkes Land (Antarctica)
193
K. D . COLLERSON, A . KERR, & W . COMPSTON:
Geochronology and Evolution of Late Archaean Gneisses in Northern Labrador: an Example of Reworked Sialic Crust. . . . . . . . . . . . . . . . . . . .
205
ARCHAEAN SEDIMENTARY AND VOLCANIC ROCKS J . S. R . DUNLOP & R . BUICK:
Archaean Epiclastic Sediments Derived from Mafic Volcanics, North Pole, Pilbara Block, Western Australia . . . . . . . . . . . . . . . . . . . . . . . ., . .
225
K. A . ERIKSSON:
Archaean Platform-to-Trough Sedimentation, East Pilbara Block, Australia
235
S. M . NAQVI, P . K. GOVIL, & J . J . W . ROGERS:
Chemical Sedimentation in Archaean-Early Proterozoic Greenschist Belts of the Dharwar Craton, India. . .
245
S. R . TAYLOR & S. M . MCLENNAN:
Evidence from Rare-Earth Elements for the Chemical Composition of the Archaean Crust. : ..
255
M . E . BARLEY:
Relations Between Volcanic Rocks in the Warrawoona Group: Continuous or Cyclic Evolution?. ...
263
C . W . GILES:
Archaean Calc-alkaline Volcanism in the Eastern Goldfields Province, Western Australia
275
A . Y. GLIKSON & A . H . HICKMAN:
Geochemical Stratigraphy and Pedogenesis of Archaean Basic-Ultrabasic Volcanic Units, Eastern Pilbara Block, Western Australia .
287
C . R . L . FRIEND, R . P . HALL, & D . J . HUGHES:
The Geochemistry of the Malene (Mid-Archaean) Ultramafic-Mafic Amphibolite Suite, Southern West Greenland ..
301
R . C . O . GILL, D . BRIDGWATER, & J . H . ALLAART:
The Geochemistry of the Earliest Known Basic Metavolcanic Rocks, at Isua, West Greenland: a Preliminary Investigation . . . . . . . . . . . . . . . . . .
313
ARCHAEAN GRANITOIDS AND GNEISSES R . E . P . FRIPP:
The Ancient Sand River Gneisses, Limpopo Mobile Belt, South Africa . .
329
L. J . ROBB:
Detailed Studies of Select Migmatite Outcrops in the Region Southwest of the Barberton Greenstone Belt and their Significance Concerning the Nature of the Early Archaean Crust in the Region ..
337
CONTENTS
ix
J . S. MYERS:
The Fiskenaesset Anorthosite Complex—a Stratigraphic Key to the Tectonic Evolution of the West Greenland Gneiss Complex 3000-2800 m.y. Ago r; . ...
351
L . F . BETTENAY, M . J . BICKLE, C . A . BOULTER, D . I. GROVES, P . MORANT, T . S. BLAKE, & B . A . JAMES:
Evolution of the Shaw Batholith—an Archaean Granitoid-Gneiss Dome in the Eastern Pilbara, Western Australia . .
361
R . DAVY & J . D . LEWIS:
The Geochemistry of the Mount Edgar Batholith, Pilbara Area, Western Australia
373
J . R . MUHLING:
Archaean Evolution of the High-Grade Gneiss Complex at Errabiddy, North West Yilgarn Block, Western Australia .
385
A . F . WILSON:
Oxygen-Isotope and Other Geochemical Attributes of some Archaean Granitoids from Southwestern Australia
393
ARCHAEAN MINERALIZATION
R . J . MARSTON & D . I. GROVES:
The Metallogenesis of Archaean Base-Metal Deposits in Western Australia
409
P . K . SECCOMBE & K . M . FRATER:
A Preliminary Study of Sulphur Isotopes and Ore Genesis at the Golden Grove Copper Deposit, Western Australia .
421
A . C . GRIFFIN:
Structure and Iron-Ore Deposition in the Archaean Koolyanobbing Greenstone Belt, Western Australia
429
P . C . THURSTON:
Economic Evaluation of Archaean Felsic Volcanic Rocks using REE Geochemistry
CONCEPTUAL MODELS
U . R COSTA, W . S . FYFE, R . KERRICH, & H . W . NESBITT:
Is Ocean Formation Synchronous with the First Preservation of Crust? . .
453
C . R . ANHAEUSSER & L . J . ROBB:
Magmatic Cycles and the Evolution of the Archaean Granitic Crust in the Eastern Transvaal and Swaziland
457
KENT C . CONDIE:
Geochemical and Isotopic Constraints on the Origin and Source of Archaean Granites —•
R . G . PARK:
Origin of Horizontal Structure in High-Grade Archaean Terrains
4 6 9
481
N . J . ARCHIBALD, L . F . BETTENAY, M . J . BICKLE, & D . I. GROVES:
Evolution of Archaean Crust in the Eastern Goldfields Province of the Yilgarn Block, Western Australia
491
INTRODUCTION
PROGRESS IN RESEARCH INTO THE EARLY HISTORY OF THE EARTH: A REVIEW, 1970-1980 G. J. H. McCall Consultant Geologist, 27 Venns Lane, Hereford HR1 IDE, England ABSTRACT
A review of aspects of the development of the Earth-Moon system and the early history of the Earth, citing data accrued during the past decade, is illustrated by diagrammatic charts. A preferred general model is outlined. Models involving an early interval of intense bombardment of the Earth and Moon are regarded as only one possible scenario, and the statement elsewhere by the author, of an alternative scenario, is referred to. Favoured models for the repeated greenstone-granite (sensu lato)/gneiss cycles of the Archaean allow an older cratonic foundation or infrastructure to the greenstone belts, but accept the dominance in the granite/ gneiss areas of newly generated and post-greenstone phase sialic material as intimate injections and bodily intrusions, of mantle derivation. It is suggested that we have never seen protocrustal rock material, and that the newly discovered, most ancient rocks in Greenland represent yet another, primitive, early greenstone-granite cycle. Arguments of Baer and Kroner have strengthened the author's previously expressed opinion that plate tectonics and immense oceans with deep subduction at their margins are a feature only of the last 1000 million years of geological history: Baer believes that the eclogite sinker required for deep subduction could not be formed until about 1000 Ma ago because of the constraint of higher heat flow and thermal gradient: Kroner has shown convincingly that both the greenstone belts and the later Proterozoic sedimentary depositories must have been entirely or almost entirely ensialic, oceanic crust foundation being rarely if ever achieved. Evidence for the lack of displacement of the cratonic areas on either side of the belts, both geological and palaeomagnetic, precludes the large-scale relative movements required by plate-tectonic regimes. The author favours an evolving planet rather than rigid application of uniformitarian thinking. He stresses the need for much more field-based research on the Archaean and makes a special plea for detailed sedimentological research, which could yield new data concerning the nature of, and physical controls on, the Archaean oceans, their configuration, and even the grey area of knowledge—the configuration of the Archaean globe itself.
INTRODUCTION Since 1970, the date of the previous international symposium on the Archaean in Perth, there has been an immense output of research on the Archaean. Whereas the first symposium was concerned mainly with greenstone belts and economic geology, other areas of research have broadened during the decade: even more ancient terrain than previously recorded has been discovered in Greenland and Labrador: there has been a substantial feed-back from planetological research: there has been considerable research on the origins of the binary Earth-Moon system, the atmosphere, the oceans, and life itself, attempting to fill in the gap in direct knowledge which now extends from about 4500-3900 Ma ago: and the most vexed question has been the applicability or otherwise of plate-tectonic regimes to the Archaean. Despite lack of consensus concerning the early history of the Earth, I propose to attempt to encapsulate a generalised model that I Spec. Pubis geol. Soc. Aust., 7 (1981)
see emerging from all this research. This carries forward discussions previously published (McCall, 1977): it is my intention to expand these ideas into book form. At a symposium held in February 1980 at the Royal Society, London, on the origin of the Earth's continental crust, it was stated from the floor that geological opinion had swung toward a uniformitarian view of the Archaean, one that extrapolates plate-tectonic regimes back through virtually the entire geological column. I suspect that this vieWj which I have paraphrased, is erroneous, and that the situation arose because many of today's eminent geologists are plate tectonicists who know nothing of the Archaean at first hand: and these were the people invited to speak, not geologists concerned with mapping the Archaean. I advance this review from the viewpoint of a believer in an evolving Earth. Though we never cease to marvel at glimpses of the familiar in the very ancient rocks—for example
4 G. J. H. McCALL turbidites with their distinctive primary sedimen- cant depth dimension in direct lunar rock studies, tation structures (Dunbar & McCall, 1971) or compound the frustration: likewise the difference stromatolites (Schopf, 1976: Henderson, 1977) in initial bulk chemistry and dynamic environsuch examples where the present is the key to the ment—the first is difficult to derive with any past should not blind us to situations where it is degree of accuracy because of certain assumpnot. The Earth is a small planet of a not very tions concerning the Earth and surface bias in the large star in a universe composed of myriads of case of the Moon (Smith, 1976). In addition, the such stars, all evolving in a unidirectional favoured "Baldwinian" models for the Moon manner. To suppose that it is special and behaves may be oversimple: I have discussed elsewhere differently is illogical, and has religious, not (McCall, in press) the major anomalies emerging scientific, roots. That there have been unidirec- from the study of the lunar highlands and maria, tional progressions in the history of the Earth is and also the problems of explaining the comquite apparent: there has been a fall off in the plexity of the craters on the Moon, Mars and the heat production from within, and a fall off in Martian satellites (this problem has been also exogenic radiation: there has been unidirectional remarked on by Moore, 1978): I suggested that biotic evolution: there have been changes in the the difficulties are such that they may eventually oceans both related and unrelated to biotic evolu- lead to the abandonment of Baldwinian impact tion: there have been changes in the character of models in favour of endogenic cratering models. the atmosphere: there have been changes in the I am certainly not alone in these doubts as a character of the magmas produced (komatiites, recent letter to me from an astronomer (R. W. granites sensu lato): there have been changes in Stevens) shows: he concluded . . the inexorable the patterns of mineralization (nickel and pri- accumulation of data and experimental results mary uranium early: porphyry coppers late, etc.). will eventually convince workers that the meteoriThe only reasonable conclusion is that the Earth tic theory is untenable". has repeated the same general processes in a For the present, my conclusion is that we can, broad sense, but against the background of the at best, obtain incomplete and ambiguous eviconstraints imposed by its unidirectional evolu- dence from planetology, and we are driven back tionary development: constraints which make the to unlocking the secrets of the Earth's dark period largely by study of the Earth itself. It is my processes vary in detail with time. We shall probably never achieve more than a thesis that in this we should not be applying the blurred, generalised model for the Earth's dark retrospective methodology favoured by Dickinperiod (4500-3900 Ma ago), unless rocks even son (1980), who favours extrapolation backwards older than those of West Greenland are dis- in time from plate-tectonic regimes of the covered. The search for such terrains should cer- Phanerozoic, but that we should be engaging in tainly be continued, for their discovery is by no detailed and sophisticated study of the Archaean means impossible, though the chance diminishes rocks themselves, which surely have many more secrets to unlock. with time There is a major paradox in planetology which The various stages in the development of the allows rocks more than 4000 Ma old to be ex- Earth-Moon system and of the Earth itself are posed over large areas of the lunar surface, but discussed below and this discussion is illustrated yet to be completely obliterated from the surface with the aid of diagrammatic charts (Tables I, II of our own planet. Wasserburg (1961) envisaged and III). a type of steady state, attributing the age of the oldest rocks to the mean rate at which normal BEGINNINGS OF THE EARTH-MOON geological processes erase the record. Smith SYSTEM (1976) cited prolonged weathering and meta- It seems generally agreed that the substance of morphism, continuing igneous activity and geo- the Earth and Moon condensed from the solar tectonic processes (especially plate tectonics) as nebula and was then accreted. However, the the main agencies of obliteration. TABLE I None of these agencies has been significantly Alternative models for the origin of the active on the surface of the Moon for the past Earth-Moon system 3800 Ma, and it was hoped that lunar research would feed back information to fill the gap. That hope has, however, been largely frustrated, mainly by the ubiquitous lunar regolith which presents the surface rocks as a breccia, a chaotic jumble. The incomplete stratigraphic and geophysical controls, and the total lack of a signifiv
ORIGINS OF EARTH-MOON SYSTEM A1 ternat ive Models
Condensation
(1) accretion + fission
from solar nebula
(2) volatilization + condensation
and accretion
(3) early capture + continued accretion
(k) late capture of fully grown Moon
(5) simultaneous accretion
P R O G R E S S I N R E S E A R C H , 1970-1980 TABLE II
Early history of the Earth-Moon system Note that there was no significant activity except minor impacts and tidally triggered seisms on the Moon after 2500 m.y. MOON ?Endogenic cratering
^Endogenic
Degassing Profuse Highland Cratering
Earth with
, . Large l a t e craters (Copernicus, Tycho)
Large Hare Basins
Mbon i n s t a t ion, ? I n t e n s e bombardment by p l a n e t e s i m a l s
near but outs i de Roche . . im,t
>
« ? ®
Global plagioclasic b a s a l t outer layer : 6 0 km thick
Mare basalts
Late b a s a l t s in c r a t e r s
4
~
EARTH
<jj jjj o, 0 -
? H o t t e r , l a r g e r : primordial c r u s t slower to form : u n s t a b l e : recycled as soon as convection s t a r t e d up in mantle
— «j -g ^ - Ei
?also possibly intense bombardment by p l a n e t e s i m a l s
1 DARK PERIOD
1
1
1
ARCHAEAN
origin of the Earth-Moon system with its strangely large satellite in proportion to the planet, though not to other satellites of the solar system, remains controversial. Smith (1976) listed the alternatives: fission (Ringwood, 1975a; Binder, 1975), volatilization and condensation, late capture (Anderson, 1973; Cameron, 1973): early capture and simultaneous accretion without capture (Smith, 1974). Although admitting that a case could be made for every one of these, Smith favoured early capture because of the less severe dynamic and chemical constraints: he believed that early capture, before the last stages of accretion allows ready explanation of the Moon's plagioclasic outer shell, mantle, and core. Late capture of the Moon would produce a very large heat pulse, whereas early capture and simultaneous accretion without capture (which he also found attractive) would allow Earth and Moon to have separate thermal regimes, for which there is much evidence. In a more recent paper, Smith (1979) has continued to favour early capture of a partly accreted Moon, both Earth and Moon continuing to accrete hot detritus after capture. However, he now considers that no one process of those put forward can adequately account for the Earth-Moon system as it is, and entertains the
5
possibility that the answer lies in a combination of two or more of these processes. Smith (1976) explained the chemical disparity by supposing that, though the Earth and Moon came from the same part of the Nebula, there was more efficient collection of Fe-rich material by Earth than Moon; loss of metal by the Moon; gain by the Earth; or a combination of all three. He also considered the possibility that the Moon experienced a separate high-temperature event or was more efficient in the separation of metal into its core. He concluded that the details would always be controversial, but that it was reasonable to suppose that the Earth had the Moon in station close in, but just outside the Roche limit, ca 4500 Ma ago. All the geochemical differences derived from 'guesstimates'' of the bulk composition could, he believed, be fitted into a model of early disintegrative capture of a proto-Moon, with continuing accretion while in Earth orbit, but he acknowledged the case for alternative scenarios, especially simultaneous accretion without capture. If I read his recent immense paper (Smith, 1979) correctly, he now has doubts that it can have been so simple! Two conclusions may be reached, relevant to this review. We can invoke no catastrophic influences arising from lunar capture after 4500 Ma ago and must expect normal tidal forces to have operated from then on. Secondly, the difference in bulk chemistry, particularly that involving volatile elements and the proportions of silicaand iron-seeking elements, suggests that we should not look for any close analogy in the pattern of early development of the Earth and Moon. EARLY INTENSE BOMBARDMENT The question whether there was a shared early intense bombardment remains. That the first 500800 Ma was a period of intense bombardment of Earth and Moon has been suggested by Green (1972), Goodwin (1976) and Smith (1976). Green suggested that the greenstone depositories were impact-originated basins analogous to the lunar maria. I have myself raised objections and considered that proto-ocean origin was more likely (McCall, 1977). Recognition that komatiites are not a uniquely restricted early product (see below, greenstone belts) has virtually destroyed Green's case. Smith suggested that the bombardment produced high temperatures that prevented a stable crust forming on Earth until about 3000Ma ago. Goodwin suggested that mega-impacts from 4500-3800 Ma ago dominated the early history of the Earth: and both he and Smith believe that the effects were greater on Earth than on the Moon, because of its larger size, thin crust and hotter,.
G. J. H. McGALL T A B L E III
Evolution of the Earth's crust, atmosphere and hydrosphere, and development of life 1 DARK PERIOD CRUSTAL
1
1
1
I
I
1
-PROTEROZOIC
ARCHAEAN
|
GREENSTONE EVENTS Largely basal t ie primordial crust:wi th ei ther partial or complete, global, veneer of g ran i te or acid vol-
PreAmi.tsoq I sua
Swaz iland Sebakwian Kola Ukraine Pilbara
Early belts arcuate around granite nodes
Bulawayan
S1 ave Superior E. Goldfields
FUTURE
PHANEROZOIC
GRAN ITIC AREAS were protoeratons: accessions of new granitic material (s.i.) followed greenstone events; products of earlier such cratonizations fringe and under!ie, in the form of thinned out remnants, the greenstone sequences (most are, however,,much modified)
XXXXXX
Very early rifting, instabi1i ty, recycli ng
Initial nodal cratonization
GREENSTONE BELTS were proto-oceans: not very deep (evidence of shallowwater clastics alternating with volcanics): no ophiolites: no sheeted dykes: no pelagic sediments like those in Phanerozoic trench environments: mainly ensial ic rifts, closely spaced: origin in chains of mantle plumes?
Late belts 1inear
GRANITIC EVENTS X
? Possibly virtually the entire globe eratonized*
Peak o f cratonization
Question-: did era toni c area of globe increase to a peak at end of Archaean, then continue to increase at a. lesser rate: or was there continuous recycling: or only recycling when deep subduction commenced?
Later cratonization restricted to 1 inear.,,active belts
Onset of mantle convection — numerous narrow, closely spaced cells - - - - - circular cells linear cells TECTONIC
1
EVOLUTION
Break-up of older eratonic expanses new convection patterns with larger and larger,eel Is ending in the present pattern of i mmense eel 1 s , f ew i n number
PATTERNS
Phase of greenstone belts: no large lateral Still no large lateral displacements: mainly displacements: quite small plates, jostling: platform seas on an ex"sag" subduction or no deep subduction at tensively eratonized all? -No large, deep, spreading oceans? Seas globe shallow, covering much of continents, with intervening zones of moderate depth, rifts analogous to marginal basins — decrease in thermal gradient ATMOSPHERE,
Phase of plate tectonics (s.s.) Large spread i ng oceans (Grenvi1le) ECLOGITE SINKER POSSIBLE?
deep subduction could occur
HYDROSPHERE
Volcanic degassing from solid phase in Earth's interior (very rapid) (1)? inefficient: most remained in the Earth (2) ? reducing. CH-, N H 3 , H 2 (3) Hz + C O --- u.v. C3O2 to ocean broth of life (A) All but a little H 2 of reducing gases lost to space Hydrosphere Reducing atmosphere condensed Weakly reducing, non-oxidizing atmosphere Oxidizing atmosphere as at present
Question: has atmosphere and hydrosphere been recycled: how much^recycling does occur in deep subduction (the only possible recycle mechanism):? Excess biogenii oxygen now going into atmosphere First red beds Biogenical1y produced oxygenic atmosphere": sea which previously consumed it now used up: .FeO (minority
First oceans global, shallow, with only shoallike land emergence? Greenstone rifts form deeper basins but no very large oceans, nor very deep oceans FirsY chemical _______ b a n d e d ironstones very . sediments important G;lastf<c sediments Minor carbonates Major turbidites (Fig Tree) Major marginal sediments (Moodies) First major platform sequences (Pongola, Wi twatersrand)
"Broth of life"
long period of primitive life forms
?anaerobic --- aerobic (mutation) Graph i te (life related?) I sua etc. Blue green algae or bacteria (Onverwacht, Pilbara) Stromatolites: no internal structure (Pilbara, Bui awaya n)
Mushrooming evo1uti on o f l i fe wi th hard -skel etons
First foss i l s wi th. preserved interna1 structure . Transvaal dol.omi te Stromatolites (n.i.s.) Slave
Question: possibility expanding Earth?
of an
* G l i k s o n (1980) has independently suggested global cratonization in the early to-middle Proterozoic and the attractiveness of an expanding Earth model. I do not believe that there is a great problem in initiating a new global disruptive phase, in Grenvilie times, 1100 Ma ago, particularly in the light of the pulses in the mantle recognized by Taylor & McLennan (1981); As for the pre-Grenvi1le dissipation of energy from the interior of the Earth,-the answer must, I believe, lie in finding another mechanism, not in retro-extrapolation o f plate-tectonic regimes.
PROGRESS IN RESEARCH, 1970-1980
convecting interior. Goodwin suggested a sort of boosting effect, accelerating differentiation of sial and culminating in its global emplacement in the crust about 2700 Ma ago, the boosting agent being the heat produced by the impacts. Such a boosting corollary remains highly speculative, and indeed all these essays into exogenic catastrophism remain no more than speculation, based foursquare as they are on acceptance of the Baldwinian model for the lunar surface, I favour a much less popularized alternative scenario, in which the lunar highlands are the product of global partial melting of the lunar mantle and the craters the product of the surface eruptivity which accompanied this phase, not impact scars (McCall, 1977: McCall, in press). In this alternative model, the maria are in a limited sense analogous to terrestrial oceans, and the large, complex craters are not considered to be modifications of simple impact-generated dish craters. I can only emphasise here that there is no real certainty that either the Moon or the Earth did suffer such a period of early intense bombardment, and that there is much valid evidence accumulating that is contrary to the Baldwin model. The bombardment scenario is just one possible scenario: all such scenarios can be no more than speculative, and others may well be equally acceptable. I myself am attracted by endogenic models. Global heating 4500 Ma ago accompanied by melting is, after all, a shared feature of Earth, Moon and the meteorite parent bodies. It is apparently a fundamental post-accretionary phase, but its origin remains obscure (Simonds, 1978). Short-lived radiogenic isotopes and gravity acceleration during accretion and core formation have been found to be unsatisfactory, the former because of the absence of decay products, and the search continues for a heat source. Electromagnetic inductive heating has been suggested, but there will no doubt be other suggestions. The Moon appears to have separated core and silicate shells just like the Earth, which had, apparently, formed its core by the end of the first 100 Ma (Murthy, 1976). It is difficult to attribute core, mantle and outer-shell segregation to anything but a shared fundamental process of development. I believe that the 60m-thick outer plagioclase shell of the Moon can be related to a different thermal regime affecting a smaller globe and a different bulk chemical starting point, rather than exogenic agencies, taking the place of the more complex, heterogeneous outer shell of the Earth. I visualise the Earth as larger, hotter and internally convecting, and forming a first, thin, either complete or incomplete, solid crust, probably a
7
basaltic material with patches of a thin, initial, sialic scum (though such scum may have been globe encircling). This initial crust was largely recycled after disruption piecemeal by convection currents, and only gradually, after hundreds of millions of years, was a stable crust enveloping the Earth achieved. In contrast, the Moon rapidly achieved a thick, stable crustal layer, which was shortly afterwards complicated by a number of ocean-like basins related to renewed subcrustal activity in a laterally inhomogeneous mantle: and after this it suffered only very minor eruptive and tectonic complications. I prefer such a model attributing the Earth's sialic crust to endogenic processes, but admittedly my preference is largely intuitive. I am not expert in this field, and the model is largely based on the realisation that the bulk chemistry of the Moon appears to have been such that it could not have secreted areas of sial on its surface. Differences in chemistry and thermal regime could, I suspect, adequately explain the plagioclasic outer shell by simply invoking a process of partial melting from the lunar mantle. THE ATMOSPHERE The planetesimals that accreted to form the Earth can have carried little primordial atmosphere from the nebula (evidence of inert gas spectra): either that component was lost in some way (Walker, 1976) or there never was a primordial gas envelope (Schidlowski, 1976). The terrestrial volatiles were initially concentrated in the solid phase of the planet and only later released by degassing. Rubey (1951) favoured gradual degassing, but there is now general acceptance of intense, early, rapid degassing, more or less simultaneously producing atmosphere and oceans (Fanale, 1971; Schidlowski, 1976; Walker, 1976, 1977; Shimizu, 1979). A minimum time of about 100 Ma is all that would be required to produce the entire atmosphere (Fanale, 1971; Walker, 1976). Degassing from volcanic sources has undoubtedly continued (Dymond & Hogan, 1973; Fisher, 1974), but the possibly counterbalancing effect of recycling to the mantle in deep subduction zones is controversial. Holland (1980) suggested that the calcitedolomite-calcite pattern recognised through the geological column may reflect an early (Archaean) phase of intense volcanic activity and recycling by subduction or some other return process: a middle stage of inactivity: and a last, mainly Phanerozoic, furiously active stage characterized by fully developed plate-tectonic regimes. There is certainly evidence of CO2 and N in rocks from depths of 200 km or more
8
G. J. H. McCALL
recovered from kimberlite pipes (Dawson, 1971; a, very small amount of hydrogen survived Mitchell & Crockett, 1971). However, there is no through the Archaean to make the Archaean proof of recycling and Fanale (1971) considered atmosphere weakly reducing. that the initial expulsion system was very inefficiShimizu (1979) leans heavily on new evidence ent, most of the planetary volatiles remaining from the Venusian and Martian atmospheres. We locked in the interior. The consensus seems to appear to be only at the dawn of such direct favour some approach to a steady state through- planetary research. It seems safe to predict that out the geological record, with no very significant many new data will accumulate in this way during variations. the next decades and will engender new and more Controversy also concerns the nature of the securely founded models for the Earth's initial initial outgassing atmosphere and its subsequent outgassing atmosphere. modification. The evidence of uraninite and pyrite in Archaean marine sediments remains valid—Holland (1980) pointed out that the dis- THE OCEANS AND SEDIMENTS covery of uraninite in very rapidly accumulated Siever (1977) has stressed that this remains a fluviatile sediments of the Indus is not relevant— very grey area of knowledge. There is evidence and the change to a biogenically produced oxy- that the character of the sea water was not unlike genic atmosphere is generally accepted as having that of the present oceans, and Armstrong (1980) coincided with the first red beds about 2000 Ma has suggested that the calculation of Wise (1974) ago (Cloud, 1973). The delay in its appearance is that the total volume of the oceans was estabwidely attributed to the need to sweep up all the lished at least by the beginning of the Proteroferrous iron from the oceans before the excess zoic, can be extended to the Archaean. Accepting could build up in the atmosphere. Kimberley & early, rapid, more-or-less complete outgassing, Dimroth (1976) differ from most authorities in the total volume of the oceans should have been believing that the Archaean atmosphere can be achieved early, but there may be a problem in modelled in actualistic terms, though they accept finding space to accommodate this great volume that there may have been an initial reducing of sea water if, as is probable (see below, Plate atmosphere before 3800Ma ago. Those who Tectonics and the Archaean), there were no favour the late appearance of oxygen disagree counterparts of the immense Phanerozoic oceans about the nature of the initial atmosphere. in the Archaean world. The problem is not eased Shimizu (1979) and Walker (1976) favour the loss by the numerous greenstone depositories, which of an initial, strongly reducing atmosphere to appear to have been narrow and not very deep. space: an atmosphere of hydrogen, methane and Recent suggestions that the volcanic material of ammonia would, Walker suggested, have had a the greenstone belts may never quite have broken short life and violent end. Shimizu favours an through to the surface (a sort of subcretion initial hydrogen atmosphere that was largely lost, model) seem to ignore this problem, but it is, in leaving behind CO and N at 1-2 atmospheres any case, difficult to see how such processes pressure: the CO being then converted to C 0 could produce basic rock suites of pillow lavas and its polymorphs by ultraviolet radiation, these associated intimately with sediments, both clearly going into the ocean and being converted in part formed at the sea bottom. Such suites go right to a broth of organic compounds, the 'broth of through the greenstone sequences. A more real life". Walker (1976) observed that the very fact possibility is that quite shallow seas enveloped the that the atmosphere must have come from the entire globe, with only small shoal-like land areas solid phase in the Earth's interior by degassing emergent. This model is favoured by some Rusrules out an extraterrestrial origin for life, sian authorities (McCall, 1977, p. 249). Another because no life form, however primitive, could possibility is that the Armstrong calculation is survive such a process. Schidlowski (1976) incorrect and there was a smaller volume of sea favours an initial highly reducing atmosphere water—though that is unlikely. rather than the very weakly reducing, nonThere have unquestionably been major changes oxidising atmosphere preferred by these two in the chemistry of the oceans. Holland (1976, authorities: he believes that the first organisms 1980) has noted the changing pattern of calcite utilised anaerobic fermentation, independence of and dolomite formation with time, and believes such a process being only later attained by muta- that endogenic decrease in the energy level and tions. Early release of oxygen would, he believes, exogenic changes (mainly in the level of solar have endangered the survival of life, and it is radiation) have significantly changed the because of the delayed release to the atmosphere character of the oceans. The evolution of biota caused by the oxidation of the ferrous iron in the must also have produced major progressive sea that life survived at all. Walker suggested that changes. The long history of chemical sedimenta3
4
2
PROGRESS IN RESEARCH, 1970-1980
tion involving banded ironstone production up to 2000 Ma ago must also have progressively altered the character of sea water. Unfortunately there is little direct evidence. Dolomite in Archaean rocks, and gypsum and halite in rocks as old as 2000 Ma indicate hypersaline regimes according to Cloud (1968). Evidence of evaporites at North Pole, Pilbara, given at this symposium, suggests that the sea water there 3500 Ma ago was hypersaline (Dunlop & Buick, 1981). Glauconite can only be used to trace diagenesis involving K-rich solutions back to the mid-Proterozoic. The presence of siderite, sulphide and bedded chert requires no great change in pH value for ancient seas. Stromatolites in Archaean rocks (Awramik, 1977), including some more than 3000 million years old in the Bulawayan and others about 2750 million years old in the Slave Province, should indicate similarity of sea-water character to the present day, and the precipitation of organic carbonate (aragonite) (Kimberley & Dimroth, 1976), but the lack of preservation of internal structure makes such inference suspect. The morphology of the proto-oceans remains obscure, as does the morphology of the entire Archaean world. Cloud (1968) could find nothing to support the existence of variable and separated oceans of intermediate depth in the Archaean and believed that there was a single interconnected ocean. Archaean sediments carry a profusion of primary sedimentation structures like their Phanerozoic counterparts, essentially of familiar character and scale. Sedimentological papers dealing with such sediments (for example Donaldson & Jackson, 1965; McCall et al.y 1970; Condie et aL, 1970; Henderson, 1975; Dunbar & McCall, 1971; Reimer, 1972; Eriksson, 1979) do not indicate different sedimentational processes or physical controls in ancient proto-oceans. The last-named author recognised macro- and microtidal regimes similar to those in the present-day seas. Marginal marine facies appear to be preserved only rarely in Archaean sedimentary sequences—the Moodies Group described by Eriksson is the most significant exception—and this contrast to Phanerozoic sequences may indicate an almost complete absence of wide shelves. I would like to couple my call for more detailed studies of Archaean rocks in the field with a plea for much more detailed sedimentological study. My recent work covering the entire Makran ranges of Iran has strengthened my belief that such studies offer a unique chance to unlock the secrets of the Archaean proto-oceans. We should endeavour to find such things as crystal pseudomorphs which establish the saline character of the sea, and we should record every primary sedimentational structure to establish a model which
9
encompasses the nature, morphology and ruling influences of these early seas. We know from the Greenland evidence (see below) that the seas existed 3800 (possibly 3900) million years ago—the higher figure may reasonably be inferred from inclusions in the gneiss. Chemical sedimentation was even then producing banded ironstones: clastic shale, quartzite and conglomerate were being deposited, and there was minor carbonate precipitation: graphitic shales suggest a primitive biota. Pillow lavas indicate interaction of basalt with sea water just as at present. Such direct geological evidence is irrefutable and unambiguous, and is the sort of evidence that I would like to see multiplied by intensive sedimentological study directed at the Archaean.
THE SIGNIFICANCE OF THE OLDEST ROCKS This subject was concisely summarised by Moorbath (1975) from a geochronologist's viewpoint, and Myers (1976) from a field geologist's viewpoint. The discovery of rocks older than hitherto recorded by McGregor (1973) must be acknowledged as the major advance of the decade in Archaean geology. First discovered at Godthaab and Isua in West Greenland, similar suites are now known from the Nain region of Labrador (Hurst et aL, 1977). Other reported developments of such ancient rocks remain unconfirmed: Hurst et aL (op. tit.) question the validity of the 3800 million year Rb-Sr isochron obtained by Goldich and co-workers for the Morton and Montevideo gneisses of Minnesota. Rb-Sr whole-rock ages for the West Greenland gneisses average about 3750 Ma (Moorbath et aL, 1972) and are supported by Pb-Pb whole-rock ages (Black et a/., 1971) and U-Pb zircon dating (Baadsgaard, 1973). The events in Greenland have been summarized by Moorbath (1975, Table I) and by Myers (1976). The age may either represent an igneous precursor to the Amitsoq gneiss or the metamorphism to amphibolite facies. The gneisses are remarkably ordinary in character, mainly quartzfeldspar gneisses, and are cut by a dyke suite emplaced shortly afterwards, the Ameralik dyke suite, which is the oldest known tensional dyke swarm. Moorbath (1975), I believe correctly, discounts suggestions that this gneiss terrain is a remnant of the Earth's protocrust. There appear to be even older rocks as rafts and enclaves within the gneiss (McGregor & Bridgwater, 1973), the "pre-Amltsoqs", which consist of layered amphibolite; striped iron-rich clinopyroxene and hornblende rocks; quartz-rich rocks with gruner-
10
G. J. H. McCALL
ite, orthopyroxene and garnet; basic hornblendebiotite rocks; and pods of enstatite-olivine ultrabasics, together with quartz-rich biotite gneisses containing garnet, plagioclase, grunerite and graphite. This metamorphically homogenized suite appears to be a sedimentary/volcanic/ plutonic suite derived from an even more ancient supracrustal greenstone belt, in which iron-rich sediments of banded ironstone character were prominent. There has been argument concerning the synclinal sequence 2-3 km thick at Isua, 150 km north of Godthaab. Metamorphosed greenstone supracrustals include quartzite, slate, calcareous rocks, conglomerate, basic volcanics, ultramafic pods, garnet-chlorite schist and quartz-gruneritechlorite banded ironstone associated with carbonate lenses. However, a Pb-Pb isochron age (metamorphic) was obtained on the ironstone, of 3760 ± 70Ma (Moorbath et al., 1973), and field evidence suggests that the igneous precursors of the Amitsoq gneiss invaded this sequence which therefore must be of pre-Amitsoq age. The pre-Amitsoq of Godthaab and Isua thus seem to provide firm evidence of both clastic and chemical sedimentation 3760Ma ago or earlier. The gneisses have been shown by Myers (1978) to represent acid intrusions slightly younger than the greenstone rocks, tectonically modified and interleaved with the older supracrustals. Appel (1980) has shown that the banded ironstones of Isua include the three fades characteristic of the later banded ironstones, and has inferred from the field relationships and geochemistry a submarine exhalative origin from brines of basaltic association. Conglomerates at Isua contain numerous granite clasts (Bridgwater & McGregor, 1974) of original acid volcanic provenance, and there is some evidence that the conglomerate may overlie an older granitic basement. All this evidence suggests that we can now extend the familiar cycles of Archaean greenstone accumulation and a slightly later invasion of new, mantle-derived granitic (sensu lato) material back to 3800 Ma ago, and possibly by inference to 3900 Ma. The patterns of proto-oceans and land masses supplying clastic detritus to them, of erosion and eruptivity of types broadly tamiliar in the later geological record, were already initiated, even then. Yet these were almost certainly very primitive proto-oceans and land masses: the latter may have been low and shoal-like, and were certainly composed of absolutely bare rock for there was no vegetation to bind soils or produce humic acid. The seas were probably much shallower than the present-day oceans, perhaps more like the Black Sea, but they were probably interconnected and globe encircling. The point I am
making is that this archaic world would appear by no means familiar to our eyes if we could journey back there by time machine, despite the familiarity of the rock types. It is necessary to extract every morsel of data from this unique region, which is characterized by near-perfect exposures. Even so, it is danger* ous to suppose that it was the only type of geological association at the time, and to develop from it alone too sweeping conclusions and apply them to the entire globe. Such a very restricted remnant might well be unrepresentative. Future research on this unique remnant should utilise the whole range of geological tools, for geochrono* logy and geochemistry can never alone provide the answers. The discovery rendered obsolete the equation proposed between komatiites and the Earth's protocrust. Neither komatiites nor the ancient gneisses of Greenland can conceivably be such & protocrust, and we may never see rocks of the Earth's protocrust directly. We now know that cratonizatioii had commenced 3750 Ma ago, though it may well have still been localized and inefficient, the cratons separating very primitive proto-oceans being small, thin, fragmentary and unstable. THE GREENSTONE BELTS Greenstone belts remain the subject of debate, which is best approached by enumerating the known facts: (1) The bulk of the pillow lavas closely resemble oceanic Krpoor tholeiites. (2) Komatiites are subordinate. Suefy primitive basalts have been identified at repeated intervals in both Western Australian and Onverwacht sequences (McCall, 1977, p. 60; Williams & Furnell, 1979), and do not have an exclusively basal position. Excellent surface sections at Munro Township, Ontario (Pyke et al., 1973) have displayed unequivocally the character of the flows. They have been recorded from the Archaean of the Kola Peninsula (Suslova, 1978) and Madras (Viswanathan & Sankaran, 1973). It is clear that komatiite is not restricted to the Archaean (McCall, 1973; Gale, 1973; Schau, 1975), but has minor occurrence in the Proterozoic and Palaeozoic. Komatiite does not appear to be generated by meteorite impacts (Green, 1972; McCall, 1977). (3) Calc-alkaline suites abound in the Archaean of Canada and are rather rare in the Archaean of Western Australia (Hallberg, Carter & West, 1976; Hallberg et al., 1976). They are not geochemically exactly
PROGRESS IN RESEARCH, 1970-1980
like analogous Phanerozoic suites—most of the rocks studied by Pearce et al. (1977) do not fit into the orogenic field. Opinion has been moving away from direct analogies with Phanerozoic plate-tectonic regimes (Goodwin, 1977; Archibald et al., 1978). Pearce et al. (1977) reject direct analogy with island-arc models and doubt if such models can be applied to the Archaean. Baer et al. (1978) note the general loss of support for such analogies. (4) Greenstone belts appear to be largely ensialic, the enclosing granitic rocks (sensu lata) not having been subsequently displaced to any degree (Kroner, 1977, 1979). The greenstone belts appear to pass down into zones of progressively increasing granitic rock content until they are engulfed, being preserved at the most as wisplike remnants (Goodwin, 1977). (5) The neighbouring granite areas appear to be a primary infrastructure to the greenstone belts, in a sense, as Goodwin (1977) suggested, but the relationship is complex. They and the belts themselves have been the sites of post-greenstone generation of new acid magmatic material from the mantle, and this emplacement, by processes of intimate injection, bodily intrusion and diapirism, has commonly largely obscured the original foundation to the greenstone belts, which was generally thinned-out sial. This foundation commonly has obscure boundaries and can only be detected by agedating survivals. (6) Kroner (1977, 1979) suggests that true oceanic basins without sialic foundation may have been rare or absent in the Archaean greenstone belts. (7) Though there are a few areas where basal unconformities are preserved (Glikson, 1976, p. 265; Henderson, 1977; Rutland, 1976), some recognitions of basal unconformities are unconvincing (Marston, 1978; Piatt et al., 1978) and there is a surprising lack of internal unconformity within the greenstone sequences compared with Phanerozoic sequences—a contrast particularly strange in view of the probable prolonged development, possibly over as much as 250 Ma, of individual greenstone belts. (8) There is a complete lack of ophiolites in greenstone suites: instead there are numerous layered intrusions which appear to be of sill form, though immense ponded flows could also be represented. Concordance with beds above and below is commonly observed in the field, and sedimen-
11
tary screens also occur within the layered sequences. There are no sheeted dyke complexes at all, and they should be preserved if pillows are preserved. The lack of ophiolites cannot as Windley recently suggested (comment from the chair at the recent Royal Society symposium) be explained by supposing that geologists simply do not know how to recognise ophiolites in deep crustal sections, for the layered plutonic rocks are commonly associated with pillow lavas and sediments which preserve their primary structures, and there is no conceivable process that could preferentially transform and mask the character of the plutonic rocks. (9) The plutonics are closely associated with banded ironstone, dark shale and coarser clastic sediments including conglomerate. Pelagic sediments of deep-ocean character are not recognized. Limestones of any sort are rare. The implication is that the greenstone depositories were not very deep, though they must have subsided to accommodate the very thick sequences recorded. Such great thicknesses are undoubtedly real, supported by consistent facings and continuity up-strike: there can be no possibility that there has been duplication to any significant extent by thrusts as suggested by Burke et al. (1976) and Gorman et al. (1978). (10) There is little progressive increase in metamorphic grade with depth of burial in these immensely thick sequences, which commonly display greenschist metamorphism (though patches of higher-grade metamorphism can be found). (11) There are no blueschists in Archaean greenstone belts: however these indicators of fossil subduction zones (Coleman, 1972) would not be expected in view of the increased thermal gradient (Fyfe, 1973). (12) The belts have the general character of numerous, not very wide, rifted depositories, founded in sial: in this they are not dissimilar to modern intracontinental rifts, but Goodwin (1977) has suggested analogy with quite small back-arc or inter-arc basins which commence with rifting. He relates them to Assuring of existing sialic crust and the action of narrow convection cells. The analogy seems to be reasonable, but it must be borne in mind that these are all that appear to have existed as proto-oceans in the Archaean, there being apparently no complimentary deep oceans of Atlantic or Pacific type. The numerous, not widely
12
G. J. H. McCALL
spaced greenstone belts suggest a series of quite narrow depositories, rather than wide seas. (13) The belts now have synclinal form which is clear from the facings. The earlier belts appear to have been arcuate and to have intervened between sialic nodes. Many had the form of triangular synclinoria and their arcuate character appears to be primary, not superimposed. Later belts were linear and more extensive. Goodwin (1977) recognized a number of individual basins composing a greenstone belt, and suggested that each such basin stems from a mantle plume in a linear array of such plumes. (14) Some belts display a late, thick development of turbiditic sediments, and at least one is in turn capped by a thick sequence of neritic sediments. (15) Goodwin recognised three intervals of intense volcanic, plutonic and tectonic activity (3800-3500; 3100-2900; 2700-2600 Ma), but it appears to be more logical to relate the rocks of Greenland to a fourth, oldest-known cycle. (16) The extension of greenstone terminology to West Greenland appears to be justified, but Gill & Bridgwater (1979) have observed that there is no exact match with the betterpreserved, later Archaean greenstone belts, nor is there any basis for deducing the geotectonic setting. (17) Any acceptable model for the greenstone belts and granite areas of the Archaean must take into account the incorporation of belts of older cycles, preserved now elsewhere, into a stable cratonic nucleus while the later belt was forming. There appears to be no other way that this could be achieved than by some form of progressive outward growth of cratonic nuclei, but this was probably not a simple pattern because of complications by the progressive change with time from nodal and arcuate (nuclear) to linear fissuring patterns. THE GRANITIC ROCKS (SENSU LATO) The areas which separate greenstone belts include granitic intrusions (s./.), developments of gneissic rocks, granulites and relics of supracrustal rocks, which may be no more than wisps (as in the granite around Johannesburg in the Kaapvaal craton). The following points can be made: (1) The boundaries with greenstone belts are commonly complex and confused: they seem to have involved considerable fault-
ing: higher grades of metamorphism than are common in the greenstone belts themselves are recorded at the peripheries of the belts, close to the granitic areas. The granite areas are sites of higher-grade metamorphism, probably from a thermal blanketing effect. (2) The principal component of these areas is a new accession of sialic material from the mantle by partial melting, such material occurring as distinct plutons and intimate injections, which on subsequent tectonic modification form banded gneiss (Myers, 1978). Hunter et al. (1978) have clearly demonstrated the abundance of mantlederived material and discounted any connection with island arcs and convergent plate margins. The banded gneisses are partly composed of pre-existing supracrustal material, partly of newly introduced mantle material, and may even incorporate material from the older sialic foundation to the greenstone belt which supplied the supracrustal enclosures. (3) The ancient sialic foundation that formed the cratonic infrastructure to the ensialic greenstone belt commonly has no welldefined geological boundaries, its existence being established only by reliable older dates from gneissic rocks, etc: for example certain Katarchaean dates obtained from the western Yilgarn Block could indicate such relics (Arriens, 1971). (4) Clearly there were areas in the pre-existing cratonic nuclei which escaped almost completely the effects of later uprise of granitic material (sensu lato) in the next cycle, i.e. modification and obscuration—otherwise the perfectly preserved Archaean rocks of Barberton and the Pilbara, more than 3000 Ma old, could not still exist. Contacts between younger greenstone belts and such older relics are rarely preserved, and if preserved are commonly obscured by tectonization or other processes. In a few areas basal unconformities are recognized, and the best examples are in, the Slave Province (Henderson, 1977, p. 49) and in Rhodesia, between the Sebakwian and Bulawayan (Glikson, 1976, p. 265). There is a surprising lack, of unconformity in the granitic areas as in the greenstone sequences. (5) Granulites are common in the granitic areas, and must have originated as supracrustal rocks for they contain indisputable sediments: thus some process of deep metamorphic dehydration must have taken place
PROGRESS IN RESEARCH, 1970-1980
after they were deeply buried in the crust, the granuhtes being the dry residue left behind by removal of granitic melts and their emplacement higher in the crust (Fyfe, 1973). (6) Anorthosite is also common in such areas: it is manifestly derived from layered intrusions which include ultrabasics, gabbro, leucogabbro, anorthosite and chromitite; anorthosite is a subordinate component of such suites which commonly preserve cumulate textures and reveal igneous differentiation trends (Windley, 1973; Hor et al., 1975; Windley & Selvan, 1975; Windley et at., 1979). The occurrences in West Greenland, the Limpopo Belt and Madras all conform to this model. There has been very strong tectonic modification, and such rocks are now strung out as extensive layers and lenses within deformed, commonly tonalitic gneisses, which volumetrically overshadow them. There is no similarity to ophiolite sequences except in the shared layered character. The origin of the anorthositic magma is related to the conversion of calcic pyroxene to hornblende in a wet magma and is not the same as the origin of lunar anorthosite, in which the crystallisation of calcic plagioclase is related to low concentration of volatile elements. (7) Tonalitic and trondhjemitic rocks generally predominate in these cratonic nuclei, but granodiorite and granite (sensu stricto) are also represented. There seems to be considerable variation in the composition and geochemistry, of the intrusive and gneissic suites of different regions (for example Greenland, Swaziland, Minnesota) and it may be unwise to attempt to derive globally applicable models from the study of any one region. There is a marked difference in the character and geochemistry of the Archaean suites to Phanerozoic batholithic suites, perhaps at least partly because of progressive change in the character of mantle partial melting products with time. (8) There does not seem to be any special quality of "grey gneisses" that restricts them to the early Archaean: such gneisses do not seem to be peculiarly archaic within the Archaean; new grey gneiss can apparently form in each cycle, though it can probably be reworked in a later cycle without much modification. What we see is cycles repeating themselves backward to the point of obscurity, and Hutton's dictum remains true; we see no vestige of a begin-
13
ning, of a special protocrust represented by a peculiar rock type. (9) New granitic material (s.l.) derived from the mantle appears to have been produced in cycles throughout the early Archaean, but there is evidence of peak production 3000-2750 Ma ago and resultant widespread cratonization, possibly on a globe-en veloping scale. This left behind a largely cratonized world to be affected by only ensialic development of mobile belts during the long span of the early and middle Proterozoic, with extreme activity and disruption of cratonic areas again at the end of the era. (10) The question whether there was a steady increase in sial globally through the geological record has been discussed by Watson (1976); alternatives suggested are a general increase with the peak mentioned above, or a steady state with compensation by recycling of sial in deep subduction zones (Armstrong, 1980). The subject remains controversial. PLATE TECTONICS AND THE ARCHAEAN The recent symposium on the origin of the Earth's continental crust (Royal Society London, Feb. 21-22, 1980) was in the nature of a platform for actualistic extrapolation of plate-tectonic regimes, familiar in the Phanerozoic, back to the Archaean. The tenor of comment from the floor suggested to me that many of the audience found the statements unconvincing. At a recent symposium in Canada (Baer et ah, 1978) there was a consensus that subduction at convergent plate margins could not provide the answer to the problems of the Archaean. Besides myself (McCall, 1977), Kroner (1977/1979) and Baer (1977) have argued convincingly in this vein and I find their arguments well founded. Baer suggested that deep subduction was not possible until about 1000 Ma ago because eclogite was not stable in the oceanic crust until then. He accepted the model of Ringwood & Green (1966) and Ringwood (1975&), and the need for an eclogite sinker to initiate deep subduction, and argued that the raised geothermal gradient must have prevented eclogite formation, placing constraints on the evolution of the Precambrian crust. He regarded plate tectonics as a necessary but transitory stage in the history of the Earth. The only possible actualistic margin of Proterozoic age older than the Grenville is the Coronation Geosyncline of Canada (Hoffman, 1973; Olade, 1975), and there is no real evidence for any Wilson cycle or collision orogeny in North America prior to the Gren-
14
G. J. H. McCALL
ville event ca 1100Ma ago (Baer, 1974, 1976). Neither the Zambezi nor the Kibara belt is a product of plate accretion (Shackleton, 1973), nor is the Namaqua belt (Kroner, 1974). Baer's model requires thickening of the lithosphere with time to its present thickness by progressive deepening of the mantle below the crust, the crust: mantle ratio of continental lithosphere decreasing with time. He accepts convection in the Precambrian (Verhoogen, 1973) and in some areas the top of the asthenosphere was colder and the lithosphere above thicker, producing a lower surface relief of about 10km. Pull exerted by asthenospheric currents would have thus varied and caused relative displacement of lithospheric blocks; and where the top of the asthenosphere was colder there would have been decreased viscosity and better coupling with the lithosphere. In such a system, Baer suggested, full instability was not achieved until about 1000 Ma ago, when the possibility of an eclogite sinker initiated deep subduction. The Archaean was dominated by rifting regimes, but the rifts did not accumulate great volumes of anhydrous gabbro beneath a thermal blanket of overlying sediments: nor was there temperature rise, transformation of gabbro to eclogite, deep subduction of ocean crust, or development of Atlantic and Pacific type ocean margins. Kroner (1977, 1979) based his arguments mainly on the overwhelming evidence, both geological and palaeomagnetic (the latter is valid at least for the Proterozoic: Dunlop, 1980), that Archaean and Proterozoic regimes involved little relative displacement of the blocks bordering greenstone belts or the Proterozoic mobile belts. His arguments seem to establish that large-scale displacements of plates laterally, the requirement of plate tectonics, was impossible during both Archaean and early-middle Proterozoic. He related the late development of plate-tectonic regimes to an increased rigidity following lithospheric thickening. Arguments concerning the applicability of horizontal tectonics (i.e. large-scale thrusts, nappes) to the Archaean have been pursued for many years. Stowe (1968) recognized such structures in the Rhodesian craton, but it is the consensus of geologists familiar with the South African, Canadian and Australian Archaean that such styles are exceptional and that the Archaean is dominated by vertical tectonics. There have been objections to this by Burke et al. (1976), but the uniqueness of the dominant Archaean tectonic style seems to be well documented, though there is disagreement concerning the controlling effect exerted by diapirism or otherwise. The most important subject of such controversy has been the Archaean of West Greenland, and
Chadwick et al. (1974) have concisely summarized the conflict between: (1) A sliced-up, interleaved system, brought together from a diversity of geotectonic settings by lateral displacements on low-angle surfaces (Bridgwater et al., 1974) and (2) A cognate association of basement and supracrustal rocks analogous to the younger Archaean greenstone and granitic rocks associations (which Chadwick et al., 1974, favour). The first model allows plate-tectonic analogies full play, whereas the second is more consistent with the vertical tectonics recognised in other Archaean greenstone-granite terrains. In the second model, any horizontal displacement did not bring rock masses of diverse provenance together early on, but was related to the much later emplacement of the Nuk gneisses 3000 Ma ago. The least that one can say is that there is no certainty that the first model is correct, and consequently this region cannot be considered to establish the operation of major lateral displacements and plate-tectonic regimes in the early Archaean. I conclude that the giant oceans and deep subduction at their margins, accompanying the immense lateral plate displacements that are a feature of Phanerozoic plate-tectonic regimes, were probably restricted to the last 1000 Ma or so of the Earth's history. I believe that there has been far too much invocation of vast oceans with fully developed mid-ocean ridges in studies of Phanerozoic ophiolites, and that most of these relate to quite small marginal basins, not fully fledged oceans. Archaean proto-oceans appear to have been related to narrow convection cells, and calc-alkaline suites in greenstone belts may be the markers of downgoing convection in such cells and recycling of basalt through the mantle. If so, only an up-and-down circulation seems acceptable. The alternative is that such rock suites originated in some other way—the origin of calcalkaline suites at Phanerozoic plate margins is by no means clear and even there more than one mode of generation of calc-alkaline suites may have operated. ARCHAEAN LIFE Schopf (1977) listed the various types of life form that have been described from the Archaean, but considered few, if any, to be wholly compelling. These supposed fossils are mainly micro-organisms that may be blue-green algae or bacteria (Nagy & Nagy, 1969): there are four reported occurrences of stromatolites
PROGRESS IN RESEARCH, 1970-1980
(Awramik, 1977). (It appears now that the Bulawayan stromatolites of A. M. McGregor are of the same age as those of the Slave Province: the oldest stromatolites, more than 3500Ma old, are those of the Warrawoona Group at North Pole in the Pilbara reported on in this symposium, see Walter, 1980; and Dunlop & Buick, 1981.) The ambiguity of such occurrences is stressed by Cloud & Morrison (1979), who are impressed by the real possibility of late microbial contamination. The lack of preserved internal structure is frustrating, and it is generally agreed that the oldest life-forms that are unequivocal because of preservation of internal structure are contained in the 2300 Ma old Transvaal Dolomite (Nagy et al., 1977). There is much indirect evidence for a prolonged development of primitive life forms through the Archaean. Nagy et al. (1977) consider that graphite schists from West Greenland imply that life was in existence 3800 Ma ago, but they admit that there is no proof. Cloud (1974) suggested that the oxygen required for the oxidation of the iron minerals in Precambrian banded ironstones, which date back to 3750 Ma ago, may well have had a biogenic origin. Carbon- and oxygen-isotope evidence (Oehler et al., 1972; Eichmann & Schidlowski, 1975; Schidlowski et al., 1975) favours the existence of Archaean life. Reimer et al. (1979) regard the total amount of carbon in the Sheba Formation of the Fig Tree
15
Group of Swaziland (0.42% of the total rock mass) as indicating microbial photosynthesis at much the same rate in the proto-oceans as in modern oceans. This fascinating field of research could yield new and exciting discoveries because such occurrences are hidden in the Archaean rocks like the proverbial needle in a haystack. Even the most confirmed hard-rock enthusiast working on the Archaean must be on the look out for traces of ancient life. CONCLUSION A review paper of this kind is not presented with the aim of presenting the final answers to the problems of the Archaean. The aim is rather to stimulate thinking about the Archaean by introducing some fresh ideas and views. There is much that I have omitted because of limitations of space and time. However, what I have considered has stemmed from a most exhaustive research into the literature, and I can only stress the enormous volume of literature now available dealing with problems of Archaean geology. I conclude with an acknowledgment to Professor W. S. Pitcher and Professor D. Flinn of Liverpool University who have generously afforded to me the status and facilities of a Fellow of the University: without that, continuance of this research would not have been possible.
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covery of a komatiite in the Precambrian of India and its significance in the nature of Archaean volcanism and of the early crust in the Indian Shield. Curr.
Sci.,
42, 2 6 6 - 2 6 9 .
WALKER, J. C. G., 1976: Implications for atmospheric evolution of the inhomogeneous accretion model of the origin of the Earth; in Windley, B. F. (Ed.) The Early History of the Earth, 537-546. Wiley, London. 1977: Origin of the atmosphere: history of the release of volatiles from th£ solid Earth; in Ponnamperuma, C. (Ed.) Chemical Evolution of the Early Precambrian, 1-11. Academic Press, New York. WALTER, M. R., 1980: Paleobiology of Archaean stromatolites. Second internat. Archaean Symposium, Perth, 1980, Extend. Abstr., 22-23. WASSERBURG, G. J., 1961: Crustal history and the Precambrian time scale. Ann. N. Y. Acad. Sci., 91, 583-594.
WATSON, J., 1976: The Earth's crust in Precambrian times. Proc. Yorks. geol. Soc., 41, 145-162. WILLIAMS, D . A . C . , & FURNELL, R . G . , 1979: R e a s s e s s -
ment of part of the Barberton type area, South Africa. Precamb. Res., 9, 325-347. WINDLEY, B. F., 1973: Archaean anorthosites: a review with the Fiskenaesset Complex, West Greenland as a model for interpretation. Spec. Pubis geol. Soc. S. Afr.,
3, 3 1 9 - 3 3 2 .
WINDLEY, B . F . , & SELVAN, T . A . , 1 9 7 5 : A n o r t h o s i t e s
and associated rocks of Tamil Nadu, Southern India. J. geol. Soc. India, 16, 209-215. WINDLEY, B . F . , BISHOP, F . C . , SMITH, J . V . , STEELE, I. M . , NEWTON, R . C . , DELANEY, J . S . , & MCCOR-
MICK, G. R., 1979: Anorthosite complexes in the early crust of the Earth: comparison of mineralogy with lunar anorthosites. Proc. Tenth lunar and planet. Sci. Conf, Houston, Texas, 1356-1358. WISE, D. U., 1974: Continental margins, freeboard and" the volumes of the continents and oceans through time; in Burk, C. A., & Drake, C. L. (Eds) The Geology of Continental Margins, 45-58. SpringerVerlag, Berlin.
REGIONAL REVIEWS OF ARCHAEAN TERRAINS
THE PATTERN OF ARCHAEAN CRUSTAL EVOLUTION IN SOUTHERN AFRICA AS DEDUCED FROM THE EVOLUTION OF THE LIMPOPO MOBILE BELT AND THE BARBERTON GRANITE-GREENSTONE TERRAIN* J. M. Barton Jr.
Bernard Price Institute of Geophysical Research, University of the Witwatersrand, 1 Jan Smuts Avenue, Johannesburg 2001, South Africa ABSTRACT Evidence relating to the tectonic evolution of the Limpopo Mobile Belt and the Barberton Granite-Greenstone Terrain is reviewed and models are proposed to explain the evolution in each area. The Limpopo Mobile Belt is believed to have formed -3570 Ma ago as a graben or aulacogen in a >3790 Ma old protocraton. The aulacogen was deformed in three separate events by Riedal shear between the adjacent cratonic blocks and a fourth time by a clockwise rotation of the Rhodesian Craton with respect to the Kaapvaal Craton —2600Ma ago. It is postulated that the Barberton Granite-Greenstone Terrain originated - 3550 Ma ago as a back-arc basin between two cratonic blocks: the ^3790Ma old protocraton in which the Limpopo Mobile Belt formed to the north and a second block to the south on which an island arc was superimposed. The back-arc basin formed and was destroyed by compression and plutonism within —60 Ma but episodic plutonism continued until - 2550 Ma ago. The plutonic rocks in the Barberton Granite-Greenstone Terrain have genetic affinities to oceanic crust. While the Limpopo Mobile Belt and the Barberton Granite-Greenstone Terrain appear to have evolved in different manners during the same time interval, plate tectonics can provide a cause-and-effect link between them.
INTRODUCTION Understanding of the tectonic evolution of crustal rocks of southern Africa during the Archaean has been hampered by a lack of multidisciplinary study aimed at resolving comprehensive patterns. Most areas containing rocks of this age have been mapped only on a reconnaissance scale and, until recently, in only a very few, such as the greenstone terrain of the Barberton Mountain Land in South Africa and Swaziland and some of the granite-greenstone terrains in Rhodesia (Figs 1, 2 & 3), have any attempts been made to combine detailed mapping with structural, chemical, petrologic and geophysical analyses. Even in these areas, a comprehensive geochronologic framework with which to view and compare tectonic evolution has been lacking. As a consequence, there has been a tendency to view events occurring during the Archaean in terms of broad generalities and to ascribe evolution to non-uniformitarian processes covering inordinately long time spans. However, two areas, the intracratonic Limpopo Mobile Belt (Fig. 2) and the granitegreenstone terrain surrounding and including the
Barberton Mountain Land (Fig. 3) have recently come under intense study, largely through the auspices of the South African Government's contribution to the International Geodynamics Project. In the Limpopo Mobile Belt and the gneiss terrain surrounding and within the greenstones of the Barberton Mountain Land, geologists (see Acknowledgments) with many different specialities have combined to try to understand the tectonic evolution of these areas. Furthermore, detailed isotopic analyses have made it possible to compare the timing of tectonic events in the two areas. In this paper, the most significant results of these studies are used in an attempt to resolve the pattern of tectonic activity in each area. These two contrasting patterns are then rationalized and a rriodel presented for Archaean crustal evolution in southern Africa. The data described in this paper come from the work of many individuals and are acknowledged, but the synthesis is the responsibility of the author, and does not represent a consensus of opinion. This paper is not a definitive study of the tectonic evolution of southern African crust during the Archaean, but is aimed at stimulating further thought and study.
* Publication number 59 of the South African contribution to the International Geodynamics Project.
Spec. Pubis geol. Soc. Aust., 7 (1981)
22
Fig. 1.
J. M. BARTON JR
A geologic map of a portion of southern Africa showing the areas where rocks of Archaean age are exposed. The diagonal pattern represents younger than ~ 2800 Ma old cover rocks. R = Rhodesian Craton. P = Palabora. K = Kaapvaal Craton. J = Johannesburg Dome. V = Vredefort Dome.
THE LIMPOPO MOBILE BELT The synthesis of the tectonic evolution of the Limpopo Mobile Belt presented here is taken from the review of Barton & Key (1981). Consult that paper for references to source materials. The Limpopo Mobile Belt (Figs 1 & 2) is a zone of high-grade, often granulite-facies, polymetamorphic rocks between the Rhodesian and Kaapvaal Cratons. It consists of a fault-bounded Central Zone containing an intracratonic sequence of sedimentary rocks, and, in the east, flanking Marginal Zones composed largely of reworked rocks of the granite-greenstone terrains of the adjacent cratons. The belt is a remnant of a larger feature. At its ends, it is buried beneath younger, mainly sedimentary rocks (Fig. 4). In the east, it is truncated by the edge of the continent that formed during the break-up of Gondwanaland. In the west, it is truncated by the Kalahari line, a crustal suture of uncertain age believed to mark the edges of the Rhodesian and Kaapvaal Cratons. The Central Zone was created by the formation of a graben or aulacogen in a protocraton. A remnant of this protocraton is preserved as the Sand River Gneisses (Fripp, 1981) that form part of the basement complex in the Central Zone. These gneisses yield a Rb-Sr whole-rock isochron age of - 3 7 9 0 Ma that is interpreted to reflect the fabric-forming event in these rocks (Do). No rem-
nants of > 3790 Ma rocks of the protocraton have yet been recognized in the adjacent Marginal Zones or cratons. The Sand River Gneisses were intruded by gabbroic dykes - 3 5 7 0 Ma ago, presumably during the initial stages of graben formation. The graben was floored by continental crust and was largely filled with clastic sediments derived from erosion of remnants of the adjacent protocraton. An aqueous depository may be deduced from the presence of chemically precipitated rocks. Carbonate rocks occur near the shallower edges of the graben and banded iron formation occurs in the centre and possibly deeper portion of the basin. There are very few volcanic or volcanoclastic rocks in the sedimentary sequence but a layered gabbroic intrusion, the Messina Layered Intrusion, was emplaced into the supracrustal rocks - 3270 Ma ago. After emplacement of the Messina Layered Intrusion, the graben and rocks of the adjacent segments of the protocraton were deformed during four major events. The first three events resulted from Riedal shear across the graben caused by differential movements between the segments of the protocraton. The first event (Dj) occurred between - 3 2 7 0 Ma ago and - 3 1 5 0 Ma ago and resulted from right-lateral motion between the segments of the protocraton, producing folding and thrusting along north-south trending axial planes. The second event (D2), the major fabric-forming and granulite-grade metamorphic event affecting the supracrustal rocks and the Messina Layered Intrusion, occurred —3150 Ma ago and resulted from left-lateral motion between the adjacent cratons, producing folding and thrusting along north-northwest-south-southeast axial planes. The third deformational event (D3) occurred between - 2 7 0 0 Ma and - 2 6 0 0 Ma ago and also resulted from left lateral motion producing refolding along similarly north-northwestsouth-southeast trending axial planes. Some - 2 7 0 0 Ma granitic plutonism preceded or accompanied the D 3 event northwest of Messina. The fourth deformational event (D4) -2600 Ma ago resulted from clockwise rotation of the Rhodesian Craton with respect to the Kaapvaal Craton around a point in eastern Botswana. This rotation thrust a portion of the Central Zone over the Rhodesian Craton and a portion of the Kaapvaal Craton over the Central Zone. It produced a series of imbricate thrust sheets in the southern part of the Rhodesian Craton, bringing granulitefacies rocks possibly from the Central Zone to the surface and creating the Northern Marginal Zone. It also rotated a portion of the Kaapvaal Craton partially on edge, bringing granulitefacies rocks to the surface and creating the
28°E
32°E
30°E
C O V E R ROCKS
GREENSTONE
f r
BELTS
FAULT AND S H E A R ' Z O N E S T H E C E N T R A L ZONE
BOUNDING
, BOUNDARY OF G R A N U L I T E F A C I E S ROCKS OF L I M P O P O M O B I L E B E L T
23°S
0 I
I
I
I
I I I I kilometres
CRUSTAL EVOLUTION IN SOUTHERN AFRICA
NORTHERN MARGINAL ZONE
2I°S L
100 1—I—J
24° S
Fig. 2.
A geological map of the Limpopo Mobile Belt and adjacent cratons (from Barton & Key, 1981, Fig. 2). 23
24
J. M. BARTON JR
The tectonic development of the Limpopo Mobile Belt is summarized in Table I and is shown schematically in Figure 5. In general, it may be viewed as the formation and destruction of an intracratonic basin with only minor associated igneous activity. Oceanic crust was not emplaced during development of the basin and subduction did not play a direct role in its destruction. Deformation of the belt resulted from externally applied stresses, and nearly every tectonic event that affected the adjacent cratons also affected the belt.
Fig. 3.
A geologic map of the Barberton GraniteGreenstone Terrain showing the locations of . the rock units listed in Table II. The horizontal pattern represents cover rocks including those of the Transvaal and Karoo Supergroups. The vertical pattern represents the rocks of the Ushushwana Complex and the Pongola Group. The black areas represent greenstone-belt material and the blank areas represent leucocratic orthogneisses and granitoid igneous rocks. Rock units I, K, Q and V plot in the same position as Unit T.
Southern Marginal Zone. This thrust movement bent the Central Zone slightly to the northeast (Figs 4 & 5) and probably resulted from crustal extension and thinning preceding the formation of the Witwatersrand Basin and the emplacement of the Great Dyke complex. It may also have been related to the coeval tectonic activity between the Matsitama greenstone belt and Pikwe (Fig. 2). Uplift and erosion followed the D 4 event and by 1950 Ma ago, the area was nearly at its present erosional level. Sediments derived from erosion of the Limpopo Mobile Belt in part filled the Witwatersrand and Transvaal Basins. Movement along the fault zones bounding the Central Zone and those conjugate to these controlled the pattern of minor sedimentation in the Belt during Waterberg, Soutpansberg and Karoo times.
THE BARBERTON GRANITEGREENSTONE TERRAIN Geographically, the Barberton Granite-Greenstone Terrain (Fig. 3) consists of the greenstone belt and plutons of leucocratic orthogneisses and undeformed granitoid rocks surrounding and intruding it. The greenstone belt has been interpreted as consisting of an over 20 km thick sequence of volcanic and associated sedimentary rocks (Onverwacht Group) overlain conformably by the predominantly sedimentary Fig Tree and Moodies Groups (Visser, et al.f 1956; Viljoen & Viljoen, 1969a; Anhaeusser, 1973; Eriksson, 1979). The volcanic rocks include komatiites which become less abundant upwards in the succession. They have undergone greenschist-facies metamorphism and relict igneous textures show
Fig. 4.
A geological map of southern Africa showing the buried extremities of the Central Zone of the Limpopo Mobile Belt (L) (after Barton & Key, 1981, Fig. 7). The diagonal pattern represents post-Archaean cover rocks. RC = Rhodesian Craton. KC = Kaapvaal Craton. K = Kalahari line. M = Makgadikgadi line. The Kalahari line is a major crustal suture marking the western boundaries of the Kaapvaal and Rhodesian Cratons. The Makgadikgadi line is probably a thrust zone along which rocks of the Damaran orogen are thrust southeastward onto the Rhodesian Craton.
25
CRUSTAL EVOLUTION IN SOUTHERN AFRICA
ferent and thinner, being repeated in recumbent folds and imbricate thrust sheets (Williams & Furnell, 1979). In addition, some lithologic differences previously attributed to separate formations may be fades differences within individual formations. Rocks of the Komati Formation (lower Onverwacht Group) yield a Sm-Nd whole-rock isochron age of —3510 Ma (Table II) which is interpreted to reflect the time of emplacement of the unit (Hamilton et al. 1979). The rocks of the Onverwacht, Fig Tree and Moodies Groups were locally deformed twice, the second time— 3450 Ma ago (Table II) in part resulting from the intrusion of the Kaap Valley and Theespruit plutons (Ramsay, 1963; Viljoen & Viljoen, 1969b) (Fig. 3). It appears, therefore, that the emplacement of the rocks of the greenstone belt and two successive deformational events occurred rapidly within a time span of - 60 Ma. This time span is similar to those encompassing Phanerozoic orogenic events. The greenstone belt near the Eureka syncline subsequently deformed a third time (Ramsay, 1963). The greenstone belt is intruded and surrounded by plutons of diverse age that are composed of leucocratic orthogneiss and undeformed granitoid rocks. Near the plutons, the metamorphic
A. > 3 5 7 0 m.y. B. 3 5 7 0 m.y. C. 3 5 7 0 - 3 2 7 0 m.y. D. 3 2 7 0 - 3 1 5 0 m.y E. 3150 m.y. 8 2 7 0 0 - 2 6 0 0 m.y.
9
F. 2 7 0 0 - 2 6 0 0 m.y. G. 1770 m.y. a 2 0 0 m.y.
Fig. 5. A schematic diagram showing the evolution of the Limpopo Mobile Belt (after Barton & Key, 1981, Fig. 7). The explanation is given in the text and in Table I.
that the rocks never were completely recrystallized during metamorphism. Gravity studies indicate that the greenstone belt is a remnant with a preserved thickness of no more than 6 km and underlain by rocks of densities similar to those that intrude it (Burley et al., 1970; Darracott, 1975). In addition, recent mapping and structural interpretations suggest that previous interpretations of the stratigraphic sequence are incorrect and that the original sequence may have been dif-
TABLE
I
Tectonic development of the Limpopo Mobile Belt * Age (Ma) * *
Description
^200
R e a c t i v a t i o n o f T u l i and S o u t p a n s b e r g T r o u g h s ; d e p o s i t i o n o f Karoo S u p e r g r o u p and N u a n e t s i p l u t o n s .
%730
Emplacement o f B e i t B r i d g e k i m b e r l i t e
M770
Formation of T u l i
M9O0
D e p o s i t i o n o f Waterberg S u p e r g r o u p and M a r g i n a l Zones.
M950
Establishment of veil erosion.
%2200
Emplacement o f p o s t - k i n e m a t i c m a f i c dykes and Mahalapye p l u t o n
2600 -
1950
5G).
i n t r u s i o n o f m a f i c dykes
.
in w e s t e r n p a r t o f B e l t and emplacement o f p o s t - k i n e m a t i c m a f i c dykes •
of r a d i o m e t r i c ages o v e r B e l t and a d j a c e n t c r a t o n s
Emplacement o f p o s t - k i n e m a t i c g r a n i t i c p l u t o n s M a r g i n a l Zone.
^2550
( F i g . 5G);
pipes.
and S o u t p a n s b e r g T r o u g h s ; d e p o s i t i o n o f S o u t p a n s b e r g Group ( F i g .
in C e n t r a l
in S o u t h e r n M a r g i n a l
Zone in
Zone and Great Dyke complex in N o r t h e r n
U p l i f t and e r o s i o n t o near p r e s e n t s u r f a c e l e v e l , r e s u l t i n g d e t r i t u s f i l l i n g Kaapvaal c r a t o n s . P o s s i b l y some d e f o r m a t i o n i n w e s t e r n p o r t i o n o f B e l t .
i n t r a c r a t o n i c b a s i n s on R h o d e s i a n and
2700 - 2600
Third period o f deformation
^2700
Emplacement o f B u l a i
pluton
3050 - 2950
Emplacement o f m a f i c
dykes.
^3150
Second p e r i o d o f d e f o r m a t i o n (D 2 ) and major f a b r i c f o r m i n g e v e n t a t g r a n u l i t e grade in s u p r a c r u s t a l r d c k s o f C e n t r a l Zone r e s u l t i n g from l e f t l a t e r a l motion between a d j a c e n t c r a t o n s ( F i g . 5 E ) .
(Fig. .
3150
(D 3 ) a f f e c t i n g C e n t r a l
F i r s t period of deformation (Fig.
in C e n t r a l
from l e f t
lateral
motion between a d j a c e n t
cratons
Zone perhaps s y n t e c t o n i c a 1 1 y w i t h D 3 .
(Di) a f f e c t i n g
the C e n t r a l
Zone r e s u l t i n g
from r i g h t
lateral
motion between a d j a c e n t
5D).
Emplacement o f M e s s i n a Layered 3270
Zone r e s u l t i n g
5E).
eratons >3270' 3570 -
and
Botswana.
F o u r t h p e r i o d o f d e f o r m a t i o n (D 4 ) a f f e c t i n g C e n t r a l Zone - c l o c k w i s e r o t a t i o n o f R h o d e s i a n c r a t o n w i t h r e s p e c t to Kaapvaal c r a t o n ( F i g , 5 F ) . FORMATION OF MARGINAL ZONES by b r i n g i n g g r a n u l i t e f a c i e s r o c k s t o s u r f a c e by t h r u s t faulting. Emplacement o f Matok and S c h i e l complexes in S o u t h e r n M a r g i n a l Zone.
^2600
3270 -
in C e n t r a l
r e l a t e d to t e r m i n a t i o n o f u p l i f t
Formation of f a u l t
Intrusion
in Central
Zone.
bounded b a s i n and emplacement o f s u p r a c r u s t a l
^ 3 570
Emplacement o f m a f i c dykes OF LIMPOPO MOBILE BELT.
^3790
Major p e r i o d of deformation o f p r o t o c r a t o n
>3790
E s t a b l i shment
(DQ).
of.protocraton.
* M o d i f i e d from T a b l e 1 o f B a r t o n & Key * * Based on the decay c o n s t a n t s
rocks
in C e n t r a l
Zone ( F i g .
in basement complex p r i o r t o graben o r a u l a c o g e n f o r m a t i o n
(1981).
recommended by S t e i g e r & J a g e r
(1977).
5C).
( F i g s 5A & 5 B ) .
ESTABLISHMENT
J. M . B A R T O N J R
26
TABLE II
Radiometric age data, Barberton Granite-Greenstone Terrain
1
Unit** (A) (A1) (B) (c) (D) (E) (F) (G) (H) (1) (J) (K) (L) (M) (N) (0) (P) (Q) (R) (S) (T) (U) (V) (w) •(X) (Y) (z) (2) (3) (4) (5) (6) (7) (8) (9) (Al). (B1) (CI) (D1) (El) (Fl) (Gl) (HI)
Leucocratic Orthogneiss Facies, Bimodal Suite, Ancient Gneiss Complex, Swaziland Komati Formation, Onverwacht Group Theespruit Pluton Granodiorite Suite, Swaziland Tsawela Gneiss, Swaziland Banded Orthogneiss near Pigg's Peak Basement Orthogneiss near Phalaborwa Hebron Orthogneiss Nelshoogte Pluton Leucocratic Orthogneiss Porphyroblastic Orthogneiss Leucocratic Orthogneiss Banded Orthogneiss Orthognei ss Lochiel (Homogeneous Hood) Pluton Homogeneous Granitic Orthogneiss Granitic Orthogneiss Leucocrat ic Orthognei ss Migmatltic Orthogneiss Salisbury Kop Pluton Leucocratic Orthogneiss Mliba Pluton Lamprophyric Dyke Boesmanskop Pluton Granophyre, Ushushwana Complex Cunning Moor Pluton Sicunusa Pluton Mgapeni Pluton Kaap Valley Pluton Garnetiferous Orthogneiss Dal me in Pluton Doornhoek Pluton Stentor Pluton Batavia Pluton Sinceni Pluton Kweta Pluton Nhlangano Gneiss Ngwempisi Gneiss Hlatikulu Pluton Nponono Layered Intrusion Quartz-Porphyry, Msauli Mine Granodioritic Gneiss, Johannesburg Dome Leucocratic Orthogneiss, Johannesburg Dome
Rock Age* 3555
1112
U-Pb Zircon & Sphene Age*
Rb-Sr BiotiteWhole Rock Age*
3250 ± 8 0
s
2698 ± 54*
3220 ± 80
5
2993 ± 60" 2880 ± 58*
2 0.6999 + 0 . 0 0 1 6
3510 ± -60 3 3432 ± 135* 3350 ± 57* 6 3323 ± 168 6 3270 ± 104 3260 ± 90* 7 3211 ± 133 3 1 8 0 ± 75* 3 1 8 6 ± 74* 3149 ± 125 7 3147 ± 2 1 * s 3138 + 112 6 3072 + 1 2 3028 ± 14* 2986 ± 69 6 8 2955 ± 104 2939 ± 75* 2927 ± 137 1 0 2927 ± 59 s 2916 ± 33* 6 2879 ± 88 2866 ± 70* 2848 ± 31* 2813 ± 59 9 * 7 2784 ± 53 2608 ± 123 6 2496 ± 176 1 0 3482 ± 166* 2734 ± 226* 3201 ± 43* 3192 ± 46* 2755 ± 51* 2769 ± 54*
0.7001 ± 0 .00062 0.7000 ± 0 .0019* 0 . 7 0 1 1 ± 0 .0008* 6 0 . 7 0 0 6 ± 0 .0024 0 . 7 0 2 2 ± 0,. 0 0 2 6 6 0.7000 + .0013* 7 0.7007 ± 0..0011 0 . 7 0 1 0 + 0.. 0 0 1 1 * 0.7014 ± 0.. 0 0 1 0 * 0.7014 ± 0..0010 7 0.7040 ± 0..0003* 0.7048 ± 0,. 0 0 2 2 6 6 0 . 7 0 6 0 + 0..0004 0.7013 ± 0..0002* 0.7054 ± 0.0036 0.7036 ± 0..0010° 0.7021 ± 0..0011* 10 0.7052 ± 0..0020 5 0.7033 ± 0.,0012 0.7018 + 0..0005* 0.7024 ± 0.,0006 6 0.7028 ± 0.,0010* 0.7040 ± 0.,0004* 9 0.7029 ± 0. 0 0 7 2 0.7034 ± 0. 0003 7 0.7006 ± 0. 0 1 8 8 6 10 0.7065 ± 0. 0 0 3 2 0.7001 ± 0. 0024* 0.7029 ± 0. 0032* ± 0.7020 0.,0006* 0.7232 ± 0. 0007* 0.7029 ± 0. 0007* 0.7023 ± 0. 0008*
2766 ± 21* 3081 ± 148 1 1 3009 ± 8 p i
0.7059 ± 0. 0003* 0.7004 ± 0. 0021 1 1 0.7027 ± 0. 0 0 1 1 1 1
* Uncertainty expressed as 2a. *<> See Figures 1 & 3 for sampling locations. 1 Determined using the decay constants recommended by Steiger & Jager 2 Barton et ai. (1980). 3 Hami1 ton et al. (1979). * J.M. Barton Jr, Unpub. Data. 5 Oosthuyzen (1970). 6 Davies (1970). 7 E.S. Barton, Unpub. Data. 8 H.L. Allsopp, Unpub. Data. 9 Davies et al. (1970).• 10 de Gasparis (1967). 11 M. Callow & J.M. Barton Jr, Unpub. Data.
grade of the greenstones has been raised to amphibolite fades and away from the greenstone belt, some of the plutons contain enclaves of what may have been greenstone-belt material (Anhaeusser & Robb, 1978). The older plutons have metamorphic fabrics and are intruded by both deformed and undeformed plutons. Some of the metamorphic fabrics appear to be related to solid-state diapiric rising of the plutons, while others may have resulted from deformation associated with diapiric emplacement of younger plutons. Igneous and sedimentary rocks of the greater than-2850 Ma old Pongola Group and Ushushwana complex (Fig. 3) were deposited on a terrain of older plutons and intruded by younger ones.
+
7 86 Initial 8 S r / S r *
0,
2889 ± 58 2
2 8 1 1 ± 56*
3160 ± 190 5
2796 ± 5 6 * 3075 ± 100 5
3048 ± 6 1 " 2601 ± 52*
3280 ± 196 5
3130 ± 188
2800 ± 168 5 3270 ± 196 s 3170 ± 190
2677 ± 54 s 2800 ± 56"
s
2320 ± 4 6 1 0
5
(1977).
Age data from the greenstones, leucocratic orthogneisses and undeformed granitoid rocks are summarized in Table II and sampling locations are shown on Figure 3. No unit yet studied has a radiometric age appreciably older than the -3510 Ma Komati Formation and only three units, the leucocratic orthogneiss component of the Bimodal Suite of the Ancient Gneiss Complex (Barton et al., 1980) and the Kaap Valley and Theespruit plutons, are coeval with it. Therefore, no substantially older continental basement to the greenstone-belt succession has yet been identified. The Rb-Sr whole-rock ages (Figs 6 & 7) may be divided into four groups with a small amount of possible overlap. Group 1 clusters at 3450 ±
27 to represent sub-oceanic mantle (e.g. Davies et al., 1970; Faure & Powell, 1972) already depleted in Rb with respect to Sr and suggests that the Komati Formation and several orthogneiss units have close affinities with sub-oceanic mantle (Barton et al., 1980). This contention is further supported by Sm-Nd studies of bodies of sodic porphyry intrusive into the Komati Formation, analytical results from which plot on the wholerock isochron for the latter unit and yield an initial i 3Nd/ Nd ratio that falls on a reasonable mantle Nd-isotopic growth curve (Hamilton etal., 1979). It has been suggested that the lower portion of the Onverwacht Group, including the Komati Formation, could represent oceanic crust (e.g. Anhaeusser, 1973; de Wit & Stern, 1980). Furthermore, REE analyses indicate that many of the tonalitic and granodioritic orthogneiss units in the Barberton Granite-Greenstone Terrain could have been derived from partial melting of basalt or amphibolite compositionally similar to the Komati Formation (Hunter et al,, 1978). It is reasonable to suggest, therefore, that those plutons with initial Sr/ Sr ratios near the postulated sub-oceanic mantle growth curve (Fig. 7) could have been derived by partial melting of oceanic crust and that the Komati Formation and probably the entire lower Onverwacht Group are
CRUSTAL EVOLUTION IN SOUTHERN AFRICA
50 Ma and is composed of the oldest intrusive plutons, ranging in composition from tonalitic to granodioritic orthogneisses. Group 2 clusters at 3250 ± 150 Ma and is composed of tonalitic to granitic orthogneisses. Group 3 clusters at 2900 ± 150 Ma and contains both deformed and undeformed rocks of tonalitic to granitic and syenitic composition. Group 4 clusters at 2550 ± 50 Ma and is composed primarily of granite. The Rb-Sr whole-rock ages are consistent with their structurally determined relative ages and are believed to reflect emplacement. All units studied thus far in the Barberton Granite-Greenstone Terrain, with the exception of the Doornhoek pluton, have low initial Sr/ Sr ratios (<0.7065) (Table II) and there is a general inverse correlation between initial Sr/ Sr ratio and age (Fig. 6) but no correlation between initial Sr/ Sr ratio and K-content of the rock unit. For example, the granite of the Lochiel (Homogeneous Hood) pluton has an initial Sr/ Sr ratio of 0.7013 but granodioritic orthogneiss (Unit K) has one of 0.7040 (Table II). Some plutons and the Komati Formation have initial Sr/ Sr ratios defining a linear Sr-isotopic growth curve passing through a present day Sr/ Sr ratio of 0.7025 (Fig. 7) and corresponding to a Rb/Sr ratio of 0.017. This growth curve closely approximates those commonly postulated 87
86
87
86
87
87
86
87
86
87
8 6
86
4
144
87
86
710 -
CO CD 00
.700 3.5
3
AGE (I0 9 years)
Fig. 6. A diagram of initial S r / S r ratio versus age for the Rb-Sr whole-rock age data presented in Table II. The crosses represent the 2a uncertainties in the data. 87
86
J. M. BARTON JR
28
.710 -
cn CO
•V / -/6 J. yCv Fl
00
w
J -705100
u
v .700-
A
^
Nj
7 V
*
B
Gpl 3.5
Gp 3
Gp2 3
I Gp 4 2.5
AGE (I0 9 years)
Fig. 7. A similar diagram to Figure 6 showing the growth vectors for the units listed in Table II. The solid line represents the postulated oceanic mantle Sr-isotopic growth curve and passes to an 8 S r / S r ratio of 0.7025 for the present. 7
remnants of that oceanic crust. The oceanic crust itself was probably formed from partial melting of sub-oceanic mantle (Ringwood, 1974). The Sr/ Sr isotopic growth vectors for the units listed in Table II are plotted on Figure 7. As noted previously by Davies & Allsopp (1976), many of these vectors are nearly colinear, suggesting either the ages represent local episodic rehomogenization of the Sr-isotopes in a relatively homogeneous unit or local melting of portions of a fairly homogeneous source in such a way that the Rb/Sr ratios of the melts and of the source were constant. Neither possibility can be proven correct in most cases. However, data from the Boesmanskop pluton (Unit W) provide some clues (Anhaeusser et al., 1979). This unit is intruded into deformed tonalitic orthogneiss plutons including Unit T. The Boesmanskop pluton yields a Rb-Sr whole-rock isochron age of -2850 Ma and Unit T one of -2916 Ma (Table II). These ages are distinct at the 95% confidence level and are consistent with their structurally determined relative ages. However, zircons from the Boesmanskop pluton yield a U-Pb age of ~3130Ma (Table II), the same age as indicated by the intersection of the Sr-isotopic growth vector for this unit with the postulated suboceanic mantle Sr-isotopic growth curve (Fig. 7). 87
8 6
86
This close age agreement suggests that the precursor of the Boesmanskop pluton formed ~3130Ma ago from material closely related to oceanic crust but that it was remobilized -2850 Ma ago and emplaced in its present position. The Boesmanskop pluton was emplaced as predominantly a liquid. Therefore, the remobilization involved melting during which the Srisotopes were homogenized but the U-Pb zircon "clocks" were not reset. The zircons probably remained crystals during remobilization, i.e. they are xenocrysts. The Salisbury Kop Pluton, Unit S on Figure 7, with its— 2930 Ma Rb-Sr whole-rock isochron and ~ 3280 Ma U-Pb zircon ages (Table II), probably has had a similar history. In fact, it is proposed that most of the Group 2 and 3 and Group 4 plutons could represent remobilized units that originally formed as partial melts of oceanic crust between -3400 Ma and -3100 Ma ago and 2600 Ma ago respectively. Those units with higher initial Sr/ Sr ratios were remobilized and emplaced at higher crustal levels during subsequent periods of deformation while those with initial Sr/ ^Sr ratios near the postulated oceanic mantle Sr-isotopic growth curve were emplaced shortly after magma genesis. The agreement between Rb-Sr whole-rock isochron 87
86
87
8
CRUSTAL EVOLUTION IN SOUTHERN AFRICA 29 and U-Pb zircon ages is best in those units with is identical to its whole-rock isochron and zircon low initial ^ S r / ^ S r ratios (Table II). ages, indicating the unit was emplaced at high Oceanic crust may be partially melted as a crustal level and cooled quickly. It has not subseresult of two tectonic mechanisms: subduction, quently been metamorphosed and, therefore, all and folding and burial to great depth. Both regional metamorphism and deformation near mechanisms could have operated during destruc- the Barberton greenstone belt, including D , had tion of the basin of which the Barberton green- occurred by — 3050 Ma ago. stone belt is a remnant. The source of the clastic The tectonic evolution of the Barberton sediments of the Fig Tree and Moodies Groups Granite-Greenstone Terrain is outlined in Table was to the south. These sedimentary units were III. The most plausible modern analogue for the deposited in a narrow continental shelf-type en- mechanism creating this terrain is the formation vironment as might have existed along the margin and destruction of a back-arc basin between two of a fault-bounded basin, and they were derived pieces of continental crust on one of which is from a rapidly eroding source area composed of superimposed an island arc (Tarney et al., 1976). both volcanic and continental rocks (e.g. Condie This situation exists today in the Sea of Japan. et at., 1970; Eriksson, 1979, 1980). The amount The komatiites of the lower Onverwacht Group of continental detritus increases upwards through would be formed during the initial break-up of the Fig Tree and Moodies Groups. Furthermore, the continental'crust and the volcanism of the boulders in conglomeratic beds within the upper Onverwacht Group would accompany Moodies Group are compositionally similar to spreading. The island arc would supply detritus the orthogneiss plutons intruding the greenstone for the Fig Tree and Moodies Groups and partial belt and yield U-Pb zircon ages that suggest that melting of subducted oceanic crust would provide their source was older than -3300Ma (van the magmas for some of the early plutons inNiekerk & Burger, 1978). This indicates that the truding both the island arc and the back-arc source terrain was composed of volcanic rocks basin. Deformation of the basin due to changes in resting on continental crust, both of which were the rate of subduction would fold some of the intruded by plutons of material derived from par- oceanic crust in the back-arc basin to depths tial melting of oceanic crust. Such a terrain can be where it could partially melt to form magmas for created by subduction of oceanic crust under con- other plutons. This process would create a zone tinental crust such as is happening in the Andes of new continental crust between older continenand Japan today. However, many of the younger tal crust and would allow for a growth of contiplutons apparently formed from partial melting nental nuclei or cratons. The younger —3540 of oceanic crust after the destruction of the basin Ma old plutons would be emplaced during subse-3450 Ma ago. The source of these plutons could quent, and probably unrelated, periods of therhave been oceanic crust subducted beneath the mal activity below the deformed back-arc basin. Barberton Granite-Greenstone Terrain during Some of this activity could be the result of introsubsequent events. More probably, however, it duction of another subduction zone beneath the was oceanic crust folded and deeply buried dur- basin. ing the destruction of the basin that was partially Remnants of the >-3550Ma old, continental melted during later thermal events. crust presumably surrounding the back-arc basin have not been positively identified. However, by It is not known how far away from the Barberton Granite-Greenstone Terrain this proposed comparing the early -3550 Ma to-3150 Ma relationship between oceanic crust and plutonism evolution of the Barberton Granite-Greenstone might apply. However, the initial Sr/ Sr ratios Terrain (Table III) with that of the Limpopo of the — 3260 Ma old basement orthogneisses Mobile Belt (Table I), it can be seen that tectonic near the Palabora complex south of the Murchi- activity in these areas has occurred sympathetison greenstone belt and the - 3100 Ma old grano- cally. It is reasonable to postulate, therefore, that dioritic orthogneisses from the Johannesburg the protocraton in which the Limopo Mobile Belt Dome (Figs 1 & 2) also plot on the postulated was formed was one portion of continental crust oceanic mantle Sr-isotopic growth curve (Fig. 7), bounding the back-arc basin. The other greensuggesting similar evolution patterns may have stone belts between the Barberton greenstone belt and the Limpopo Mobile Belt (Figs 1 & 2) would occurred in these parts of the Kaapvaal Craton. Rb-Sr biotite-whole rock ages (Table II) for the probably be remnants of the same basin. The other piece of continental crust on which Barberton Granite-Greenstone Terrain range from — 3050 Ma to -2300Ma, generally reflect- the island arc presumably was situated must be south of the Barberton greenstone belt. Signifiing post-emplacement cooling after metamorphism. However, the-3050 Ma biotite-whole rock cantly, the metamorphic grade increases to age of the Lochiel (Homogeneous Hood) pluton granulite facies going southward in Swaziland 3
87
86
J. M. BARTON JR
30 TABLE III
Tectonic development of the Barberton Granite-Greenstone Terrain Age (Ma)*
Description
250O- to Present
This area has been a stable portion of the Kaapvaal Craton during this time but it was probably covered and subsequently denuded of sediments and volcanic rocks of the Transvaal and Karoo Supergroups.
2550±50
Fourth period of pluton
^2850
Deposition of the sedimentary rocks of the Pongola Group and emplacement of the Ushushwana Complex.
emplacement
in the region.
2900 ± 150
Third period of. pluton emplacement
3 2 5 0 ± 150
Second period of pluton emplacement
in the region.
3 4 5 0 ± 50
First period of pluton emplacement
in the area
3500-3450
occurring contemporaneously with D 2 in the greenstone belt. Filling of the basin with volcanic and sedimentary rocks (the upper Onverwacht, Fig Tree and Moodies Groups) followed by two periods of deformation during which the basin was destroyed.
in the region.
3550-3500
Formation of a basin floored by oceanic crust (the lower Onverwacht Group) by rifting apart of the protocraton. Some plutonism manifested in the Ancient Gneiss Complex may have occurred at this time.
>3550
Formation of a protocraton, possibly the same one in which the Limpopo Mobile Belt was formed.
Based on the decay constants recommended by Steiger & Jager (1977).
(Jackson, 1979). Although the geological history of this region has been complicated by younger tectonism related to emplacement of the rocks of the Pongola Group and the Ushushwana complex, the granulite-facies metamorphism preceded this activity. Therefore, it is possible that this granulite-facies metamorphism occurred in rocks deeply buried beneath the island arc. The granulite-facies rocks include some of the Ancient Gneiss complex which Hunter (1970) proposed on structural evidence to be a basement to the Barberton greenstone belt. Not all of the rocks grouped in the Ancient Gneiss complex are reasonable candidates for this basement (Barton et al., 1980) but those in the type area near Mankayane may be. CONCLUSION From the foregoing discussion, it may be concluded that evolution of both the Limpopo Mobile Belt and the Barberton Granite-Greenstone Terrain began more or less coevally -3550Ma ago by rifting of continental crust to form fault-bounded basins. The Limpopo Mobile Belt probably formed as an aulacogen while the Barberton Granite-Greenstone Terrain likely
originated as a back-arc basin between two pieces of continental crust. In the Limpopo Mobile Belt, the basin was filled primarily with detritus of continental origin while the back-arc basin, being genetically linked to an island arc, was filled first with volcanogenic rocks and only in the later stages by continental sediments. The back-arc basin was destroyed fairly rapidly after formation by a combination of compression between cratons and plutonism resulting in part from partial melting of oceanic crust during subduction. The Limpopo Mobile Belt, on the other hand, existed as a basin until at least -3270 Ma ago. It was ultimately destroyed by Riedal shear resulting from strike-slip motion between the adjacent cratons. Only in the latest stages was compression across the belt a major tectonic mechanism. The last period of plutonism in the Barberton Granite-Greenstone Terrain and the last major period of deformation in the Limpopo Mobile Belt occurred -2600 ± 100 Ma ago. It appears, therefore, that plate tectonic processes can explain the patterns of evolution of the Limpopo Mobile Belt and the Barberton GraniteGreenstone Terrain during the Archaean. Although this evolution was not parallel in the two areas, plate tectonic interpretation can provide a cause-and-effect link between them. It may not be necessary, therefore, to invoke special nonuniformitarian mechanisms to explain early crustal evolution in southern Africa. ACKNOWLEDGMENTS During the past five years, the author has had the privilege to work with the following persons on the study of the Limpopo Mobile Belt and the Barberton Granite-Greenstone Terrain during the term of the International Geodynamics Project: H. L. Allsopp, C. R. Anhaeusser, E. S. Barton, M. C. du Toit, K. A. Eriksson, R. E. P. Fripp, P. C. Horrocks, D. R. Hunter, M. P. A. Jackson, R. M. Key, L. J. Robb, B. Ryan, D. D. van Reenen, M. K. Watkeys, H. J. Welke and A. C. Wilson. Ideas do not form in a vacuum and most of those presented in this paper were formulated through interaction with these people. I. D. M. Robertson, C. B. Smith and D. A. C. Williams are thanked for critically reviewing this manuscript.
REFERENCES
ANHAEUSSER, C. R., 1973: The evolution of the early
Precambrian crust of southern Africa. Phil. Trans. R. Soc. Lond., A273, 359-388. ANHAEUSSER, C. R . , & ROBB, L. J., 1978: Regional and detailed field and geochemical studies of Archaean
trondhjemitic gneisses, migmatites and greenstone xenoliths in the southern part of the Barberton Mountain Land, South Africa. Inform. Circ. Econ. Geol. Res. Unit, Univ. Witwatersrand, Johannesburg, 125.
CRUSTAL EVOLUTION IN SOUTHERN AFRICA ANHAEUSSER, C . R . , ROBB, L . J . , & BARTON, J . ML,
1979: Mineralogy, petrology and origin of the Archaean Boesmanskop syenite, Barberton Mountain Land, South Africa. Inform. Circ. Econ. Geol. Res, Unit, Univ. Witwatersrand, Johannesburg, 139.
BARTON, J . M L , J R . , HUNTER, D . R . , JACKSON, M . P . A., & WILSON, A. C., 1980: Rb-Sr age and source
of the Bimodal Suite of the Ancient Gneiss Complex, Swaziland. Nature, Lond., 283, 7 5 6 - 7 5 8 . BARTON, J . M., J R . , & KEY, R . M., 1981: The tectonic development of the Limpopo Mobile Belt and the evolution of the Archaean cratons of southern Africa; in Kroner, A., (Ed.), Precambrian Plate Tectonics, 185-212. Elsevier, Amsterdam.
FAURE, G., & POWELL, J . L., 1972: Strontium
Geology. Springer-Verlag, Berlin.
31 Isotope
FRIPP, R. E. P., 1981: The ancient Sand River Gneisses,
Limpopo Mobile Belt, South Africa. Spec. Pubis geol. Soc. Aust., 7, 329-335.
HAMILTON, P . J . , EVENSEN, N . M . O'NIONS, R . K . , SMITH, H . S . , & ERKLANK, A . J . , 1979: S M - N d
dating of Onverwacht Group volcanics, southern Africa. Nature, Lond., 279, 298-300.
HUNTER, D. R . , 1970: The Ancient Gneiss Complex in
Swaziland. Trans, geol. Soc. S. Afr., 73, 107-150.
HUNTER, D . R . , BARKER, F . , & MILLARD, H . T . , J R . ,
1978: The geochemical nature of the Archaean Ancient Gneiss Complex and Granodiorite Suite, Swaziland: a preliminary study. Precamb. Res., 7, BURLEY, A . J . , EVANS, R . B . , GILLINGHAM, J . M . , & 105-127. MASSON SMITH, D . , 1970: Gravity anomalies in M. P. A., 1979: High-strain deformation Swaziland. Bull. geol. Surv. Mines Dep. JACKSON, history of the Ancient Gneiss Complex in the ManSwaziland, 7, 4-16. kayane area, Swaziland: a preliminary report. CON DIE, K . C . , MACKE, J . E . , & REIMER, T . O . , 1970: Abstr. 18th Congr. geol. Soc. S. Afr., Port ElizaPetrology and geochemistry of early Precambrian beth, 210-216. graywackes from the Fig Tree Group, South OOSTHUYZEN, E. J., 1970: The geochronology of a suite Africa. Bull. geol. Soc. Am., 81, 2759-2776. of rocks from the granitic terrain surrounding the DARRACOTT, B. W., 1975: The interpretation of the Barberton Mountain Land. Ph.D. Thesis, Univ. gravity anomaly over the Barberton Mountain Witwatersrand [unpublished]. Land, South Africa. Trans, geol. Soc. S. Afr., 78, RAMSAY, J . G., 1963: Structural investigations in the 123-128. Barberton Mountain Land, eastern Transvaal. DA VIES, R . D . , 1970: Geochronology and isotopic Trans, geol. Soc. S. Afr., 66, 353-398. evolution of the early Precambrian crustal rocks in RINGWOOD, A. E., 1974: The penological evolution of Swaziland. Ph.D. Thesis, Univ. Witwatersrand island arc systems. J. geol. Soc. Lond., 130, [unpublished]. 183-204. DA VIES, R . D . , & ALLSOPP, H . L., 1976: Strontium iso- STEIGER, R. H . , & JAGER, E., 1977: Subcommission on topic evidence relating to the evolution of the lower geochronology: convention on the use of decay Precambrian granitic crust in Swaziland. Geology, constants in geo- and cosmochronology. Earth 4, 553-556. planet. Sci. Lett., 36, 359-362. DA VIES, R . D . , ALLSOPP, H . L . , ERLANK, A . J . , & MANJ . , DALZIEL, I. W . D . , & DE WIT, M . J . , 1976: TON, W. I., 1970: Sr-isotopic studies on various TARNEY, Marginal basin "Rocas Verdes" complex of S. layered mafic intrusions in southern Africa. Spec. Chile: a model for Archaean greenstone belt forPubis geol. Soc S. Afr., 1, 576-593. mation; in Windley, B. F. (Ed.), The Early History DE GASPARIS, A . A . A . , 1967: Rb-Sr isotopic studies of the Earth, 131-146. Wiley, London. relating to problems of geochronology on the Nelspruit and Mpageni granites. M.Sc. Thesis, Univ. VAN NIEKERK, C . B . , & BURGER, A . J . , 1978: T h e a g e o f the Moodies conglomerate boulders. Spec. Pubis Witwatersrand [unpublished]. geol. Soc. S Afr., 4, 99-106. DE W I T , M. J . , & STERN, C . R., 1980: A 3 5 0 0 M a ophiolite complex from the Barberton greenstone belt, VILJOEN, M . J . , & VILJOEN, R . P . , 1969a: An introducSouth Africa: Archaean oceanic crust and its geotion to the geology of the Barberton Granitetectonic implications. Extended Abstracts, 2nd Greenstone Terrain. Spec. Pubis geol. Soc. S. Afr., internal. Archaean Symposium, Perth, 1980, 2, 9-28. 85-87. , 1969b: A proposed new classification of the ERIKSSON, K. A., 1979: Marginal marine depositional granitic rocks of the Barberton region. Spec. Pubis processes from the Archaean Moodies Group, geol. Soc. S. Afr., 2, 153-180. Barberton Mountain Land, South Africa: evidence VISSER, D. J . L. (COMPILER) et al., 1956: The geology of and significance. Precamb. Res., 8, 153-182. the Barberton area. Spec. Pubis geol. Surv. S. Afr., , 1980: Transitional sedimentation styles in the 15. Moodies and Fig Tree Groups, Barberton MounD. A. C., & FURNELL, R. G., 1979: Reassesstain Land, South Africa: evidence favouring an WILLIAMS, ment of part of the Barberton type area, South Archaean continental margin. Precamb. Res., 12, Africa. Precamb. Res., 9, 325-347. 141-160.
CRUSTAL STRUCTURE IN THE PILBARA AND NORTHERN YILGARN BLOCKS FROM DEEP SEISMIC SOUNDING B. J. Drummond, R. E. Smith, & R. C. Horwitz 1
2
2
Bureau of Mineral Resources, Geology and Geophysics, PO Box 378, Canberra City, ACT 2601: currently at Research School of Earth Sciences, Australian National University, PO Box 4, Canberra, ACT2600 CSIRO Institute of Earth Resources, Division of Mineralogy, Private Bag, PO Wembley, Western Australia 6014 1
2
ABSTRACT Two seismic refraction profiles in the Pilbara and northern Yilgarn Cratons of northwest Australia reveal major velocity-depth features in the crust. The Pilbara Craton has a 28-33 km thick, two-layered crust, but in the northern Yilgarn Craton the crust is thicker (50 km) and three-layered. The lower crust in the Proterozoic mobile belt between the cratons has high velocity gradients implying a dense lower crust. Along the axis of the Hamersley Basin, the crust has a velocity inversion close to the surface, and lateral changes in near-surface seismic velocities correspond to zoning in the burial metamorphism of the Hamersley Basin volcanic and sedimentary rocks. Partial melting and metamorphic grade are used to explain velocities in the lower crust that are different from those under the eastern Pilbara Block. The differences in the crustal thicknesses and the number of crustal layers in the Pilbara and Yilgarn Cratons are evidence of separate crustal histories of the cratons during the Archaean.
INTRODUCTION The Pilbara region of northwest Australia is one of the most geologically interesting and important areas of Australia—it has the oldest isotopically dated rocks in Australia and therefore may represent the primeval continental crust of Australia; it is also one of Australia's principal mineral provinces. In 1977, the Australian Bureau of Mineral Resources, Geology and Geophysics (BMR) and the Research School of Earth Sciences, Australian National University (ANU) undertook a survey to study the crustal and upper mantle structure in the Precambrian shield of northwest Australia. The aims of the survey were to define the structural relations between the Archaean Pilbara and Yilgarn Cratons, and to compare the crustal parameters of the exposed parts of the cratons with those of the crust under the Proterozoic sedimentary basins in the region. This paper sets out some of the results and develops crustal models for the region.
vinces follows that of Gee (1979a) and amendments are discussed in the text. The oldest rocks crop out in the Archaean Pilbara Block which has a granite/greenstone stratigraphy. Rb-Sr isotopic ages of the bulk of the granites (Arriens, 1971; de Laeter & Blockley, 1972; de Laeter et al., 1975) are about 3100 Ma, although some granites were emplaced later at about 2700 Ma. U-Pb ages (Oversby, 1976; Pidgeon, 1978a; 19786; 1978c) imply older ages of about 3300-3400 Ma for the bulk of the granites. The Yilgarn Block in the south also has a granite/greenstone stratigraphy, but both Rb-Sr and U-Pb isotopic ages of granites (Arriens, 1971; Oversby, 1975) suggest that granites equivalent in age to the younger granites of the Pilbara dominate in the Yilgarn Block. Volcanics, banded iron formations, and sediments of the Hamersley Basin overlie the southern half of the Pilbara Craton. Archaean inliers within the Hamersley Basin, and trends in the Bouguer gravity anomaly field (BMR, 1975) suggested to Horwitz & Smith (1978, Fig. 1, p. 294) that the craton extended south to include the GEOLOGY Sylvania Dome, and hence that the Hamersley Several Precambrian provinces occur in north- Basin, as later redefined by Gee (1979a, 1979c), is west Australia; those discussed here are shown in within the confines of the Pilbara Craton. Figure 1. Most nomenclature used for the pro- The slightly younger Ashburton Trough Spec. Pubis geol. Soc. Aust., 7 (1981)
B. J. DRUMMOND, R. E. SMITH & R. C. HORWITZ
34
(Doust, 1975) abuts the southwest margin of the Hamersley Basin and contains the sediments of the Wyloo Group. Smith & Horwitz (1975) demonstrated that it coincides with the western part of a gravity feature which Fraser (1976) labelled the Ashburton Regional Gravity Ridge. Horwitz & Smith (1978) named it the Wyloo Trough and extended it to the Glengarry Subbasin of the Nabberu Basin (Hall & Goode, 1978) where similar sediments, beheved to be of the same general age, overlap the northern Yilgarn Block. Sedimentary rocks of the Bangemall Basin unconformably overhe the rocks of the Hamersley and Nabberu Basins and mask their structural relations. The Gascoyne Province, in the west of the surveyed region, contains metamorphosed and folded rehcs of Archaean basement, lower Proterozoic sediments and volcanics, and rocks of the Bangemall Basin (Daniels, 1975; de Laeter, 1976; Williams et al., 1978). Horwitz & Smith (1978, Fig. 1) extended the Gascoyne Province eastward to include the Goodin and Marymia Domes of Gee (1979^?).
Gravity
road
traverse
point
Recording
Piibara
THE CRUSTAL SURVEY During the crustal survey, seismic refraction data and data from road and helicopter gravity traverses were collected (Fig. 1). The Pilbara region has seven open-cut iron ore mines which regularly fire large quarrying blasts suitable for use as seismic sources. To extend the survey
traverse
Gravity helicopter Shot
Palaeogeographic reconstructions by Horwitz & Smith (1978) showed that, during the early sedimentation in the Hamersley Basin, a basement high extended north-northwest from about the northern edge of the Sylvania Dome, and that the southern margin of the Pilbara Craton was tilting or subsiding toward the Ashburton Trough. This is allied to an increase in metamorphic grade towards the south (Smith, 1975a) and an increase in deformation (Daniels, 1966). A similar but reversed picture occurs in the sedimentary cover of the Yilgarn Craton (Hall & Goode, 1978; Horwitz & Smith, 1978), and thus folding and metamorphism indicate that the Gascoyne Province lies along an east-west axis of mirror-image symmetry.
station
and Yilgarn
Cratons
Hamersley
Basin
Ashburton
Trough, Glengarry
Earaheedy
Sub-Basin
Bangemall
Basm
Gascoyne
Province
Phanerozoic
sedimentary
Sub-Basin
Bas/ns
Fig. 1;
Geology and survey design.
CRUSTAL STRUCTURE FROM SEISMIC SOUNDING
farther south, a specially prepared blast was fired at Meekatharra on the northern Yilgarn Block. The field work was described by Drummond (19796, 1979c). Portable field seismographs were deployed at intervals of less than 20 km along three lines (Fig. 1). The principal line (GBC) extended from Goldsworthy to Meekatharra, crossing the eastern Pilbara Block and Hamersley Basin, the Bangemall and Nabberu Basins, and the northern Yilgarn Block. Farther west, a second line (GHDE) extended south to the Gascoyne Province. The third BMR line (FDB) lay along the axis of the Hamersley Basin outcrop area. Several methods were used to interpret the data. The models presented here were derived using the intercept method (Mota, 1954). They are two-dimensional representations of the crust and have several layers in which the velocities generally do not vary with depth. However, studies using synthetic seismogram computer modelling indicate that velocity gradients are likely in the crust, at the crust/mantle boundary, and in the upper mantle. They are also evident from the positions of reflected phases and traveltime cusps in some of the data presented here.
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200
l
SECTION GBC The first model for line GBC was derived and presented by Drummond (1979a), but because line GBC is the principal seismic profile in the survey, this model is reviewed here. Because of the approach to the original interpretation, some of the layers had small steps which have now been smoothed out. Three possible models were presented by Drummond (1979a), and their revised equivalents are shown in Figure 2. Seismic velocities in the upper crust along the profile are between 6.0 and 6.2km s - 1 , except to the north of Newman, where a slightly lower velocity (5.86km s - 1 ) and small negative intercept indicate thickening Hamersley Basin rocks to the north. This agrees with the palaeogeographic reconstruction for the region of Horwitz & Smith (1978). A lower crustal layer (6.4 km s" 1 ) is present along the profile, at depths ranging from 13 km in the north (at G) to 15 km in the south (at C). From the data available, the differences between the three alternative models for the lower crust along segment BC in Figure 2 cannot be resolved. While they differ in detail, they all show a thickening of the crust from north to south. Velocity gradients and material with a seismic velocity of about 7 km s " 1 imply a dense lower crust in the south of the segment. The crust is 28 km thick in the north of the
I;
A
--
T
+
+
. . F . . . T I . . " T . . X
AT- A v I I 0(b) 100 +
Bangemall Basin j
- R — , — , — - I — +
+
+
|
Hamersley Basin
35
U • V
A
8 , - 8 velocity
34 gradient
Alternative models along seismic lines ABC and GBC, based on those of Drummond (1979a).
Pilbara Block, and thickens to 33 km to the south of B, where the steps in models a and b increase its thickness to 37, 43, and finally 52km at C. The thicknesses may be revised slightly by synthetic seismogram modelling; results so far suggest that increases of up to 10% may be required. The upper mantle velocity along segment GB is 8.34km s _ 1 . Travel-times from the Meekatharra blast to seismological observatories to the south indicated an upper mantle velocity of 8.1 km s~ 1 under the Yilgarn Block. Consequently, the upper mantle velocity may decrease from north to south along segment BC. The outcrop areas of the geological provinces have been marked in Figure 2. The Pilbara Craton, whose rocks crop out along much of segment GB, is considered to extend about 80 km south of B, to the most northerly step in the lower crust. This is about 40 km south of the Sylvania Dome, and corresponds to the southern limit of the Northern Fades of the Bangemall Basin (see Brakel & Muhling, 1976). The northern Yilgarn Craton is three-layered and its northern margin is near the Marymia Dome (Fig. 1). The zones of dense lower crust in
B. J. DRUMMOND, R. E. SMITH & R. C. HORWITZ
36 Pannawonica
(a)
(b)
(c)
DISTANCE
200
IN
K!
.WA/BI0-83A
LOMETRES
Fig. 3. Seismic record sections along line FDB of blasts at (a) Pannawonica, to the southeast towards Newman; (b) Tom Price, to the northwest towards Pannawonica (LHS) and to the southeast towards Newman; and (c) Newman, to the northwest towards Pannawonica. The amplitudes of the traces are as recorded, except in (c) where they have been trace-normalized to enlarge the phases beyond 270 km. The travel time curves superimposed on the record sections are for the model in Figure 4.
the centre of segment BC correspond to the seismic phases is as follows. Pj arrivals have an Bangemall and northern Nabberu Basin outcrop apparent velocity of about 5.9km s~ and have travelled directly through the near-surface sediareas. ments. The P phases have a higher apparent velocity (6.3 to 6.7 km s^ ), and have travelled SECTION FDB through near-surface rocks which are underlain Some of the data from line FDB are presented by lower-velocity rocks. Pg arrivals are refracted as record sections in Figure 3. The velocities and through the top of the crystalline basement. The intercept times for the observed phases were P phases are reflected off a lower crustal interscaled from the record sections, and the travel- face, and the P* phases (6.5 km s ) are refracted time equations are listed in Table I. A prelimi- through the lower crust. P arrivals are widenary model derived from these data is shown in angle reflections from the Moho, and Pn arrivals Figure 4. The nomenclature used to describe the are Moho refractions. Not all phases were 1
2
1
1
- 1
M
TABLE I
Pi p Pg * p* Pn 2
#
Pannawonica eastward absent absent tt = D/6.07 s tt = D/6.51 + 1.2 s tt = 0/7.64 + 5.5 s
Seismic travel time equations absent tt = D/6.32 absent tt = D/6.54 + 0.6 s tt = D/6.57 + 1.2 s tt = D/8.0 +6.1 s
tt = D/5.88 - 0.31 s tt = D/6.49 + 0.42 s absent tt=D/6.39+0.91 s absent
absent tt = D/6.76 + 0,1 s absent tt = D/6.77 + 3-1 s tt =0/7.96 + 6.0 s
CRUSTAL STRUCTURE FROM SEISMIC SOUNDING
observed at each shot-point (Table I, Fig. 3), and with differences of up to 2 s in the times for reciprocal pairs of near-surface phases there must be lateral variations in the upper crust. The Pi phase was observed to the east of Tom Price, and the P2 phase both to the east of Tom Price and the west of Newman. Interpretation of the travel-time data shows that the P2 phase results from a shallow, high-velocity layer 0.5 to 1 km deep under Newman and 2 to 3 km deep to the east of Tom Price. These depths are approximately the thicknesses of the Hamersley Basin volcanic and sedimentary rocks in the region (Daniels & MacLeod, 1965; MacLeod & de la Hunty, 1966; de la Hunty, 1965). For 60km west of Tom Price, the high-velocity layer is modelled as 1 km thick, but may be thicker, and beyond 60 km no estimate of the basin thickness is possible because it has a seismic velocity indistinguishable from that of the basement. In Figure 5, the laboratory measurements of seismic velocities and densities in hand specimens from the region are plotted as bars, in their approximate stratigraphic order. Basement under the basin consists of granitic rocks and their metamorphosed equivalents, gneiss and, in places, quartzite. The predominantly basic volcanic Fortescue Group forms the basal unit of the Hamersley Basin, and is overlain by the banded iron formations, mudstone, siltstone, and sandstone of the Hamersley Group. The seismic
5 9 km
6
Fig. 4.
07
s-l
37
velocities are likely to increase with increasing pressure, and therefore with increasing depth of burial, as the microfractures and cracks in the rocks close up, but because this will affect all of the rock types, the relative order of the velocities and densities should be unaltered. The basic volcanics and banded iron formations have the highest velocities and densities, and velocity inversions are likely where they overlie granitic basement. Four zones of burial metamorphism (subgreenschist to greenschist fades) have been recognised in the rocks of the Hamersley Basin (Smith, 1975^; Smith et ai, 1978). The lowest-grade metamorphic zones are recognised on the northern side of the basin. With increasing grade they are now called: pumpellyite-prehnite zone ZI; pumpellyite-epidote zone ZII; pumpellyiteactinolite zone ZIII; and the greenschist facies proper, the actinolite zone ZIV. Rocks of zones ZIV and ZIII occupy most of the southern side of the basin. The intersection of three of the metamorphic zones with the section FDB is shown diagrammatically in Figure 6. The progressive increase in metamorphic grade may explain the observed increase in seismic velocity from west to east in the near-surface high-velocity layer. ZII corresponds to the observed velocity of 6.07km and is indistinguishable from basement; ZIII corresponds to the 6.3 km s ~ ^ velocity, and ZIV to the
6 3 -
6 54 -
6-7km
s-i
7 79 km
6-56
Crustal model along line FDB based on the seismic data in Figure 3 and Table 1.
s-l
38
B. J. DRUMMOND, R. E. SMITH & R. C. HORWITZ
Sedimentary Rocks
lionstones Basic Volcanics Granites Gneisses Quartzites
Fig. 5.
Bar graph of corresponding laboratory measurements of velocities and densities of 49 hand specimens from the survey region. Seismic velocity was measured both parallel and at right angles to the layering in layered samples, and is illustrated in the figure by the split ends of some of the velocity bars.
6.5 to 6.7km s " ^ velocity material. The surface layer of 5 . 9 k m s - ^ between Tom Price and Newman may correspond to the Hamersley Group. No estimate of the velocity of the material below the Hamersley Basin is possible. In the model, 6.07 km s"^ was used; it is the value measured east from Pannawonica and generally agrees with the upper crustal velocities from line GBC (Fig. 2). Subsequent inversion of the P* data gives a lower crustal layer 12 km deep under Pannawonica. About 60 km west of Tom Price, it shallows to 7 km over a maximum distance of 20 km (Fig. 4). The region where the lower crustal layer shallows around Tom Price corresponds to areas of Archaean basement outcrop and doming of the Hamersley Basin rocks. East of Tom Price, the lower crustal layer deepens, possibly reaching 20 km before shallowing again to 13 km at Newman, the depth determined from line GBC. The velocity in the lower crustal layer is 6.55kms-l
Fig. 6.
The 7.8km velocity layer is 31km deep under Pannawonica, and 32 km deep under Newman. Although the depth under Newman agrees well with the depth of the crust/mantle boundary under Newman along profile GBC, the 7.8km velocity is very much less than the upper mantle velocity of 8.3 km s ~ ^ along section GB. Upper mantle P-wave anisotropy of 0.5 km s ~ ^ may therefore be implied in the lithosphere below the Pilbara Craton. Anisotropy of this magnitude has been observed in both oceanic {e.g. Morris et al., 1969) and continental {e.g. Bamford, 1977) lithospheres, and is generally attributed to alignment of olivine crystals. Confirmation of anisotropy in the lithosphere below the Pilbara Craton will have to await more comprehensive analysis of all the available data. An alternative explanation is that the lower 7.8 km s " ^ velocity is confined only to the axis of the Hamersley Basin. Fuchs (1979), in a review of seismic studies of continental lithospheres, reports that lateral changes of several tenths of a kilometre per second can be observed in upper mantle velocities from regions only several tens of kilometres apart. The lateral velocity heterogeneity in the Pilbara Craton is unHkely to be due to high present-day crustal temperatures in the Hamersley Basin, because heat flow in the region is uniformly low (Cull & Denham, 1979). However, it may be due to lower crustal and upper mantle chemical inhomogeneities which pinch out towards the eastern edge of the basin. Drummond (1979a) suggested a tectonic model for the region based on crustal thicknesses and dips in which he inferred that the Pilbara Craton was tilted to the south. The dip of 0.8° on the base of the crust is similar to the dip on Hamersley Basin remnants in the north of the Pilbara Block and represents a subsidence of at least 5 km along the southern boundary of the craton. However, the presence
Diagrammatic cross-section along line FDB illustrating the correspondence of the increase in metamorphic grade to the increase in near-surface seismic velocity. The metamorphic grades in the figure are represented as: ZII pumpellyite-epidote zone; ZIII pumpellyite-actinolite zone; and ZIV actinolite zone.
CRUSTAL STRUCTURE FROM SEISMIC SOUNDING
of Archaean inliers within the Hamersley Basin, which correspond to the shallowing of the lower crustal layer near Tom Price in Figure 4, the sedimentary thicknesses (Horwitz & Smith, 1978, Figs. 6 & 7), and the low-grade burial metamorphism of the Hamersley Basin rocks all suggest that the subsidence was much greater, especially along the southern margin of the craton, and that the crust has rebounded. A maximum rebound, based on metamorphic grades, of 10 to 12 km to the west of Tom Price, and perhaps 5 to 7 km near Newman, has been estimated. The 7.8 km s - 1 velocity may therefore represent lower crustal rocks depressed sufficiently during the Proterozoic for partial melting to have occurred. The partial melting could account for the acid Woongarra Volcanics (Trendall, 1975). The remaining lower crustal rocks would then be anomalously dense, giving rise to seismic velocities of 7.8 km s " 1 . As well, the lower crustal velocity (6.55 km s _ 1 ) along profile FDB is slightly higher than to the northeast under the eastern Pilbara Block (6.4km s _ 1 ) . This probably results from a higher metamorphic grade caused by the depression of the material into a higher temperature and pressure regime. DISCUSSION These studies have considered mainly the Pilbara Craton and its Proterozoic cover rocks, and we are therefore able to suggest a possible scenario of events for the tectonic evolution of the Pilbara Craton from the end of the Archaean. We have not tested all aspects of the scenario and propose it mainly as a basis for further discussion. The movements associated with the vertical tectonics interpreted along Section FDB are approximately half the thickness of the present-day crust in the Pilbara Craton. The forces required to produce such movements are unlikely to have originated within the crust, and a high heat flux is suggested as the factor influencing the creation of the Hamersley Basin. High heat flux would have caused partial melting of the upper mantle, producing the basalts of the Fortescue Group. The basalts spilled out over the Archaean basement, which, weakened by the high heat flux, would have subsided; the metamorphic pattern suggests that the subsidence was greater in the south than in the north. The Hamersley Basin sedimentary and volcanic sequences accumulated and were progressively metamorphosed as they were in turn depressed and buried. The rocks of the lower crust would have been depressed sufficiently to have partially melted, giving rise to acid volcanism. With the passing of the thermal event,
39
basin subsidence ceased. The crust, depressed into the more dense mantle, began to rise as local isostatic compensation was replaced by lowerstress regional compensation. The uplift would not be by the full extent of the original subsidence because of the redistribution of mass from the mantle onto the top of the crust, and the ability of the cooler and therefore stronger crust to support the load. The balance of the subsidence is represented by the 5 km of tilt measured by the seismic profiling. Models for the evolution of the region must reconcile major differences in the geology and structure between the Pilbara and Yilgarn Block. Gee (1979c) listed four of them: (i) The western Yilgarn Block has large areas of high-grade Archaean gneiss terrain—no similar rocks have been found in the Pilbara Block; (ii) granites form the bulk of the rocks in the crust of both the Pilbara and Yilgarn Blocks, but all reliable dates for those of the Yilgarn Block fall in the range 2600 to 2700 Ma, while most of those of the Pilbara Block have ages greater than 3000 Ma; (iii) the clastic sequences within the greenstone belts of the Pilbara Block are more mature than those of the Yilgarn Block, implying to Gee (1979c) that the Pilbara Block developed more toward crustal stability than the Yilgarn Block (such stability also being evident from the broad, simple burial-metamorphic pattern: see Smith, 19756) and (iv) the Pilbara Block has nearcircular granite batholiths, but those of the Yilgarn Block have distinct linear trends. To these four points can now be added two more, which may be related. Firstly, the crust of the Pilbara Craton is 28 to 33 km thick, but that of the Yilgarn Craton is more than 50 km thick in the north (Fig. 2) and 40km thick in the south (Mathur, 1974). The crust in the Yilgarn Craton is therefore much thicker than the crust in the Pilbara Craton. Secondly, excluding the rocks of the Hamersley Basin, the Pilbara Craton is twolayered (Figs 2 & 4), but three layers are present in the Yilgarn Craton (Fig. 2, and Mathur, 1974). A third, thin, lower crustal layer can be used to explain anomalous amplitudes of some seismic phases immediately to the north of Newman, but. it appears to be a local feature and not of significance to the structure of the Pilbara Craton as a whole. The difference in the crustal thickness, especially in the northern Yilgarn Craton, is due mainly to the thick, high-velocity (7 km s _ 1 ) lower crustal layer. The lower crustal layer of Mathur (1974) in the southern Yilgarn Craton has a higher seismic velocity (7.4km s" 1 ), and in the southwestern Yilgarn Craton is very close to the surface. Glikson & Lambert (1976) proposed that
B. J . D R U M M O N D , R . E . S M I T H & R . C .
40
it was formed by the injection of basic m a g m a during the break u p of Gondwanaland in the Phanerozoic. However, D r u m m o n d (1979a) suggested that the layering within the cratons represents the metamorphic effects on a crust of broadly uniform acid to intermediate chemical composition so that the third crustal layer may be a metamorphic effect and not a change in chemistry. Certainly any tectonic model which proposes the injection of magma, whether basic or acidic, into the base of the crust has to inject it selectively at the base of the Yilgarn Craton and not at the base of the Pilbara Craton, which, f r o m palaeomagnetic data (McElhinny & McWilliams, 1977; Embleton, 1978), was in its present position relative to the Yilgarn Block f r o m about 2400 M a ago.
HORWITZ
ACKNOWLEDGMENTS W e acknowledge the invaluable assistance of our colleagues at the Bureau of Mineral Resources, Geology a n d Geophysics (BMR), the Research School of E a r t h Sciences at the Australian National University (ANU) in Canberra, and the C S I R O Division of Mineralogy in Perth. K. J. Muirhead, R. Arculus, a n d D. M . Finlayson critically read the manuscript. The seismic interpretation was p e r f o r m e d while B J D was on full-time study leave at the A N U under the Australian Public Service P o s t g r a d u a t e Scholarship Scheme. This work is published with the permission of the Director of the B M R . J a n e Prowse typed the manuscript a n d Gill Clark drew the figures.
REFERENCES DRUMMOND, B. J., 1979A: A crustal profile across the Archaean Pilbara and northern Yilgarn Cratons, northwest Australia. BMR J. Aust. Geol. Geo-
ARRIENS, P. A., 1971: The Archaean geochronology of Australia. Spec. Pubis geol. Soc. Aust., 3, 11-23. BAMFORD, D., 1977: Pn velocity anisotropy in a continental upper mantle. Geophys. J. R. astr. Soc., 49, 29-48. BMR, 1975: Gravity map of Australia, 1:5000000. Bur. miner. Resour. Geol. Geophys. Aust., Canberra. BRAKEL, A . T . , & MUHLING, P . C . , 1976: S t r a t i g r a p h y ,
sedimentation and structure in the western and central part of the Bangemall Basin, Western Australia. Ann. Rep. geol. Surv. West. Aust. for 1975, 70-79.
CULL, J. P., & DENHAM, D., 1979: Regional variations
in Australian heat flow. BMR J. Aust. Geol. Geophys., 4, 1-13. DANIELS, J. L., 1966: The Proterozoic geology of the North-West Division of Western Australia. Proc. Australas. Inst. Min. Metall., 219, 17-26. DANIELS, J. L., 1975: Gascoyne Province; in Geology of Western Australia. Mem. geol. Surv. West. Aust., 2,107-114. DANIELS, J . L . , & MACLEOD, W . N . , 1965: N e w m a n ,
W.A.—1:250000 Geological Series. Explan. Notes geol. Surv. West. Aust., SF50-10. DE LAETER, J. R., 1976: Rb-Sr whole-rock and mineral ages from the Gascoyne Province. Ann. Rep. geol. Surv. West. Aust. for 1975, 126-130. DE LAETER, J . R . , & BLOCKLEY, J . G . , 1972: G r a n i t e
ages within the Archaean Pilbara Block, Western Australia. J. geol. Soc. Aust., 19, 363-370. DE LAETER, J . R . , LEWIS, J . D . , & BLOCKLEY, J , G . ,
1975: Granite ages within the Shaw Batholith of the Pilbara Block. Ann. Rep. geol. Surv. West. Aust. , for 1974, 73-79. DE LA HUNTY, L. E., 1965: Mount Bruce, W. A. —1:250000 Geological Series. Explan. Notes, geol. Surv. West. Aust., SF50-11. DOUST, Q., 1975: Economic implications of the geology of the Duck Creek Syncline, Ashburton Trough, -Western Australia. Abstr. First Aust: geol. Conv. Prot. Geol., geol. Soc. Aust., 11.
phys.,
4, 171-180.
, 19796: Pilbara Crustal Survey, 1977: Operational Report. Rec. Bur. miner. Resour. Geol. Geophys. Aust., 1979/54 [unpublished]. , 1979c: Structural Relations Between the Archaean Pilbara and Yilgarn Cratons, Western Australia, from Deep Seismic Sounding. MSc. Thesis, Aust. Nat. Univ. [unpublished]. EMBLETON, B. J. J., 1978: The palaeomagnetism of 2400 m.y. old rocks from the Australian Pilbara Craton and its relation to Archaean-Proterozoic tectonics. Precamb. Res., 6, 275-291. FRASER, A. R., 1976: Gravity provinces and their nomenclature. BMR J. Aust. Geol. Geophys., 1, 350-352. FUCHS, K., 1979: Structure, physical properties, and lateral heterogeneities of the subcrustM lithosphere from long-range deep seismic sounding observations on continents. Tectonophys., 56, 1-15. GEE, R. D. (COMPILER), 1979a: Geological map of Western Australia, 1979, 1:2500000. Geol. Surv. West. Aust., Perth. , 19796: The geology of the Peak Hill area. Ann. Rep. geol. Surv. West. Aust. for 1978, 55-62. , 1979c: Structure and tectonic style of the Western Australian shield. Tectonophys., 58, 327-369. GLIKSON, A. Y., & LAMBERT, I. B., 1976: Vertical zona-
tion and petrogenesis of the early Precambrian crust in Western Australia. Tectonophys., 30, 55-89. HALL, W . D . M . , & GOODE, A . D . T . , 1978: T h e Early
Proterozoic Nabberu Basin and associated iron formations of Western Australia. Precamb. Res., 7, 129-184. HORWITZ, R . C . , & SMITH, R . E . , 1978: Bridging the
Yilgarn and Pilbara Blocks, Western Australia. Precamb.
Res.,
6, 293-322.
CRUSTAL STRUCTURE FROM SEISMIC SOUNDING MACLEOD, W . N . , & DE LA H U N T Y , L . E . , 1 9 6 6 : R o y
Hill, Western Australia—1 :250000 Geological Series. Explan. Notes geol. Surv. West. Aust., SF50-12. 1 9 7 4 : Crustal structure in southwest Australia from seismic and gravity data. Tectonophys.,24, 151-182.
MATHUR, S . P . ,
MCELHINNY, M .
W.,
& MCWILLIAMS, M . O . ,
Precambrian geodynamics—a view. Tectonophys., 40, 137-159.
1977:
palaeomagnetic
MORRIS, G . B . , RAITT, R . W . , & SHOR, G . G . ,
1969:
Velocity anisotropy and delay-time maps of the mantle near Hawaii. J. geophys. Res., 74, 4300-4316. MOTA, L., 1954: Determination of dips and depths of geological layers by the seismic refraction method. Geophys., 19, 242-254. OVERSBY, V. M., 1975: Lead isotopic systematics and ages of Archaean acid intrusives in the KalgoorlieNorseman area, Western Australia. Oeochim. cosmochim. Acta, 39, 1107-1125. _ _ _ , 1976: Isotopic ages and geochemistry of Archaean acid igneous rocks from the Pilbara, Western Australia. Geochim. cosmochim. Acta, 40, 817-829. PIDGEON, R. T., 1978a: 3450M.y. old volcanics in the Archaean layered greenstone succession of the Pilbara Block, Western Australia. Earth planet. Sci. Lett., 37, 421-428. ,19786: Geochronological investigations of granite batholiths of the Archaean granite-green-
41
stone terrain of the Pilbara Block, Western Australia; in Smith, I. E. M.f & Williams, J. G. (Eds) Proceedings of the 1978 Archaean Geochemistry Conference, 360-362. University of Toronto. 1978c: Big Stubby and the early history of the earth. Open-File Rep. U.S. geol. Surv., 78-701, 334-335. SMITH, R. E., 1975a: Exploration guides for Keweenawan-type copper deposits in the Proterozoic Fortescue Group of Western Australia, Preliminary Report. Rep. CSIRO miner. Res. Lab., FP 8. , 19756: Metamorphism of the Proterozoic Fortescue Group, Western Australia—a reconnaissance study. Rep. CSIRO miner. Res. Lab., FP 9. SMITH, R . E . , GREEN, A . A . , ROBERTS, G . , & HONEY,
F. R., 1978: Use of Landsat-1 imagery in exploration for Keweenawan-type copper deposits. Remote Sens. Environment, 7, 129-144. SMITH, R . E., & HORWITZ, R . C., 1975: Structural syn-
thesis for the deposition of the Wyloo and Bangemall Groups. Abstr. First Aust. geol. Conv. Prot. Geol., geol. Soc. Aust., 77. TRENDALL, A. F., 1975: Main areas of Proterozoic sedimentary rocks, Hamersley Basin; in Geology of Western Australia. Mem. geol. Surv. West. Aust., 2, 119-143. WILLIAMS, S . J . , ELIAS, M . , & DE LAETER, J . R . , 1 9 7 8 :
Geochronology and evolution of the eastern Gascoyne Province and the adjacent Yilgarn Block. Ann. Rep. geol. Surv. West. A ust. for 1977, 50-56.
CRUSTAL DEVELOPMENT IN THE ARCHAEAN YILGARN BLOCK, WESTERN AUSTRALIA R. D. Gee, J. L. Baxter, S. A. Wilde, & I. R. Williams
Geological Survey of Western Australia, Mineral House, 66 Adelaide Terrace, Perth, Western Australia 6000 ABSTRACT
The remarkable topographic flatness of the Yilgarn Block appears to be the result of exhumation of a Proterozoic unconformity, suggesting that the present crustal layering is an inherited Archaean feature. Seismic data indicate a two-layered sialic crust 32 km thick in the Kalgoorlie region. Considerably thicker crust (ca 45 km) to the west is due to an additional lower, denser layer consistent with mafic granulite. The thicker crust corresponds with surface distribution of high-grade (amphibolite to granulite fades) paragneiss and orthogneiss, forming an arc around the western and northwestern margins of the Yilgarn Block. We refer to this area as the Western Gneiss Terrain, and contrast it with the granite-greenstone terrains of the Murchison, Southern Cross and Eastern Goldfields Provinces that are underlain by thinner sialic crust. The Western Gneiss Terrain consists of repeatedly deformed and metamorphosed sediments, intruded by mafic and ultramafic rocks, and infolded with sheets of orthogneiss. The sedimentary lithofacies is commonly a shallow-water sequence indicating stable-shelf conditions. The rocks date back at least to 3.3 b.y., and are interpreted as basement upon which the greenstone sequences to the east were deposited. The three granite-greenstone provinces apparently represent fundamental depositories in which typical layered greenstone sequences were laid down in broad down warps during a unique event. They continued to develop synchronously with the domal uprise of granitic magma. The rising granite diapirs restricted late-stage sedimentation to localized basins and fault-bounded troughs, and ultimately pinched the greenstone sequences into narrow belts. Sheets of discordant granite, less than 5 km thick, were subsequently emplaced throughout the Yilgarn Block. Both forms of granite were derived by remobilization of a pre-existing gneissic basement: there was no rem el ting or assimilation of greenstone to produce granite and little introduction of sialic material to the crust. Available evidence suggests that greenstone deposition and granite emplacement took place over a narrow time span and were related to a major tectono-thermal event between 2.8 and 2.6 b.y. ago, the final stages of which led to the cratonization of the Yilgarn Block.
INTRODUCTION The Yilgarn Block, one of the largest intact segments of Archaean crust on the face of the Earth, has attracted considerable attention from students of the Archaean in the formulation of their models of crustal evolution. Yet because of its large size and paucity of outcrop, it is not well understood, and many models suffer because they have been unable to draw upon sufficient indigenous data, or because they have incorporated concepts more relevant to Archaean terrains elsewhere. Systematic regional mapping of the Yilgarn Block by the Geological Survey of Western Australia has just been completed and considerable regional data are now available. This symposium follows that achievement too closely for all data to be assessed fully and incorporated here, but Spec. Pubis geol. Soc. Aust., 7 (1981)
coupled with the many recent detailed studies of critical regions by several research institutes, it does allow examination of the Yilgarn Block as a whole. This paper outlines the crustal evolution of the Yilgarn Block, using the regional data supported by other recent studies: it is not intended to be an exhaustive review of the earlier literature. MORPHOLOGICAL FEATURES OF THE YILGARN BLOCK Post-Archaean features defining the present, roughly rectangular shape of the Yilgarn Block are the Darling Fault to the west, unmetamorphosed and flat-lying Phanerozoic sediments to the east, Proterozoic mobile belts to the south and northwest, and unconformable Proterozoic sediments to the northeast.
44
R. D. GEE, J. L. BAXTER, S. A. WILDE & L R. WILLIAMS
A remarkable feature of the Yilgarn Block is the extreme flatness. Proterozoic sedimentary rocks around and within the margin overlie the basement on smooth, flat unconformities, implying that the flat geomorphic surface is an exhumed Proterozoic peneplain (Fairbridge & Finkl, 1980). The precise age of the peneplain is uncertain, but its undoubted antiquity, predating the earliest known Proterozoic orogenic events, suggests that the Yilgarn Block has not undergone significant regional tilting since its stabilization about 2.5 b.y. ago, contrary to the suggestion of Glikson & Lambert (1976, p. 62). A corollary is that the attitude of its crustal layering is an inherited Archaean feature. CRUSTAL STRUCTURE Seismic reflection and refraction traverses (Mathur et al., 1977) indicate a two-layered sialic crust consisting of granite (d 2.78 gm c m - 3 , Vp 6.13km sec- 1 ) and an underlying granulite layer (d 2.94, Vp 6.72). Together these are 14km thick on the western margin, where gneissic terrain is now exposed, but thicken to 32 km in the Kalgoorlie area, a region composed of granitegreenstone terrain. However, the crust in the west also contains a lowermost layer of probable mafic granulite (d 3.0, Vp 7.49) giving a total thickness of 46 km. A recent seismic profile (Drummond, 1979) that intersects the northern margin of the Yilgarn Block, an area also underlain by gneiss, indicates crustal layering of similar nature and thickness to that along the western margin. Provisional conclusions are that the crust thins inward from the marginal gneiss terrains, and that the crust beneath the gneiss terrains is fundamentally different to that under the greenstone terrains. Since we interpret this as an Archaean feature, it further suggests an original basinal structure for the Yilgarn Block. REGIONAL SUBDIVISIONS IN THE YILGARN BLOCK An understanding of the crustal development must take into account the regional distribution of gneiss and greenstone terrains, metamorphic patterns, deformation styles, granite emplacement, and available geochronology. A review by de Laeter et al. (1981) of the Archaean geochronology, and the considerable number of isotopic analyses in the last decade, since the pioneering work of Arrlens (1971), provides the chronological framework for this paper. A major tectono-thermal event at 2.7 to 2.6b.y. was responsible for emplacing vast masses of granite within the whole extent of the Yilgarn Block, a
phenomenon which more than any other obscures the earlier geological history. Ignoring for the moment the granites and their associated thermal and structural effects, a major subdivision is evident (Fig. 1A) between predominantly gneiss terrain (the Western Gneiss Terrain) forming an arc around the western periphery, and three large areas containing greenstone belts (the Murchison, Southern Cross and Eastern Goldfields Provinces). This scheme modifies that proposed by Williams (1974), since each greenstone province, as now defined, apparently has a unified lithology, structural history, and perhaps even stratigraphy. WESTERN GNEISS TERRAIN Rock Types The dominant rock type in the Western Gneiss Terrain is banded quartz-feldspar-biotite gneiss, locally containing garnet, cordierite and hypersthene. Planar banding on centimetre to metre scale distinguishes it from the emplaced granitic rock so common throughout the Yilgarn Block. Although some banded gneisses were derived from earlier granitic rocks by progressive deformation of porphyritic granite to produce augen gneiss, many represent metamorphosed quartzofeldspathic sediments, whereas others may represent metamorphosed layered migmatite. The most diagnostic feature of the gneiss terrain is the inclusion of belts of schist and gneiss of undoubted sedimentary origin, the major areas of which are shown on Figure 2. These serve to identify four metamorphic belts called Balingup, Chittering, Jimperding and Narryer. Metasedimentary rock types include quartzfeldspar-cordierite-biotite-garnet granofels after greywacke; quartz-plagioclase-muscovite-sillimanite (or kyanite or andalusite) pelitic schist; muscovite-sillimanite-quartz psammitic schist; orthoquartzite with preserved cross bedding; banded quartz-magnetite-amphibole (-orthopyroxene) rocks representing BIF; diopsideepidote-garnet-hornblende-plagioclase-microcline gneiss representing calcareous sediments; and sillimanite-garnet-muscovite metaconglomerate. Small bodies of mafic and ultramafic rock are distributed throughout the paragneiss. Highgrade metamorphism and deformation have destroyed most original features, although some bodies appear to have been altered intrusions: W. R. Morgan (pers. comm.) has described rhythmic layering in a metamorphosed harzburgite-lherzolite-anorthosite body at Kondinin; Davidson (1968) interpreted mafic granulite at Quairading as a differentiated layered mafic in-
45
CRUSTAL DEVELOPMENT IN YILGARN BLOCK A
B
200 km l||p] Norseman «—' Wiluna Belt
GSWA
18591
Fig, 1. A: Regional subdivision of the Yilgarn Block. B: Metamorphic patterns in the Yilgarn Block. Low-grade domain varies from prehnite-pumpellyite to upper-amphibolite facies, but is predominantly greenschist or low-amphibolite facies. Medium-grade domain is mid-amphibolite to amphibolite-granulite transition facies. Medium- and high-grade domain varies from mid-amphibolite to granulite facies.
trusion; and similar layered mafic-ultramafic bodies occur in the Narryer Metamorphic Belt (Williams et al., in press). These bodies, which are metamorphosed to granulite facies, contain cumulate chromite and prominent anorthosite layers. No discrete anorthosite bodies have yet been encountered. Evidence for large thicknesses of ultramafic (komatiitic), mafic or felsic volcanics is lacking, and there is no direct evidence of any volcanogenic rock in the gneiss terrain. The metasedimentary terrains therefore contrast markedly with the lithologies in the greenstone belts, and appear to represent a rather conventional marineshelf sequence. Possible sedimentary lithofacies variations are represented by contrasts between the Jimperding and Chittering Metamorphic Belts. A BIFquartzite-psammite facies suggesting shallowwater deposition is identified in the Jimperding Belt, whereas a greywacke trough facies is likely in the Chittering Metamorphic Belt. The separation of lithofacies in thfse two belts is also reflected in contrasting metamorphic facies and structural styles. Both the shelf and trough litho-
facies seem to be present in the Balingup Metamorphic Belt, but their spatial separation is not as obvious as in the Jimperding and Chittering Metamorphic Belts. The lithofacies in the Narryer Metamorphic Belt is a shallow-water sequence, represented by coarse conglomerate, quartzite, pelite and calcareous sediments. Structures in the Gneiss Terrain The structure of the gneiss terrain is complex and difficult to interpret, and neither a coherent regional structure nor a consistent structural sequence has been identified. Most orthogneiss appears to post-date paragneiss, although the sedimentary nature of portions of the metamorphic belts requires the presence of a preexisting granitic basement. In the extensive banded quartz-feldspar-biotite gneiss, the gneissic foliation has been impressed upon a rock that was already banded. This gneissic foliation is axial-planar to near-isoclinal folds of the banding so that the banding is independent of, and earlier than the first recognizable penetrative deformation. It may be an inherited sedimentary banding, modified by varying
R. D. GEE, J. L. BAXTER, S. A. W I L D E & I. R. W I L L I A M S
46
GREENSTONE
BELTS
r + ^+1 "POST-TECTONIC" D I S C O R D A N T G RANITE 200 km
"SYNTECTONIC" DOMAL GRANITE I ^ M
B A N D E D GNEISS and MIGMATITE PREDOMINANTLY METASEDIMENT
GSWA 18592 Fig. 2.
Interpretation of relationships between greenstone belts and granitic and gneissic rocks in the Yilgarn Block. Narryer, Jimperding, Chittering and Balingup Metamorphic Belts are shown. Trends in domal granites and banded gneisses are diagrammatic but follow known foliation trends. Greenstone belts referred to in the text are shown by numbers: 1-Jack Hills, 2-Tallering, 3-Koolanooka, 4-Wongan Hills, 5 Saddleback. Regional centres are: T-Toodyay, S-Southern Cross, K-Kalgoorlie, N-Norseman,. L—Leonora, W—Wiluna, M—Meekatharra.
47 sions, to form complexes which record the oldest published Rb/Sr whole-rock dates (about 3.0-2.9b.y.) in the Yilgarn Block. Geological evidence requires that these gneiss-forming events predate the widespread tectono-thermal event at 2.6b.y., to which the evolution of the greenstone terrains is closely related. The U/Pb zircon dates of ca 3.3 b.y. from gneiss and quartzite at Toodyay (Nieuwland & Compston, 1981) show that these gneiss-forming processes had started at least 500m.y. before the major tectono-thermal event in the Yilgarn Block. We interpret the gneiss terrain as exposed samples of deeper crustal material, possibly representing the granulite layer (density 2.94) referred to previously, and thus a basement to the greenstone terrains. This interpretation accords with that of Rutland (1976, p. 189), but differs from that of Glikson & Lambert (1976) who believed that the greenstone and gneiss -terrains were "the manifestation at different crustal levels of the very same igneous and metamorphic events". There are five areas, whose importance has only recently become apparent (Jack Hills, 1; Tallering, 2; Koolanooka, 3; Wongan Hills, 4; and Saddleback, 5—the numbers are shown on Fig. "2), where greenstone belts come into proximity or contact with gneiss terrain. These areas therefore should present critical evidence bearing on the relationship of gneiss to greenstone belts. Experience to date, however, is that contact relations are generally unexposed, tectonized, or obliterated by later granite intrusion. In the Jack Hills greenstone belt (1) (Williams et al. in press), mafic sediments and minor volcanics near to the greenschist-amphibolite transition are in proximity to granulite facies gneisses. A similar situation exists at Tallering (2) (Baxter, 1971), although the metamorphic contrasts are not as great. The enigmatic Koolanooka (3) (Baxter & Lipple, 1978) and Wongan Hills (4) (Carter et al., 1978), sequences show no contrasts in metamorphic grade with adjacent gneiss although there are significant lithological differences. The Saddleback Group (5) (Wilde, 1976), of felsic and mafic volcanics lies well within the Western Gneiss Terrain, yet is metamorphosed only to greenschist facies. It must represent a supracrustal greenstone belt, but most of its contacts are in faulted or intrusive relation with granite, and its affinities with the greenstone sequences to the east are uncertain. Additional clues are revealed elsewhere in the Yilgarn Block where banded gneiss within some of the domal granites, is juxtaposed with greenstone belts. Wherever seen, the contact is a high-
CRUSTAL DEVELOPMENT IN YILGARN BLOCK
degrees of partial melting, or it may be stromatic banding in migmatite, of which the palaeosome could be metasedimentary, and the neosome could register a very early tectono-metamorphic event. In the obviously sedimentary rocks at Toodyay, where BIF and quartzite can be used as structural markers, three phases of folding can be recognized (Wilde & Low, 1978), of which at least the earlier two have penetrative fabrics and predate the intrusion of the 2.6b.y. granites. The earliest folds at Toodyay are thought to be recumbent, the vestiges of which are zones of flat undulatory foliation and bedding (in places overturned), preserved between areas of vertically transposed foliation. Metamorphism Metamorphism in the gneisses is complex and polyphasal, but clearly predates the intrusion of the 2.6 b.y. granites. In general, high-amphibolite facies prevails, but pockets of granulite-facies assemblages are widespread, and are particularly abundant in the Narryer and Jimperding Metamorphic Belts. Granulite assemblages include cordierite-garnet-sillimanite-quartz in pelite, orthopyroxene - clinopyroxene - plagioclase in mafic rock, and sillimanite-microcline-plagioclase-biotite in psammite. PT conditions of 5-6.5kb and 750-840°C have been calculated from electron-microprobe data on the cordieritegarnet pair by D. Blight (pers. comm. 1979), using the calibration of Currie (1971), The widespread presence of cordierite, and the usual association of andalusite with sillimanite indicates moderate pressure over most of the gneiss terrain. In the high-grade metamorphic terrains in the Narryer Metamorphic Belt, sillimanite and cordierite are present and andalusite is absent. The Chittering Metamorphic Belt is unusual in that kyanite is associated with sillimanite and staurolite, and cordierite is absent. This is therefore a belt of higher pressure adjacent to the Jimperding Metamorphic Belt of low pressure. Together with the contrasts in their original lithofacies, the two belts reveal themselves as paired baric belts not unlike those in the younger part of the geological column, perhaps pointing to a uniformitarian aspect in the development of the gneiss terrain (Gee, \919b). Relationship of Gneiss and Greenstone Terrains In summary, the Western Gneiss Terrain contains high-grade metasedimentary rocks originally of ensialic shelf lithofacies, devoid of volcanics, and quite unlike the Archaean greenstone sequences. Paragneisses are interfoliated with migmatite, orthogneiss and mafic intru-
y
48
R. D. GEE, J. L. BAXTER, S. A. WILDE & I. R. WILLIAMS
strain zone, commonly with mylonite. Where banding in the gneiss is discordant to the contact, it is also discordant to the lineation and foliation developed in both gneiss and greenstone sequences during emplacement of granite diapirs. Although the origins of the banding and the gneiss itself are uncertain, a metamorphic (or plutonic, or migmatitic) event that is registered in the banded gneiss is not found in the greenstone sequences. This argument was used by Archibald & Bettenay (1977) to postulate a pre-greenstone sialic basement in the Eastern Goldfields Province, and is supported by observed relationships in the central and northern part of the Yilgarn Block where gneiss belts are present. The banded gneisses therefore emerge as important rocks in the understanding of the evolution of the Yilgarn Block, and as they form an integral part of the vast granitic areas, it is convenient to treat the granitic rocks before discussing the greenstone terrains.
GRANITES OF THE YILGARN BLOCK Granites (sensu la to in this paper, except where indicated otherwise) account for about 70% of the crustal exposure of the Yilgarn Block. Neglecting the enclaves of paragneiss within granitic gneiss, it is evident that by far the most common compositional type is in the modal field of adamellite and granodiorite (Libby, 1979). Tonalite, trondhjemite, quartz monzonite, syenite and granite (sensu stricto) are volumetrically unimportant. This contrasts markedly with many other greenstone-granite terrains where more basic plutonic rocks predominate, but its significance cannot yet be interpreted. Classifications based on fabric, particularly the mapping of magmatic or tectonic fabrics, have proved to be more useful, and serve to identify syn-kinematic and post-kinematic types. However, this approach may be unsatisfactory where batholiths show transitional fabrics. It is also confusing where more than one tectono-thermal event is present (as in the banded gneiss), or where later penetrative events affect the younger granites. The classification used here identifies three types of granite which could be termed pre-, synand post-tectonic when related to the 2.6b.y. tectono-thermal event, but which are here given descriptive terms. Banded Granitic Gneiss On a simple "granite-greenstone" classification, these rocks clearly group with the granites. However as they have already been described in
previous sections, no further description is warranted. Tectonically Emplaced Domal Granite The greenstone belts in the Murchison, Southern Cross and Eastern Goldfields Provinces outline a multitude of discrete or coalescing ovoid plutons, reminiscent of the tectonic style of the Pilbara (Hickman, 1981). In most bodies strong linear and planar fabrics are present in the outer parts, and these fabrics are congruent and presumably cognate with those in the enclosing greenstone sequences. The contacts are usually high-strain zones, mostly without evidence of magmatic intrusion. The intensity of the tectonite fabric decreases toward the core of the plutons, progressively becoming a granoblastic fabric without dimensional preferred orientation. In the centre of the larger masses, an allotriomorphic granular texture may be developed. Domal emplacement into greenstone sequences by crystal-plastic strain at elevated temperatures is envisaged. The domes spread and squash the greenstone sequences into distinctive stellate or 4 'hour-glass" shapes, while retaining essentially concordant relationships. Some passive melting may have occurred in the cores of larger domes although subsequent recrystallization may have taken place locally. Important relationships can be observed where domal granites are emplaced alongside major areas of banded gneiss. Direct evidence that the domal granites are derived by remobilization of banded gneiss is given by Elias (in press) who describes a general progression from banded gneiss with migmatite leucosomes, through granoblastic granite with ghost banding, to nebulitic allotriomorphic-textured granite. Discordant Intrusive Granite Discordant intrusive granite forms extensive bodies of magmatic-textured granodiorite, adamellite and granite (s.s-.) that have an obvious intrusive relationship with the domal granite, with the banded gneiss, and with the greenstone sequences. Contacts are sharp but irregular, the intruded rocks bearing evidence of rupture, often at a high angle to the tectonite fabric. Agmatite (migmatite consisting of enclosed blocks of gneiss in magmatic granite) characterizes many contacts (Gee, 1979a). The most extensive areas of discordant granite are in the southwestern portion of the Yilgarn Block, more or less coincident with the rise in the Bouguer anomaly in this area, and the granites must therefore be shallow sheets. Calculations from measurements of heat flow and radioactive heat generation (Jaeger, 1970; Lambert, 1971)
CRUSTAL DEVELOPMENT IN YILGARN BLOCK
suggest a thickness of about 5 km for the granites. Large areas of intrusive granite also occur as corridors between the three main greenstone provinces, and in the north-east part of the Eastern Goldfields Province, where fractionated granite (s.s.) is common (Bunting & Williams, 1979). The granites contain high K 0 / N a 0 (1.3), U (up to 20 ppm), Th (up to 90ppm), fluorine (0.2%) and low K/Rb (130). Large masses of coarse porphyritic or seriatetextured adamellite and granite (s.s.) are included in this category of discordant intrusive granite. Phenocrysts are commonly abraded or cracked, and aligned in magmatic flowplanes, and the rocks were probably intruded as a crystal mush. Chevron-style kink folding of banded gneiss is common near some of these discordant granites (Gee, 1979a), and the appearance of magmatictextured leucosomes in pressure-release zones (e.g. fold cores and axial planes), ultimately leading to the formation of agmatite, clearly indicates derivation of magma from gneiss. We see little of the sites of generation of these magmas on a regional scale since the magmas appear to have been quite mobile and intrude crustal levels some distance from the sites of generation. However, one example has been described by Baxter & Lipple (1978). The emplacement of the domal granites and the intrusion of the discordant granites took place over a narrow time-range between 2.7 and 2.6 b.y. ago and records a remarkably widespread and intense thermal event in which a granitic substrate was remobilized and greenstone sequences were deformed and metamorphosed. To help understand the processes involved in the accumulation of the greenstone sequences, we work back through the metamorphism and deformation of the rocks. 2
2
THE 2.6B.Y. TECTONO-THERMAL EVENT IN THE GREENSTONE SEQUENCES Regional Metamorphic Patterns Regional metamorphic patterns of the greenstone sequences are shown in Figure IB. A lowgrade domain which extends over a large portion of the Yilgarn Block is characterized mainly by greenschist and low-amphibolite facies. In the Eastern Goldfields Province, Binns et at. (1976) demonstrated that variations in metamorphic grade within greenstone belts are related to structural and hence stratigraphic level. Temperatures of up to 600 °C were reached at the base of sequences that are at least 15 km thick. Binns et al. (1976) also showed that the metamorphism was imposed quickly and generally contempor-
49 aneously with deformation, but in places outlasted the bulk deformation of the greenstone belts. Medium-grade metamorphism affects some major greenstone belts, but is largely confined to the small greenstone remnants in the central and southern part of the Yilgarn Block (Fig. IB). Mid-amphibolite to amphibolite-granulite facies with temperatures up to 750 °C and pressures equivalent to 17 km of crustal overburden are recorded here. Metamorphism in greenstone sequences is generally a single-phase event of low-pressure type, and indicates very high geothermal gradients.in excess of 40 °C/km. Blight (1978) presents evidence that geothermal gradients in the west, at Wongan Hills, may have been as high as 70°C/km. This suggests that the two metamorphic domains reflect a continuous gradient from east to west, which corresponds with an increase in the Bouguer anomaly, a progressive decrease in the thickness of the upper crust, a thickening of the total crust, and a diminution in the proportion of greenstone to granite. Small greenstone remnants in the mediumgrade domain (Fig. IB) tend to be encased in gneiss, intrusion of the discordant granite seemingly playing no significant role in the fragmentation of the greenstone sequences. Although metamorphism in these remnants is relatively high-grade, they reveal no evidence of anatexis, and contacts with the gneisses are co-planar with the metamorphic fabric. The higher metamorphic gradient in the southern part of the Yilgarn Block may be due to higher geothermal gradients in an area where originally thin greenstone cover developed on a thicker sialic crust. It appears unnecessary to regard these greenstone remnants as 'root'' zones of once larger greenstone belts that became tectonically elevated to higher crustal levels. 4
Deformation of the Greenstone Belts Except for some greenstone belts dismembered by intrusion, most are revealed as crumpled synforms interstitial to the domal granites. Diapiric movement of a granitic substrate, with compensatory sinking of greenstone sequences, must be considered as the most important deformation mechanism in the Yilgarn Block. Patterns attributed to diapirism are well shown in the Murchison and Southern Cross Provinces. Regional upright folds which are clearly outlined by gross lithological layering have variable axial plunges and curvilinear axial-plane traces. Evidence of repeated deformation is common, but is best explained as progressive deformation during
50
R. D. GEE, J: L. BAXTER, S. A. WILDE & I. R. WILLIAMS
diapir growth, and interaction between adjacent diapirs. In the Eastern Goldfields Province, a distinct north-northwest-trending linear element modifies the diapiric style of deformation. The linearity is expressed by chains of elongate plutons, and numerous tectonic lineaments that can be followed for distances in excess of 300 km. The lineaments are expressed as lines of mis-match of regional structure, ductile shear zones, strike faults and lines of dunite intrusions. They also appear to have controlled the distribution of the domal granites and the orientation of regional folds. No consistent direction of movement has been identified, although substantial strike-slip movements are likely. Gee (19796) has noted that folds and faults in parts of the Eastern Goldfields Province demonstrate a primitive style of basement-cover tectonics whereby contrasting styles of deformation in the greenstone belts seem to reflect the differing behaviour of basement blocks. These lineaments may have opened up as a series of en echelon, mantle-tapping fractures during the accumulation of the greenstone sequences and then developed into crustal sutures which became high-strain zones in the granitic basement during later stages of the tectonothermal event. ACCUMULATION OF THE GREENSTONE SEQUENCES Individual greenstone belts consist of thick volcanogenic sequences, all displaying the same distinctive suite of rock types, and all characterized by large-scale layering of groups of genetically related rock types. The groups have been called associations by Williams (1970) and are useful in the analysis of greenstone deposition, as they have environmental and stratigraphic significance. Mafic-Ultramafic Volcanic Associations Tholeiitic basalt forms thick (commonly 5 km) sheets which may extend along strike for several hundred kilometres. Pillowed and non-pillowed basalts are present, but individual flows are difficult to identify (pyroclastic rocks are rare). Consanguineous dolerite and gabbro sills are common. Continuous beds of oxide-silicate fades iron formation within the sequence indicate quiescent periods during volcanism. Volcanism itself is seen as quiet, extensive outpourings on the sea floor, possibly emanating from fissures rather than vents. There are no reliable indicators of the depth of water. Geochemically, the lavas closely resemble modern low-K oceanic tholeiite (Hallberg, 1972). In the Murchison Province
some .tholeiitic sheets are interbedded with a high proportion of tuff and breccia, and may have been extruded into shallow water, or even subaerially. Outpourings of komatiitic basalt and peridotite were of similar character to the tholeiitic outpourings, although it seems that the volcanic piles were neither as extensive, nor as thick (2-3 km). Individual flows of about 10 m thick are recognized by quench textures (Barnes et al., 1974) and by interflow sediments. Ultramafic lapilli tuff is commonly encountered, and in fact many ultramafic bodies, initially thought to be intrusions, and later reinterpreted as flows, are now identified as clastic sedimentary rocks (Bunting & Williams, 1979). They are considered to result from the physical breakdown of ultramafic volcanics in contact with the sea immediately upon extrusion, and the transport of the fine debris by bottom currents. Apart from the concomitance of mafic and ultramafic volcanics in major sheets, no consistent spatial or temporal relationships have been recognized between tholeiite, komatiitic basalt and peridotite. Close interlayering at the individual flow level is absent, and there is a tendency for the komatiitic and tholeiitic rocks to develop separately. Examples of superposition of tholeiite on komatiite and vice versa are known, and contrary to the popular Barberton model (Anhaeusser et al., 1969), there is no preferential occurrence of komatiite at the lowermost stratigraphic levels. Although the bulk of komatiitic and peridotitic volcanism is associated with the major mafic-ultramafic layers, thin peridotite flows (or their proximal clastic debris) are known in felsic volcanic complexes, fluvial conglomerates, and pelagic sediments. Felsic Volcanic Associations Some instances of repeated cyclic mafic-tofelsic intercalations are known (Williams et al.» 1976), but intervals of prolonged and widespread felsic volcanism, with little associated basaltic activity, were normal. Thick extensive blankets of felsic volcanic and related epiclastic sediments could thus build up. Within these major units, which may have extended for hundreds of kilometres, are discrete volcanic complexes up to 1000 km 2 in area, composed of agglomerate, tuff, rare lavas, and sub volcanic intrusives. Some volcanic centres contain ignimbrite and ash-fall tuff and therefore grew from submarine to subaerial complexes. Thin beds of graded ash tuff and chert in otherwise coarse agglomerate are possibly caldera deposits. A distinctive .feature of these complexes is the clear evidence of their explosive, volatile-charged nature.
51 erosional events unrelated to contemporaneous volcanism. Particularly, but not exclusively, at higher stratigraphic levels, sequences of conglomerate, sandstone and shale mark local fluvial or shallow marine basins into which immature epiclastic material was deposited and rapidly buried. Water-worn clasts of all greenstone rock types, plus granite, are represented in the conglomerates. Examples of such sequences occur at Merougil Creek, Kurrawang and Jones Creek in the Eastern Goldfields Province, Ravensthorpe and Diemals in the Southern Cross Province, and at Warriedar in the Murchison Province (Fig. 3). Unconformities can be established or inferred at the base of the sequences, and there is clear or inferred evidence for the emplacement and unroofing of some granite plutons before sedimentation. The most prolific occurrence of granite detritus is at Jones Creek (Marston & Travis, 1976) in unconformable contact with granite. Naldrett & Turner (1977) invoked a local graben model for sedimentation in the Jones Creek area which can be applied to many of these epiclastic sequences. It is worth noting however that granite detritus is also found at a low stratigraphic level within the greenstone sequences (e.g. Murphy Hills, Bunting & Chin, 1979; Gower, 1976; and Meekatharra, Elias, pers. comm.). Such rare examples are important as they establish the emplacement and unroofing of some granite plutons during the build-up of the greenstone sequences. However despite growing evidence of a granitic basement to the Archaean succession, none of the granite detritus mentioned can be taken as direct evidence of it. Another association, rare in the greenstone sequence, comprises cross-bedded orthoquartzite and psammitic schist. These are known only in the Southern Cross Province where they form discontinuous units no more than a hundred metres thick at the lowermost stratigraphic levels. The rocks have been recently found and their importance remains to be evaluated. They appear to be well-sorted, shallow marine sands derived from a silicic source and deposited on a stable platform, and may provide the most direct evidence of a sialic basement to the greenstone sequences. All known contacts with adjacent granites are either tectonic or intrusive, but significantly some contacts show angular discordance of quartzite against banded gneiss. No convincing evidence of a basal unconformity has been found. Stratigraphy and Lithofacies of the Greenstone Sequences Some workers consider that spatial and temporal variations in volcanic and sedimentary
CRUSTAL DEVELOPMENT IN YILGARN BLOCK
Surrounding the larger complexes are extensive waterlain volcanogenic sediments in which proximal and distal facies can be identified. Oligomictic conglomerate, derived from agglomerate, occurs in submarine and perhaps subaerial scree fans peripheral to some complexes. Proximal and distal turbidite structures can be found in the resedimented crystal-lithic tuffaceous greywackes and argillaceous laminites. Away from the volcanic centres, resedimentation was by dilute turbidity currents. Extensive beds of ferruginous chert and iron-formation, at least 100 km in strike length, abound with these volcanogenic sediments. Rare carbonate beds have also been recorded in these sequences. The tectonic environment for this type of felsic volcanism can be summarized as discrete explosive vents on an otherwise quiet, flat sea floor. There is no suggestion of trough sedimentation associated with the volcanism, and water depth need not have been great. Depth of water is difficult to estimate, and recently encountered indicators of shallow water such as carbonatereplaced gypsum (Golding & Walter, 1979), mud cracks (A. T. Brakel, pers. comm. 1979), and pisolitic carbonates are consistent with the facies model outlined above. These extensive marine-type felsic volcanics are predominantly dacitic and do not appear to be calc-alkaline. There is no field evidence of partial melting or wholesale assimilation of greenstone sequences by the granites to generate them. In fact, a recurring theme in the Yilgarn Block is the absence of evidence of magma generation from within the greenstone sequences themselves. Another environmental type of felsic volcanism is the isolated, largely subaerial centre, unrelated to extensive blankets of marine volcanogenic sediments. It is associated with polymictic conglomerate and fluvial or shallow marine sediments of restricted extent. The relatively few known examples all seem to be located in down-warps in the upper part of the greenstone sequence. A number have calc-alkaline affinities, containing andesite, dacite and rhyolite, but the total geochemical data on the felsic volcanics are insufficient to assess their significance. Hallberg et al. (1976) suggest that at least one of these centres (Marda) reflected a 'hot spot" in the crust, over which lavas were produced by partial melting of the 'lower crust". 4
4
Sedimentary Associations In the foregoing discussion, reference has been made to chemogenic and volcanogenic sediments, but other sedimentary facies are evident in the greenstone sequences, representing important
?
52
R. D. GEE, J. L. BAXTER, S. A. WILDE & I. R. WILLIAMS
GSWA 18593
Fig. 3.
Regional distribution of lithofacies and reconstruction of major greenstone basins in the Yilgarn Block. Felsic volcanic complexes are shown in black. The lightly shaded area in the Eastern Goldfields Basin is the NorsemanWiluna Belt, which developed as a graben. It contains a chain of felsic volcanic complexes, strike faults, abundant volcanic peridotite and rare BIF. Stratigraphically high epiclastic sequences are shown by circles; localities referred to in the text are: W—Warriedar, R— Ravensthorpe, D—Diemals, M—Merougil Creek, K—Kurrawang, J—Jones Creek.
facies are so complex and transient that only local stratigraphic sequences can be identified. Others however have proposed stratigraphic sequences extending over hundreds of kilometres. The ambiguity arises from lithofacial and structural complexities in the greenstone belts. But despite portents of possible structural repetition and thickening of stratigraphic sequences (Burke et ah, 1976), we are unaware of evidence to dispel the belief that multilayering of thick mafic and felsic associations is a stratigraphic phenomenon. The regional variation in lithofacies in the Yilgarn Block is shown in Figure 3. Despite widespread tholeiite, the virtual absence of komatiite in the Murchison Province is notable. Komatiitic basalts, with but few peridotites, are present in the southern part of the Southern Cross Pro-
vince, and both komatiitic basalt and peridotite abound in the Eastern Goldfields Province. The presence of basal orthoquartzite in only the Southern Cross Province is another contrast. These features alone negate the possibility of stratigraphic correlation across the Yilgarn Block, but recognition of coherent stratigraphic sequences within each province is likely. Thus, in the Murchison Province, stratigraphic consistency has been recognized within and between individual fold belts. The sequence comprises a lower mafic unit, a felsic volcanic-sedimentary unit, an upper mafic unit, and an unconformable overlying sedimentary unit with some felsic volcanics. Stratigraphic continuity appears to be greatest at lower levels, and the basal tholeiite unit could have covered an area of 100000 km2. In the Southern Cross Province, the generalized sequence consists of a thin lower orthoquartzitic unit, overlain by a thick (5 km) mafic unit, and an equally thick volcanogenic sediment tary unit, which near Diemals is unconformable overlain by epiclastic sediments. Because of its economic importance, the Eastern Goldfields Province, particularly the southern part around Kalgoorlie, has been the subject of many attempts at stratigraphic synthesis: the most recent has been that of Gemuts & Theron (1975) who summarized much of the early work. Lithofacies variations across the Eastern Goldfields Province are expressed by the absence of banded iron formation and the presence of a chain of felsic volcanic centres within a central graben-like structure, called the NorsemanWiluna Belt (Fig. 1A). A chain of felsic volcanic centres of the extensive marine type is axially disposed within this zone. Williams (1974) suggested that vigorous volcanic activity and tectonic instability inhibited the deposition of iron formation, which would require quiescent intervals. Both Gemuts & Theron (1975) and Williams'" (1970) have proposed remarkably similar sequences in different regions of the NorsemanWiluna Belt. A generalized sequence comprises three mafic-ultramafic units, two felsic volcanic units, and an uppermost epiclastic sequence. The total thickness is at least 15km, and could be twice that amount. One of the most stratigraphically significant units, locally and informally known as the "Widgie Chert Marker", is a 100m thick cherty horizon continuous over at least 100 km. Similar multilayering involving only two major mafic layers, has been reported northward along the length of the Norseman-Wiluna Belt (Williams et al., 1976; Gower, 1976). In the Sir Samuel area, Bunting & Williams (1979) recognized two mafic layers, separated by a felsic unit,
CRUSTAL DEVELOPMENT IN YILGARN BLOCK
53 ferences between basins (i.e. Provinces) are original features. GEOLOGICAL MODELS FOR THE EVOLUTION OF THE YILGARN BLOCK Understanding of the overall tectonic controls on greenstone-basin development is limited by the absence of firm data on the age relationship between individual basins and the gneiss terrains. Granite detritus in greenstone sequences establishes that some sequences were accumulating while some of the domal granites were being emplaced, apparently tying both processes to the 2.7-2.6 b.y. tectono-thermal event. This is supported by a single Rb/Sr isochron of 2718 m.y. Nature of the Greenstone Depositories (minimum age) on the Kathleen Valley Gabbro Highly metamorphosed greenstone belts in the (Cooper et al., 1978), a sill emplaced into a stratisouthwestern part of the Southern Cross Pro- graphically low tholeiite sheet in the Eastern vince are probably remnants of a once more Goldfields Province. No data exist on the maxiextensive depository. Conversely, the northward mum age of greenstone sequences but the singular wedging out of greenstone belts into granite in the nature of the tectono-thermal event suggests that northern part of the Norseman-Wiluna Belt may all Yilgarn greenstone belts (volcanism, sedimenhave been an original feature. This is suggested tation, metamorphism and deformation) developed over a narrow time interval between 2.8 and by the decrease in metamorphism (Binns et 1976), the appearance of banded gneiss selvages 2.6b.y. From the regional patterns described in this around granite domes, and the development of fractionated high-level granite. Such belts cannot paper, it could be suggested that widespread therefore be interpreted as 'root zones", and volcano-sedimentary basins developed laterally Bunting & Williams (1979) suggest that they to much older gneissic terrain extant since about developed in separate rift valleys branching off a 3.2 b.y. The broad linearity of greenstone provinces may further suggest an accretionary model main basin to the south. With the severe disruption by granite doming, and invites comparison with modern continental and in the absence of comparable isotopic dating margins of the Pacific type. Several writers between provinces, lithofacies and stratigraphy (Tarney et al., 1976; Burke et al., 1976) advocate provide the only evidence of the nature and extent a general model whereby greenstone sequences of the greenstone depositories. We interpret the develop in a 'marginal basin" between a contithree greenstone provinces as representing separ- nental mass and an oceanic island arc. However, ate basins in which thick but simple layered this model is not applicable to the Yilgarn Block, sequences accumulated. Because assimilation of as (a) the granite corridors between basins do not greenstone by granite played no part in the verti- resemble island arcs; (b) the structural style of cal doming, erosion is the only process that could greenstone belts shows no asymmetry toward (or substantially affect the gross outline of the basins away from) the cratonic basement, and (c) the as seen on geological maps. An envelope around greenstone basins lie against a concave rather the main greenstone areas would therefore define than a convex-shaped continental mass. Although sequential development of elongate the minimum extent of the greenstone basins shown on Figure 3. Only in the fades changes basins away from the gneissic basement is posacross the Eastern Goldfields Province do we see sible, we see no evidence for invoking accretion evidence of the approaching marginal facies of onto a continental margin. Rather, the basins are seen as broad elongate downwarps, and only in one of these basins. From the data, it is impossible to reconstruct the Eastern Goldfields Basin is there evidence of the maximum extent of the greenstone basins. concurrent tectonism, and it takes the form of rift The overall variation of the greenstone-granite faulting. Some form of stable basement is required ratios throughout the Yilgarn Block is capable of opposing interpretation. The lowermost under the greenstone basins to preserve the layered volcano-sedimentary sequences, and all sequences possibly extended over vast areas. However, lithofacies of the higher stratigraphic the evidence points to the basement being predominantly sialic. Remnants of banded gneiss units suggest that there were discrete basins of deposition (Fig. 3), and that the stratigraphic dif- (both ortho and para) are a widespread, and in
and unconformably overlain by an epiclastie sequence. Whatever the precise nature of these correlations, an extensive "layer-cake" stratigraphy in the Norseman-Wiluna Belt is likely. One feature of regional importance may be the localized high-level presence of conglomerate, sandstone and shale. Although these rocks now appear to be restricted to local fault troughs, they possibly represent remnants of extensive alluvial fans and marginal sea deposits that marked a period of vigorous erosion initiated by block faulting in the final stages of greenstone belt evolution.
4
4
R. D. GEE, J. L. BAXTER, S. A. WILDE & I. R. WILLIAMS
54
GSWA 18594
Fig. 4. Diagrammatic crustal profile during accumulation of the greenstone sequences. The crust is sialic granulitic gneiss with some mafic granulite and metasedimentary material. The thick wedge of lower crust to the west is probably mafic granulite and may represent a very ancient proto-continental nucleus. Arrows indicate scale of subsequent basement doming. Vertical and horizontal scales approximately equal,
places, considerable component of the vast sea of granite throughout the Yilgarn Block. Many of the domally emplaced granites show transitional relationships with banded gneisses, and a large portion of the granite is apparently simply gneissic sialic basement remobilized during a rapid rise in the isotherms at the onset of the major thermal event. Basement remobilization was achieved by solid-state creep and by generation of "allochthonous" and "autochthonous" magmas. Density inversion is seen as the main tectonic driving force, resulting in diapiric rise of less-dense granite and compensatory sinking of more-dense greenstone sequences in a manner analogous to salt-dome tectonics. Tensional and lateral movements of independent crustal slabs, illustrated in the Norseman-Wiluna Belt, may be superimposed on this tectonic style. Present data are inadequate to estimate the contribution, if any, of granite directly from the mantle during this tectono-thermal event. The range in §r/ Sr initial ratios of the ca 2.6 b.y. granite is 0,702-0.709 and permits at least 200m.y., but no more than 400m.y. of crustal history since generation of sialic material from the mantle, unless expulsion of radiogenic strontium or some form of isotopic equilibration between lower crust and mantle is invoked (Armstrong, 1968). Most initial ratios of granite lie significantly above the mantle growth line (approx. 0.701 at 2.6b.y.) and therefore require at least two stages of fractionation from the mantle (Collerson & Fryer, 1978). This conclusion is consistent with the two-stage model of derivation of sialic melts from mantle material based on the experimental work of Green & Ringwood (1968). Oversby (1975) considers that high initial 238u/204pb (ju-value) recorded in some Eastern Goldfields granites indicates derivation from granitoid material that had resided in the crust for at least 300m.y. and possibly 700m.y., before the 2.6 b.y. tectono-thermal event. It is likely that during the 2.6b.y. event there was no major influx of sialic material into the crust, and that magma genesis was the result of 87
86
metamorphic differentiation and partial melting at intermediate as well as lower levels within the crust, which probably was, as it is now, only 30km in thickness. According to the deduced thermal gradient of at least 40°C/km, temperatures well in excess of 1200 °C could be experienced in the crust. These conditions, aided by volatile fluxing, and acting upon granulite-facies rocks, would lead to the formation of granitic melts. The subsequent rise of the sialic components leaving behind a more mafic restite, would account for the present two-tier crustal layering. The events in the interval 2.8-2.6 b.y. are seen as a unique episode that considerably modified the crust, but which added little new sialic material. The actual crustal-building processes considerably predated this event and in the Yilgarn Block we know very little of their timing and nature. The contribution of Nieuwland & Oompston (1981) adds considerably to our meagre knowledge and indicates a pre-history back to at least 3.3 b.y. The maturity of some sediments in the Jimperding Metamorphic Belt, and the presence of zircons which yielded the old ages, confirms that stable sialic basement already existed then. Similarly the lithofacies of these ancient sediments has features in common with more modern analogues. Notwithstanding the uniqueness of both the nature and scale of the 2.6 b.y. tectono-thermal event, uniformitarian principles can be applied usefully to the early Archaean evolution of the Yilgarn Block. ACKNOWLEDGMENTS The authors are indebted to their many colleagues in the Geological Survey of Western Australia who have contributed to the mapping of the Yilgarn Block. Their broad knowledge of the regional geology has been freely used in the synthesis presented here. Margaret Italiano drew the figures. This paper is published with permission of the Director of the Geological Survey of Western Australia.
CRUSTAL DEVELOPMENT IN YILGARN BLOCK
55
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- - , \919b: Structure and tectonic style of the Western Australian Shield. Tectonophys., 58, 327-369. GEMUTS, I., & THERON, A., 1975: The Archaean between Coolgardie and Norseman—stratigraphy and mineralization; in Knight, C. L. (Ed.) Economic Geology of Australia and Papua New Guinea, 1, Metals, 66-74. Australas. Inst. Min. MetalL, Parkville. GLIKSON, A. YI, & LAMBERT, I. B., 1976: Vertical zonation and petrogenesis of the early Precambrian crust in Western Australia. Tectonophys., 30, 55-89. GOLDING, L. Y., & WALTER, M. R., 1979: Evidence of
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MER, G. E., 1978: Crustal development in the Agnew region, Western Australia, as shown by Rb/Sr isotopic and geochemical studies. Precamb. Res., 7, 31-59. CURRIE, K. L., 1971: The reaction 3 cordierite = 2 garnet + 4 sillimanite + 5 quartz as a geological
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The Archaean Marda igneous complex, Western Australia. Precamb. Res., 3, 111-136.
HICKMAN, A. H . , 1981: Crustal evolution of the Pilbara
Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 57-69.
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R. D. GEE, J. L. BAXTER, S. A. WILDE & I. R. WILLIAMS
LIBBY, W. G., 1979: Regional variation in granitic rock;
in Contributions to the Geology of the Eastern Goldfields Province of the Yilgarn Block. Rep. geol. Surv. West. Aust., 9, 53-109. MARSTON, R. J., & TRAVIS, G. A., 1976: Stratigraphic implications of heterogeneous deformation in the Jones Creek Conglomerate (Archaean), Kathleen Valley, Western Australia. J. geol. Soc. Aust., 23, 141-156. MATHUR, S. P . , Moss, F . J . , & BRANSTON, J . C., 1977: Seismic and gravity investigations along the geotraverse, Western Australia, 1969. Bull. Bur. Mirier. Resour. Geol. Geophys. Aust., 191. NALDRETT, A. J . , & TURNER, A. R . , 1977: The geology and pedogenesis of a greenstone belt and related nickel sulfide mineralization at Yakabindie, Western Australia. Precamb. Res., 5, 4 3 - 1 0 3 . NIEUWLAND, D . A., & COMPSTON, W., 1981: Crustal evolution in the Yilgarn Block near Perth, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 159-171. OVERSBY, V . M . , 1975: Lead isotope systematics and ages out of Archaean acid intrusives in the Kalgoorlie-Norseman area, Western Australia. Geochim. cosmochim. Acta., 39, 1 1 0 7 - 1 1 2 5 . RUTLAND, R . W. R . , 1976: Orogenic evolution of Australia. Earth Sci. Review, 12, 161-196.
TARNEY, J . , DALZIEL, I. W . D . , & D E WITT, M . J.,
1976: Marginal basin "Rocas Verdes" complex from S. Chile: a model for Archaean greenstone belt formation; in Windley, B. F. (Ed.) The Early History of the Earth, 131-146. Wiley, London. WILDE, S. A., 1976: The Saddleback Group—a newly discovered Archaean greenstone belt in the southwestern Yilgarn Block. Ann. Rep. geol. Surv. West. Aust. for 1975, 92-95. WILDE, S. A., & Low, G. H., 1978: Perth, Western Australia—1:250 000 Geological Series. Explan. Notes geol. Surv. West. A ust., SH50-14. WILLIAMS, I. R., 1970: Kurnalpi, Western Australia— 1:250000 Geological Series. Explan. Notes geol. Surv. West. Aust., SH51-10. , 1974: Structural subdivision of the Eastern Goldfields Province, Yilgarn Block. Ann. Rep. geol. Surv. West. Aust. for 1973, 53-59.
WILLIAMS, I. R . , GOWER, C . F . , , & THOM, R . , 1976:
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Byro 1:250000 Geological Sheet, Western Australia. Rec. geol. Surv. west. Aust.
CRUSTAL EVOLUTION OF THE PILBARA BLOCK, WESTERN AUSTRALIA A. H . Hickman
Geological Survey of Western Australia, Mineral House, 66 Adelaide Terrace, Perth, Western Australia 6000 ABSTRACT The tabular geometry of the Warrawoona Group across the greater part of the Pilbara Block rules out deposition in separate geosynclines located in the present positions of the greenstone belts, and casts serious doubt on the credibility of any model involving migrating depositional basins (as in island arc-marginal basin systems). Deposition of the Pilbara Supergroup commenced at about 3600-3550 m.y. when tholeiitic basalt was extruded over an area at least 450 km by 200 km. The basement to this layer was probably a 10-20 km sialic crust composed of reworked granitic and volcanic-sedimentary material. Subsequent calc-alkaline volcanism (3550-3450 m.y.) and contemporaneous intrusion of trondhjemite-granodiorite was followed by widespread deposition of chert and carbonate sediments, and extrusion of pillowed tholeiite, high-Mg basalt and local peridotitic komatiite. Dome and syncline structures then began to develop, but some areas remained stable with accumulation of rhyolite, clastic sediments and banded iron-formation. Deformation became more widespread and erosion resulted in the deposition of thick sandstone, conglomerate and turbidite beds, probably in separate basins. Deformation culminated at about 2950m.y. and was accompanied by syntectonic intrusion of granodiorite-adamellite. Subsequent erosion was followed by local deposition of volcanic and sedimentary rocks, and post-tectonic intrusion of adamellite-granite.
INTRODUCTION The Pilbara Block occupies 60000 km of northwestern Australia, and is an Archaean lowgrade terrain in which domal granitic batholiths up to 100 km across are separated by synclinal "greenstone belts'' containing greenschist- to lower amphibolite-facies, metavolcanic, and metasedimentary rocks. Abundant exposure, combined with a general simplicity of major structures, has permitted a stratigraphic analysis of the entire greenstone succession, which is now referred to as the 'Pilbara Supergroup' (Hickman, in press). This paper briefly reviews previous interpretations, summarizes the regional geology, and presents a new interpretation of the crustal evolution of the block. 2
PREVIOUS INTERPRETATIONS Geological investigation of the Pilbara Block commenced in 1890, but until about 1960 most studies were concentrated on particular mineral deposits or were very general. This paper mentions only the most important recent contributions, but a comprehensive account is provided by Hickman (in press). Since 1960, emphasis has been on detailed mapping and stratigraphic interpretation, and
Spec. Pubis geol. Soc. Aust., 7 (1981)
recently there have been rapid developments in geochronology and geochemistry. The age relationship between the granites and the greenstones is fundamental, and over the years two general schools of thought have emerged. One school regards the Pilbara Supergroup as the oldest material present, and the granitic batholiths as later magmatic diapirs (Montgomery, 1907; Clarke, 1923; David, 1932, 1950; Ryan, 1964; Horwitz, 1966 and in Fitton et at, 1975; Glikson, 1972, 1978; Blockley, 1975). The other considers that the Pilbara Supergroup was deposited on a granitic basement, remnants of which probably exist within the granitic batholiths (Maitland, 1904; Noldart & Wyatt, 1962; Brandt, 1964; Brown et al., 1968; Hickman, 1975; Hickman & Lipple, 1975; Fitton et al. 1975; Miller, 1975; Ingram, 1977). Other questions concern the depositional and structural history of the area and the precise stratigraphic relationships of certain units within the Pilbara Supergroup. Two principal phases of deformation were recognized by Noldart & Wyatt (1962): the Warrawoona orogeny which predated deposition of the Mosquito Creek succession (Gorge Creek Group) and produced dome and syncline structures by superimposed folding, and secondly the Mosquito Creek orogeny which produced easttrending folds. They regarded most of the granite 9
A. H. HICKMAN
58
masses as intrusive, but added that gneissic rocks such as those between Port Hedland and Marble Bar had formed by in situ granitization of prevolcanic sedimentary rocks constituting an older basement. In the west Pilbara, Ryan (1964, 1965, 1966), Kriewaldt (1964), Ryan & Kriewaldt (1964), Kriewaldt & Ryan (1967) and subsequently Milier (1973, 1975), recognized a eugeosynclinal succession deposited in a northeasterly trending trough between Port Hedland and Croydon (Fig. 1). The centre of this trough was filled with about 12 km (Ryan, 1965) of turbiditic sediments but the upper part of this succession passed laterally northwestward into contemporaneous volcanics between Roebourne and Whim Creek, and southeastward into similar volcanics at Pilbara Well. The volcanic facies was correlated with tjie Warrawoona succession (Warrawoona Group) of the east Pilbara. By drawing analogies between the greenstone belts of the Pilbara Block and the Barberton greenstone belt of South Africa, Anhaeusser (1971#, b) suggested a spatial relationship between depositional troughs and the present location of greenstone synclines. Anhaeusser also referred to volcanic cycles near Marble Bar, and
INDIAN
the concept of volcanic cyclicity was used by Ingram (1977) to subdivide the east Pilbara succession into five formations. Between 1972 and 1975 the Geological Survey of Western Australia remapped about 70% of the Pilbara Block while the greater part of the remaining area was independently re-assessed by other workers. This mapping demonstrated that the stratigraphic succession at Marble Bar could be traced, with only minor variations, from one greenstone belt to another across distances of over 100 km (Hickman & Lippie, 1975; Lipple, 1975). At the same time, Fitton et at. (1975) reported that the succession in the west Pilbara, between Roebourne and Pilbara Well, was so similar to that established in the Marble BarNullagine area that stratigraphic correlation was possible over 300 km. Fitton et al. (1975) solved a long-standing problem in regional correlation across the Pilbara Block by discovering that structural complications had caused earlier workers to present the west Pilbara succession the wrong way up. In the east Pilbara, Hickman (1975) recognized four principal phases of Archaean deformation, the domes and synclines originating during the second phase, D2. He concluded that the domes were formed by essentially
OCEAN
c3 ^
Nunyerry
1
1 POST-ARCHAEAN
I X S I
ROCKS
WHIM CREEK GROUP GORGE CREEK GROUP
1
I H
HLHLLI |V*V]
WARRAWOONA GROUP, UPPER PART TOWERS FORMATION (DASHED = ABSENT) WARRAWOONA GR0UR LOWER PART POST-TECTONIC GRANITIC ROCK FOLIATED-GNEISSIC GRANITIC ROCK
Fig. 1.
Stratigraphic map of the Pilbara Block.
CRUSTAL
EVOLUTION
IN PILBARA
BLOCK
59
TABLE I
Stratigraphy and geochrortology of the Pilbara Supergroup
WHIM CREEK GROUP
Group a n d Subgroup
F o r m a t ion
Main
Thickness (km)
Lithology
Negri
Volcanics
Basalt and andesite
Louden
Volcanics
Basalt and ultramafics
1.0
Rushall
Slate
Slate, minor
tuff
0.2
volcanics
0.5
M o n s Cupri - Warambie
Volcanics
Felsic
Basalt
- V e s icular
Geochronological Formation
Lai la R o o k h
Sandstone
Honeyeater
Basalt
WARRAWOONA GROUP
Taiga Taiga Subgroup
Sal gash Subgroup
<_>
cd ec o CJJ
Soanesv i 1.1 e Subgroup
Mosquito Creek F o r m a t ion
ta
C1eavervi11e F o r m a t ion
basalt
Psammitie - pelitic schist Sandstone and conglomerate
2 6 1 0 ± 80 (1)
Basalt
1.0
B a n d e d i ron f o rma t i on
1.0
Basa1t
1.0
Formation
Metasedimehts
1.5
Formation
Rhyolite
1.0
B a s a l t , k o m a t i ite
2.0
Felsic
1.0
Apex
F o r m a t ion
Basalt
Towers
Formation
Duffer
3 0 5 0 ± 180 (4); 3 0 4 0 ( 5 ) , 2 9 2 0 (2)
volcanics
2840 - 2930
2915 ±207 (1) 2600
Basalt, komatiite
2.0.
3 3 4 0 - , 3 2 0 0 * (10)
0.5
3 4 2 0 (10)
5.0
3 5 0 0 (11); 3452 ± 16 (12); 3 1 2 5 ± 3 6 6 (13); 2 8 8 0 ±66 (14) ; 3 4 7 0 a , 3 4 7 0 b (lo); 3 2 8 0 ± 20 (3); 3 2 7 9 ±169 (I) 3 5 5 0 ± 3 0 , 3 5 2 0 ±30 (18)
Felsic Basalt
2.0
McPhee
Carbonate schist and chert
0. 1
Basal t
2.0
N o r t h Star
Basalt
from galena
< w w v w v w v \
unconformity,
regional
(15) :
2 9 3 6 ± 9 , 3 0 7 0 ±12 (2) ; 2951 ± 83 ( 9 ) , 3280 ±20, 3227 ± 5 0 , 2 9 5 0 ±50 (17)
Chert and basalt
Formation
* Minimum a g e obtained
(6); 2 9 1 0 (5)
3 3 4 0 * (8)
Mount Ada Basalt Formation
(2) - 307C> (3)
5.0
Corboy
Basalt
(climax) 2 9 5 0
3.0
C h a r t e r is B a s a l t
Euro
Formation
0.2
Wyman
Panorama
into
0.2
D2 Q. ©
Evidence
Intrusions
in vein
volcanics
2 9 0 0 - 3 0 0 0 (10)
3 5 6 0 ±32 (16); 3 5 7 0 ± 1 8 0 , 3 4 6 0 ±120, 3 4 9 0 ±60, 3 5 9 0 ±100, 3 5 9 0 ±90 (7)
quartz. local
D 2 is t h e m a i n e p i s o d e o f d e f o r m a t i o n a n d metamorphisrn Group and the Whim Creek Group.
unconformity.
(see t e x t ) w h i c h
led to a n important u n c o n f o r m i t y b e t w e e n
the Gorge
Creek
References: 1 , J . R . d e L a e t e r , p e r s . c o m m . ; 2 , O v e r s b y , 1 9 7 6 ; 3 , P i d g e o n , 19782?; 4 , C o m p s t o n & Arriens,. 1 9 6 8 ; 5 , L e g g o e t a i . , 1965; 6 , G r e e n h a 1 g h & J e f f e r y , 1 9 5 9 ; 7 , J a h n e t a i . , u n p u b l i s h e d d a t a ; 8 , R i c h a r d s , 1 9 7 8 ; 9 , d e L a e t e r et al., 1 9 7 5 ; 1 0 , J . R . R i c h a r d s , p e r s . c o m m . (a a n d b a r e f r o m s e p a r a t e l o c a l i t i e s ) ; 1 1 , S a n g s t e r & B r o o k , 1 9 7 7 ; 1 2 , P i d g e o n , 1978a; 1 3 , d e L a e t e r & B l o c k l e y , 1 9 7 2 ; 1 4 , d e L a e t e r & T r e n d a l l , 1 9 7 0 ; 1 5 , d e L a e t e r e t a i . , 1 9 7 7 ; 1 6 , H a m i l t o n e t al., 1 9 8 1 ; 1 7 , C o o p e r e t al., 1 9 8 0 ; 18, M . M c C u l l o c h , p e r s . c o m m . T h e s t r a t i g r a p h i c s i g n i f i c a n c e o f t h e data
is i n t e r p r e t e d
by t h e w r i t e r .
solid-state uplift rather than diapiric magmatic intrusion, and that the greenstone succession was regionally tabular and not deposited in troughs positioned along the present greenstone belts. Hickman & Lipple (1975) confirmed observations of previous workers (e.g. Blockley, 1975) that granite-greenstone contacts were either intrusive or tectonic, and the 1972-1975 mapping revealed no basal unconformity to the Pilbara Supergroup. Structural evidence, however, suggested that remnants of a pre-Pilbara Supergroup sialic basement might be preserved within the batholiths. Geochronological evidence obtained since 1975 is summarized in Table I, and establishes that the greenstone succession of the Pilbara Block is
chiefly between 3600 and 2950m.y. old. Within this range, the age of the Gorge Creek Group is still poorly defined and could be as great as 3400-3300 m.y. Palaeontological discoveries (Dunlop, 1976; Dunlop et al., 1978; Walter, 1978; Hickman et al., 1980; Lowe, unpublished data; I. Martin, pers. comm.) and sedimentological observations (Dunlop, 1978; Barley, 1978; Barley et al., 1979) have indicated that parts of the succession were deposited in shallow water. REGIONAL GEOLOGY Main
Subdivisions
About 60 percent of the Pilbara Block is occupied by granitic rocks (Fig. 1). They fall into
A. H. HICKMAN 60 three main categories: (1) 3500-2900m.y. migma- (the area is complicated by faulting) yet the fortitic, gneissic and foliated granodiorite and mation itself remains continuous; therefore some adamellite with minor tonalite and trondhjemite, of the North Pole cherts, which represent silici(2) foliated porphyritic granodiorite and adamel- fied carbonate and evaporite units (Dunlop, lite of various ages but chiefly about 3000 m.y. 1978; Dunlop & Buick, 1981), may have been old, and (3) unfoliated, post-tectonic granite and deposited en echelon. In other areas the more adamellite (commonly tin-bearing) 2700m.y. to typical grey/white banded cherts of the forma2600m.y. old. The granitic rocks form domal tion are far more continuous and probably repreplutonic complexes separated by synclines and sent primary silica precipitates or silicified oozesynclinoria containing the Pilbara Supergroup, like sediments similar to those described by Lowe the succession of which is summarized in Table I. & Knauth (1977) in the Onverwacht Group of Although the maximum total thickness of the Pil- South Africa. At Marble Bar the Marble Bar bara Supergroup is 30 km, its true thickness in Chert Member can be traced for about 80 km individual areas is generally only 10 to 15km. along strike, and its eventual termination is due Metamorphic grade is principally greenschist to faulting. fades, but varies from sub-greenschist to lower Stratigraphic markers also occur within the amphibolite. Taiga Taiga Subgroup, the Salgash Subgroup and the Gorge Creek Group. The McPhee Formation Pilbara Supergroup is approximately 80 m thick and consists of an Stratigraphic correlation across the Pilbara impersistent lower unit of grey/white chert and Block is facilitated by good exposure and con- an upper unit of carbonate schist. In the type siderable continuity between greenstone belts. At section at McPhee Reward Mine this schist is North Pole and McPhee Creek (Fig. 1) succes- composed of carbonate, chlorite and quartz, and sions of neighbouring synclinal belts are directly contains thin intercalations of ferruginous chert. linked, without interruption or attenuation, High nickel (500-1000 ppm) and chromium (up to across broad greenstone domes underlain by 2000 ppm) contents suggest that the rock is an granitic substratum (Hickman, 1975). Thus, simi- altered ultramafic, possibly a tuff or lava. The larities in the lithological sequences of the various formation separates the North Star Basalt from belts are not merely due to common evolutionary the Mount Ada Basalt north of Marble Bar and patterns , of separate depositories. Where indi- 50 km southwest of Marble Bar , and is correlated vidual belts are separated by granitic rocks or with a similar unit at Mount Sholl in the west Proterozoic cover, correlations are based on Pilbara. Another central marker separates the marker units, combined with similarity of Apex Basalt from the Euro Basalt in most areas sequence. Clastic sedimentary rocks of the Gorge where the upper part of the Warrawoona Group Creek Group exhibit important lateral facies is shown on Figure 1. This central unit, the changes, but within the Warrawoona Group all Panorama Formation, consists of felsic volformations other than the Duffer Formation are canics, chert and local metasediments. The essentially tabular. As to be expected in such an Cleaverville Formation of the Gorge Creek extensive volcanic pile as the Warrawoona Group is a regionally persistent unit of banded Group, detailed mapping of individual beds and iron-formation and retains its thickness of about members has revealed lateral interfingering of 1000m over most of the Pilbara Block. Apart flows and sediments, but such variations are in- from true banded iron formation, the Cleaversignificant on a regional scale. ville Formation also contains ferruginous chert, The Towers Formation, which is about 0.5 km grey/white chert, tuff and minor ferruginous thick and can be traced for 430 km across the shale, siltstone and sandstone. Pilbara Block, is the main stratigraphic marker Basaltic formations of the Pilbara Supergroup unit of the Warrawoona Group. It generally conprincipally composed of pillowed tholeiite sists of three prominent chert members separated are and consanguineous dolerite sills, but peridotitic by pillow tholeiite and high-Mg basalt, felsic tuff komatiite, basalt, dacite and rhyolite are or clastic sediments. In the east Pilbara the cen- present at high-Mg levels. It is notable that the tral and thickest (100 m) chert, the Marble Bar ultramafic several components not concentrated in Chert Member, is exceptionally well exposed at the lower part of the aresuccession; peridotitic Marble Bar Pool where it consists of red/white komatiite is most common the Salgash and grey/white banded chert. Elsewhere grey/ Subgroup, and high-Mg basaltwithin occurs in all the white banded chert predominates, although green basaltic formations except the Warambie fuchsitic chert is common at some localities. At and the Negri Volcanics. Volcanic cyclicity Basalt is not North Pole certain individual chert units of the feature of the Pilbara Supergroup, Towers Formation appear to terminate laterally aandprominent although mafic-felsic-chert cycles occur at
CRUSTAL EVOLUTION IN PILBARA BLOCK member scale, especially within the Salgash Subgroup, andesite is rare and the felsic sections are generally thin. The thickest felsic formation, the Duffer Formation, occurs in the lower part of the Pilbara Supergroup, and has been dated by Pb/Pb (Sangster & Brook, 1977; Richards, pers. comm.), U / P b (Pidgeon, 1978a) and Sm/Nd (M. McCulloch, pers. comm.) methods at between 3550 and 3450 m.y. At Marble Bar the formation is 5 to 8 km thick and is chiefly composed of dacitic lava, tuff, agglomerate and debris flows with rhyolite in its upper sections and andesite and minor basalt near its base. The formation is absent in many areas of the Pilbara Block and was clearly deposited in a series of wedge-shaped volcanic piles. As noted by Barley et al. (1979), at the edges of these piles felsic units pass laterally into more mafic material belonging to the upper part of the Taiga Taiga Subgroup, but the Duffer Formation always occurs beneath the Towers Formation and the writer is not aware of any evidence to support the interpretation of Barley et al. (1979, Fig. 1) that parts of it are locally equivalent to mafic units of the Salgash Subgroup. The Wyman Formation is chiefly composed of porphyritic rhyolite (ultra-potassic in certain areas) with minor tuff, agglomerate and very rare basalt. The Gorge Creek Group contains six formations, three of which (Corboy Formation, Charteris Basalt and Honeyeater Basalt) are confined to the east Pilbara. The two basaltic formations are Iithologically similar to those in the Warrawoona Group, but are wedge-shaped and of only local importance. The Corboy Formation includes sandstone and psammopelitic rocks, and the Lalla Rookh Sandstone (up to 5000 m thick) is composed of well-sorted, slightly feldspathic sandstone, quartzite and conglomerate. The sandstone is bedded at intervals of 0.1 to 1.0 m and commonly exhibits cross bedding. It is notable that where the Lalla Rookh Sandstone is thickest the overlying Mosquito Creek Formation, a succession of turbidite deposits, is thin or absent, and vice versa. In certain areas, such as Eastern Creek (60 km northeast of Nullagine) and Croydon, sandstone at the base of the Mosquito Creek Formation passes upward into psammopelitic and pelitic units and, as suggested by Eriksson (1981), the Mosquito Creek Formation is probably partly contemporaneous with the Lalla Rookh Sandstone. The Whim Creek Group is composed of volcanic and sedimentary formations and is restricted to a relatively small area of the west Pilbara. It is not deformed by the main deforma-
61
tion (D 2 , Hickman, 1975) and unconformably overlies the Warrawoona and Gorge Creek Groups. Two basaltic formations which unconformably overlie the Whim Creek Group (Table I) are of only local significance.
Granitic Batholiths The large domal granitic complexes of the Pilbara Block are conveniently termed 'batholiths', but it is emphasized that they owe their present geometry to tectonic processes rather than simple magmatic intrusion. These batholiths are not magmatic diapirs but were formed by solid-state upward doming of a pre-existing plutonic complex (Hickman, 1975; Bickle et al., 1980). The dominant foliation is tectonic, crosses internal plutonic contacts, and is generally parallel to batholith margins and the regional schistosity of adjacent greenstone belts. It varies from a weak biotite- alignment to a stronger parallel orientation of quartz, sericite and mafic minerals. Granitic rocks near granite-greenstone contacts are commonly transformed to schist containing zones of protomylonite. Few of the domes are simple structures; most, especially the larger ones, contain synforms in which the foliation converges downwards. These intradomal synforms (Hickman, 1975) are marked by trains of greenstone material and, in some places, by shear belts. Such features probably indicate root zones of eroded greenstone synclines. Geochemical patterns across these zones (Davy & Lewis, 1981) ^re consistent with the interpretation that the synforms deform pre-dome granite-greenstone contacts. Oversby (1976) recognized an important metamorphic event at 2950m.y., and suggested that it accompanied doming. Cooper et al. (1980) reach a similar, conclusion, and Pidgeon (1978b) considered that doming took place at about 3000 m.y. Geochronolgical data (Rb/Sr, U/Pb and Sm/Nd methods) are accumulating to indicate that large parts of the batholiths are composed of rocks between 2900 and 3500m.y. old, ages in excess of 3400m.y. being indicated by zircon U / P b (Pidgeon, 1978Z?) and whole-rock Sm/Nd (M. McCulloch, pers. comm.) methods. Thus, at the time of doming, some of the granitic rocks were already about 500 m.y. old, and simple magmatic diapirism cannot be responsible for the batholiths. The entire range of granite rock types in the batholiths is described by Hickman (in press) and individual batholiths and plutons are discussed by Blockley (1980), Davy & Lewis (1981) and Bettenay etal. (1981).
A. H. HICKMAN
62
CRUSTAL EVOLUTION Evolutionary Models Evolutionary models proposed to explain other granite-greenstone terrains fall into five general categories: oceanic (e.g. Glikson, 1972), geosynclinal (e.g. Anhaeusser et al., 1969; Anhaeusser, 1971a), rift-zone (e.g. Windley, 1973), platetectonic (e.g. Goodwin & Ridler, 1970) and extraterrestrial (e.g. Green, 1972). Stratigraphic, geochemical and structural evidence limit the application of such models to the Pilbara Block. Stratigraphic Evidence The stratigraphic succession of the Warrawoona Group is tabular and extends over an area of at least 60000 km . Individual formations can be traced continuously between greenstone belts, establishing that the major synclines are entirely tectonic in origin and bear no direct relationship to depositional basins. The tabular nature of the Warrawoona Group rules out deposition in troughs or trenches located in the present positions of the greenstone belts. Also, individual belts reveal no evidence of faeies changes or primary stratigraphic thinning toward syncline margins. Attenuation, where present, is the result of tectonic flattening. The. stratigraphic continuity of the Warrawoona Group also precludes models involving migrating depositories; many of the proponents of island-arc and marginal-basin models have suggested the process to explain broadly parallel greenstone belts separated by granitic terrain (as in the Yilgarn Block and the Canadian Shield). Comparisons between the Warrawoona Group and examples of Phanerozoic oceanic crust reveal several important differences. The former contains no lower section of ultramafic and gabbroic rocks and very few dolerite dykes, whereas it does include a 5 to 8 km thick felsic formation (Duffer Formation), shallow-water sediments and silicified evaporites. If, despite its paucity of ultramafics, gabbro and dykes, the Taiga Taiga Subgroup is considered in isolation, there remains the problem that this unit bears less resemblance to oceanic crust than the Salgash Subgroup because the latter contains more peridotitic komatiite and high-Mg basalt. By modern analogy, the Salgash Subgroup is unlikely to be oceanic because it overlies shallow-water sediments and the Duffer Formation. General objections to the oceanic model, which apply to all greenstones, are discussed by Windley (1977). Several workers have suggested that Archaean geothermal gradients were abnormally high, but Burke et al. (1976) have pointed out that if this were so the basal sections of 10-15 km thick 2
greenstone piles would show effects of high-grade metamorphism. This is not so in the Pilbara, and it can reasonably be inferred that the North Star Basalt was deposited on a crust which was at least 10km thick. Moreover, the Rb/Sr method of estimating crustal thickness (Condie, 1973) suggests that the Pilbara crust was 15-20 km thick during deposition of the North Star Basalt. It is improbable that this crust was basaltic or ultramafic because no such remnants are preserved at the base of the Warrawoona Group, and mafic xenoliths within the foliated granite complex reveal no evidence of high-grade metamorphism. If the pre-Pilbara Supergroup crust were essentially granitic, however, it could have been incorporated into the granitic complex, and would now be relatively difficult to identify. Thin sandstone units in the Taiga Taiga Subgroup near McPhee Reward Mine, and in the Salgash Subgroup about 15 km south of Marble Bar, could provide evidence of underlying or adjacent granitic basement, but remain to be studied in detail. The sandstone south of Marble Bar is composed of 95% quartz grains ( about 1 mm in diameter), feldspar, biotite, muscovite and rare zircon. Granitic pebbles in the Gorge Creek Group 40 km northwest of Marble Bar establish erosion of granitic rocks during deposition of this unit which, as noted earlier, could be as old as 3400-3300m.y. Geochemical Evidence The origin of tonalite-granodiorite batholiths is widely debated, possible sources being the mantle or earlier crustal material of basaltic, granitic or sedimentary composition. On the basis of criteria used elsewhere (Glikson & Sheraton, 1972; Chappell & White, 1974; Tarney, 1976; White & Chappell, 1977) derivation from sediments is improbable for the Pilbara batholiths (Hickman, in press; Davy & Lewis, 1981). The geochemistry of rare-earth elements (Jahm et al., unpublished data) indicates that the felsic volcanic and granitic rocks of the Pilbara Block did not originate by direct partial melting of the upper mantle or by single-stage partial melting of amphibolite; either the source material was heterogeneous or, assuming an amphibolite source, variable fractional crystallization was involved. Various workers (e.g. Glikson, 1976) have argued that granitic rocks with low concentrations of the large-ion lithophile elements (with the exception of Sr and Ba) are unlikely to have been derived by reworking of earlier .sialic crust because initial partial melting would produce rocks rich in large-ion lithophile elements. Certainly, concentrations of these elements are
63 Sr ratios are as low as 0.702 (Glikson, 1979, tabulates examples). Rb metasomatism and losses of Sr, unless uniform throughout an intrusion, would normally result in a 'poor-fit' isochron, and most of the Pilbara Rb/Sr isochrons are reasonably well defined. However, such processes have been recorded in zones of contact metamorphism. De Laeter & Blockley (1972) suggest that the isotopic compositions of foliated granitic rocks close to the post-tectonic Moolyella Adamellite have been changed, and Lewis et al. (1975) report losses of radiogenic Sr in Archaean granitic rocks close to a Proterozoic dolerite dyke. The existing Rb/Sr data provide no clear evidence on the basement problem, not only for the reasons outlined above, but also because most of the few rocks dated are known to intrude the Pilbara Supergroup. If remnants of pre-Pilbara Supergroup basement exist within the batholiths, initial identification of them will probably be achieved by using a combination of structural, petrographic and geochemical evidence. Some preliminary results of a detailed geochemical study of volcanic rocks in the Pilbara Supergroup are reported by Glikson & Hickman (1981). One aim of that study was to examine the feasibility of determining Archaean tectonic environments using geochemical comparisons between Archaean and modern volcanics. The various criteria which were applied, for example individual element concentrations and ternary diagrams such as Ti-Zr-Y and Ti-Zr-Sr (Pearce & Cann, 1973), Ti0 -K 0-P 0 (Pearce et al., 1975) and Fe0 Mg0-Al 0 (Pearce et al., 1977), gave conflicting indications of environment. In general, however, basalt from the Warrawoona Group more closely resembles basalt from modern oceanic to island-arc environments than does basalt from the Gorge Creek Group. The geochemistry of the felsic volcanic formations is discussed by Hickman (in press), Glikson & Hickman (in press), Barley (1981) and Jahn et al. (unpublished data). The highly fractionated chemistry of the Wyman Formation is consistent with relatively low proportions of partial melting of sialic crust whereas the calc-alkaline Duffer Formation was probably genetically related to some of the granodioritic rocks in the batholiths (Glikson & Hickman, iii press).
CRUSTAL EVOLUTION IN PILBARA BLOCK
generally low in the foliated granitic rocks of the Pilbara Block, but exceptions occur, about 8% having Rb/Sr ratios greater than 1.0, and 15% K/Rb ratios of less than 150 (Hickman, in press). It is possible that the pre-Pilbara Supergroup crust was tonalitic and that it underwent a high degree of partial melting. Davy & Lewis (1981) present evidence that the Mount Edgar Batholith, east of Marble Bar, originated by partial melting of acid to intermediate granulite in the lower crust. It is commonly argued that low initial Sr/ Sr ratios, as indicated by several of the Rb/Sr isotopic studies in the Pilbara Block, are evidence that the granitic rocks must either be derived directly from the Archaean mantle or from some pre-existing rock of almost the same age. Most of the Rb/Sr work has indicated ages between 3300 and 2900m.y. (Compston & Arriens, 1968; de Laeter & Blockley, 1972; de Laeter etaL, 1975; de Laeter, pers. comm.; Cooper et al., 1980) with initial Sr/ Sr ratios generally below 0.703. The significance of these ratios depends on sample population, the Rb/Sr ratio of the source material, and the assumption that there has been no Rb metasomatism or losses of radiogenic Sr since intrusion. In the first three of the five references quoted above, sampling sites were separated by distances of over 10 km, and the samples used in each work included a wide range of rock types. In view of the plutonic complexity of the batholiths, the various isochrons generated may include rocks of different plutons and different ages (a possibility noted by de Laeter et al 1975). Under these circumstances the ages calculated may be closest to those of the most fractionated (probably youngest) components, whereas the initial Sr/ Sr might relate to the older, low-Rb samples. A Sm/Nd age determination of 3490 ± 30m.y. (M. McCulloch, pers. comm.) on a sample of gneiss from the Shaw Batholith appears to support this possibility; this sample was used by de Laeter et al. (1975, Fig. 42) and is positioned on the low- Rb/ Sr end of a 2951 ± 83 m.y. isochron. The average Rb/Sr ratio of tonalitic to granodioritic rocks in the Pilbara batholiths is 0.3 (Hickman, in press) and the Rb/Sr mode for 265 such rocks from the Mount Edgar Batholith is approximately 0.18 (data from Dr R. Davy). Thus, if these rocks were derived from preexisting sialic crust the Rb/Sr ratio of this crust would probably have been about 0.1 or below (assuming enrichment of Rb through partial melting). This low Rb/Sr ratio has important implications since it generally extends the probable crustal history of the source materials by several hundred million years, even where initial Sr/ 87
86
87
86
87
86
87
86
87
86
87
2
r
2
2
2
5
3
Structural Evidence The granitic domes and the intervening greenstone synclines are not the product of superimposed folding because there is no evidence of interference at greenstone belt intersections: instead, the associated cleavage passes continu-
A.
H.
HICKMAN
A
ca
B
3450 m.y. Local deformation and deposition of Duffer Formation
C
ca
D
3400-2950 m.y. Deformation, local erosion and deposition of Wyman Formation to Cleaverville Formation
F i g . 2.
3550 m.y. Deposition of Taiga Taiga Subgroup
3400 m.y. Deposition of Towers Formation and Salgash Subgroup
Stages i n t h e e v o l u t i o n o f t h e P i l b a r a S u p e r g r o u p - a d i a g r a m m a t i c r e p r e s e n t a t i o n .
CRUSTAL EVOLUTION IN PILBARA BLOCK
E
ca 2 9 5 0 m . y . D e f o r m a t i o n , local erosion a n d deposition of Lai la R o o k h S a n d s t o n e and Mosquito Creek F o r m a t i o n . Major d e f o r m a t i o n .
F
2 7 0 0 - 2 6 0 0 m.y.
Deformation, erosion and deposition of W h i m Creek G r o u p
GRANITIC ROCKS
PILBARA SUPERGROUP H
Shale and mudstone
| P ° | Porphyritic adamellite
Siltstone
I x x | Granodiorite to adamellite
Sandstone and conglomerate Chert
a
:
0
1 Post-tectonic adamellite
Turbidites
Banded ir<
m
|
^
Tonalite to granodiorite
F • ; • -I Early crust, dioritic to granodioritic I- • ' - I and probably gneissic with enclaves of pre- 3600 my supracrustals
Felsic lava and tuff Approximate vertical scale in kilometres
Felsic agglomerate Mafic volcanics HO
65
66
A. H. HICKMAN
ously from one syncline to another, generally changing strike by 50 to 130 degrees and remaining parallel to the nearest granite-greenstone contact. With some exceptions (Bickle et al., 1980) the oldest preserved structures in the Pilbara Supergroup belong to the doming event whereas earlier structures are common within the batholiths (Hickman & Lipple, 1975; Hickman, 1975 and in press). This suggests that parts of the batholiths pre-date the Pilbara Supergroup, but it must be remembered that many of these early structures could have developed in the granitic substratum during the 600m.y. interval between deposition of the North Star Basalt and the culmination of the doming event. Bickle et al. (1980) and Bettenay et al. (1981) present evidence for local horizontal tectonics during this interval. The present structure and composition of the Pilbara crust (Drummond, 1979) precludes derivation of the batholiths by partial melting of basaltic rocks in the Warrawoona Group, a process suggested in an evolutionary model proposed by Glikson (1972). Assuming 25% partial melting, the 200000 km 3 of basaltic material now absent from the Warrawoona Group could have contributed only about 5% of the granitic rocks in the batholiths, the total volume of which exceeds 900000km 3 (Hickman, in press). This does not rule out the possibility that pre-Pilbara Supergroup mafic crust was the source, although no remnants of such a crust occur on the surface, and seismic profiles (Drummond, 1979) reveal no evidence of such material or its partial-melt residue in the block's present crust. Stages in the. Evolution of the Pilbara Block The evidence suggests that none of the models used to explain crustal evolution in other parts of the world can be satisfactorily applied to the Pilbara Block. Rather than invoke yet another model, it is more appropriate to interpret the available information in terms of a stage-by-stage development of the area (Fig. 2). Ages designated to these various stages (below) are based on currently available geochronological data (Table I). As mentioned previously, most uncertainty relates to the precise ages of stages 5 and 6, and more work is being undertaken in this area. (1) 3900-3600m.y. The existence of 3800 m.y. granitic, volcanic and sedimentary rocks in Greenland and Minnesota establishes that greenstone-belt-type volcanism, sedimentation and granitic intrusion could have preceded deposition of the Pilbara Supergroup. Relics of rocks formed at this time may be included in the mega-xenoliths and gneissic
rocks of the batholiths, although they remain to be positively identified. (2) ca 3550 m.y. Tholeiitic basaltic rocks (Taiga Taiga Subgroup) were extruded (partly subaqueously) over most of the Pilbara Block (Fig. 2A). Thin sedimentary and rhyolitic units could have been derived by erosion and partial melting of the early sialic crust. (3) 3550-3450m.y. Local deformation was followed by the development of calc-alkaline volcanic centres (Duffer Formation) in a generally shallow-water environment. Intrusion of granodiorite into the underlying sialic crust and the Taiga Taiga Subgroup probably occurred at this time (Fig. 2B). (4) Post-3450, pre-2950m.y. (probably ca 3400 m.y.) Stable conditions prevailed during which time a 4 km thick tabular succession of pillowed tholeiite, high-Mg basalt, peridotitic komatiite and chert units (Saigash Subgroup) was deposited over the Pilbara Block. Part of the succession included evapprites and shallow-water carbonates. Felsic volcanism with associated chert and volcanogenic sediments (including local conglomerate) interrupted the accumulation of the basaltic pile in most areas (Fig. 2C). (5a) 3400-2950m.y. The process of "diapiric" (solid-state) doming and synclinal sinking involved progressive gravity deformation over a long period of time (probably several hundred million years). Some parts of the craton became unstable before others, probably because of regional variations in the thicknesses of both the granitic substratum and the supracrustal Pilbara Supergroup. Domes and synclines therefore commenced development at different times in different areas. Evidence is provided by local unconformities in the Pilbara Supergroup (Table I), and by the distinctive chemistry of individual batholiths (Hickman, in press). (5b) Partial melting of the granitic substratum resulted in extrusion of rhyolite (Wyman Formation). Erosion of the uplifted granitic substratum and folded sections of the Warrawoona Group produced sandstone, shale and turbidites (Corboy Formation and Soanesville Subgroup) outside early uplifted zones. (5c) Peneplanation was followed by submergence and deposition of banded iron forma-
CRUSTAL EVOLUTION IN PILBARA BLOCK
67
tion (Cleaverville Formation) across the currently believed to be 2950 m.y. (the main greater part of the Pilbara Block (Fig. 2D). metamorphic event recognized by Oversby, 1976). (6) Deformation recommenced and large areas (8) Post-2950 m.y. of granitic terrain were again exposed to Erosion of uplifted areas was followed by erosion. In certain areas the Cleaverville intrusion of post-tectonic granitic plutons Formation was also eroded. Considerable ("tin-granites"). At Whim Creek there was thicknesses (up to 5 km) of quartzofeldisland-arc volcanism (Whim Creek Group) spathic sandstone, grit and conglomerate with associated intrusion of porphyry (Fig. (Lalla Rookh Sandstone) were deposited in 2F). . sub-basins between Croydon, Shay Gap, (9) Open folding and strike-faulting in the Budjan Creek and Eastern Creek. Rapid Whim Creek area was followed by erosion downwarping, which occurred in a northand extrusion of basalt and andesite (Negri east trending basin (or trough) between Volcanics). Nunyerry and Port Hedland and in an eastnortheast trending basin east of Nullagine, ACKNOWLEDGMENTS resulted in deposition of turbiditic sediI thank the Director of the Geological Survey ments (Fig. 2E). of Western Australia for permission to publish (7) 2950m.y. this paper. Some of the basic information was Infilling of the Gorge Creek Group basins obtained by S. L. Lipple, R. Thom and R. J. was followed by the major episode of Chin during regional mapping, and constructive deformation that completed the develop- criticism and advice was provided by J. G. Blockment of most of the dome and syncline ley, R. D. Gee, and R. Davy. However, the writer structures. The age of this deformation is is entirely responsible for the ideas expressed. REFERENCES
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dating of the North Star Basalt, Warrawoona Group, Pilbara Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 187-192. HICKMAN, A. H., 1975: Precambrian structural geology of part of the Pilbara region. Ann. Rep. geol. Surv. West. A ust. for 1974, 68-73. , in press: Geology of the Pilbara Block and its environs. Bull. geol. Surv. West. Aust., 127. HICKMAN, A . H . , HORWITZ, R . C . , DUNLOP, J . S. R., &
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notes on the Marble Bar 1:250,000 Geological Sheet, Western Australia. Rec. geol. Surv. West. Aust., 1974/20 [unpublished]. HORWITZ, R. C., 1966: Notes on the legend of the geological map of Western Australia 1966. Ann. Rep. geol. Surv. West. Aust. for 1965, 46-41. INGRAM, P. A. J., 1977: A summary of the geology of a portion of the Pilbara Goldfield, Western Australia; in McCall, G. J. H. (Ed.) The Archaean, Search for the Beginning, 208-216. Dowden, Hutchinson & Ross, Stroudsburg, Pennsylvania. KRIEWALDT, M. J. B., 1964: Dampier and Barrow Island, Western Australia—1:250,000 Geological Series. Explan. Notes, geol. Surv. West. Aust., SF50-1, SF50-2.
C R U S T A L E V O L U T I O N IN P I L B A R A BLOCK KRIEWALDT, M . J . B . , & RYAN, G. R . , 1967: Pyramid,
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A REVIEW OF THE STRUCTURAL EVOLUTION AND GEOCHRONOLOGY OF THE ARCHAEAN NAPIER COMPLEX OF ENDERBY LAND, AUSTRALIAN ANTARCTIC TERRITORY P. R. James & L. P. Black 1
2
Department of Geology, University of Adelaide, GPO Box 498D, Adelaide, South Australia 5001 Bureau of Mineral Resources, Geology and Geophysics, PO Box 378, Canberra City, ACT2601 1
2
ABSTRACT
The Archaean craton of Enderby Land (Napier Complex) like other Archaean cratons is surrounded by younger mobile belts, but unlike most well-described granite-greenstone cratons is characterised by predominantly high-grade granulite-facies gneisses. The characteristic and complex pattern of tectonothermal evolution of the Enderby Land Craton bears a remarkable similarity to that of other high-grade Archaean terrains (e.g. North Atlantic Craton). In Enderby Land, a very well developed horizontal layering and a parallel intense flattening and extension (LS) fabric is the prevailing characteristic in what may be termed a recumbent gneiss terrain. The large-scale layering is made up of thick units of paragneissic supraerustal metasediments of largely epicontinental style (quartzite and garnet-quartz-feldspar gneiss: Tula Series) interlayered with further finely laminated acidic, intermediate and mafic granulite of more indeterminate origin and also with homogeneous charnockitic orthogneiss (Raggatt series). It is clear that this layering reflects a modification by intense late-stage horizontal tectonic activity, of a combination of sedimentary bedding, discordant and concordant magmatic intrusions and early structures. The tectonothermal evolution subsequent to the development and interlayering of the various units reveals a complex series of superposed deformation events of varying intensity and style, including the production of overprinted tectonothermal fabrics in a largely highgrade environment. This culminated in the production of a dome-and-basin pattern which characterises the craton. The evolution of the terrain concluded with the development of retrogressive mylonitic shear zones, faults and the intrusion of a suite of basic dykes. Isotopic resetting, which has apparently occurred during the discrete events producing recognizable penetrative fabrics, has allowed U/Pb zircon and Rb/Sr age estimates of 3000 Ma and 2500 Ma for the early horizontal tectonism and the culmination of the dome- and basinforming events respectively.
INTRODUCTION The first main mapping of the Enderby Land area was undertaken in the nineteen-sixties by USSR geologists. They recognised the unity and distinctive granulite-facies metamorphic character of the Napier Complex, subdividing it on lithostratigraphic grounds into a lower and older predominantly orthogneissic "Raggatt" series and an upper and younger predominantly paragneissic "Tula" series, though providing little information regarding the structural and tectonic features within and between these "series" (Kamenev, 1972, 1975; Ravich & Grikurov,< * Bureau of Mineral Resources, Geology and Geophysics
Spec. Pubis geol. Soc. Aust., 7 (1981)
1976). From this work the great antiquity of the Napier Complex was first suggested by —4000 m.y. U-Th-Pb ages of Sobotovich et ah (1976) from the Fyfe Hills near the western margin of the complex. Since 1974-75 BMR* geologists have undertaken systematic reconnaissance mapping of the whole of Enderby Land and have outlined the petrography and structure of the Napier Complex and have emphasised the unique nature of the metamorphic assemblages present (Ellis et al., 1980; Sheraton et al., 1980). They have been able to confirm the unity of the Napier Complex as a
72
P. R. JAMES & L. P. BLACK
high-grade Archaean craton by their assertion that the unmetamorphosed Amundsen dykes are restricted to the complex (Sheraton et al., 1980). We follow the usage of Sheraton et al. in distinguishing the Raggatt and Tula "series" purely on lithological grounds. It is not possible, either from the BMR work cited above or from this study, to determine the stratigraphic relationship between the two series and confirm or deny the assertion of Ravich & Grikurov (1976) that the orthogneissic Raggatt series comprises 4 'a lower part of the section" than the paragneissic Tula series. The geochronological study of Sobotovich et al. (1976) appears to support an old age for the Raggatt series but there is some doubt about the validity of the geochronological methods employed (see Grew & Manton, 1979). However the Raggatt series appears to overlie the allegedly younger Tula series at the outcrop we discuss in detail and elsewhere in the Napier Complex. STRUCTURAL EVOLUTION OF THE NAPIER COMPLEX Tectonic Pattern The unique tectonic features which characterise the Napier Complex as a distinct cratonic unit are well exposed around Amundsen Bay (Fig. 1). Most important is the development, early in the structural evolution, of a dominant compositional layering formed on all scales (mm-km in thickness) as a product of a complex evolution which included an intense flattening component perpendicular to the layering. The second characteristic is the typical horizontal nature of this compound layering which developed, probably at the culmination of the regional granulite-facies metamorphism, a largely recumbent gneiss terrain. This layering was subsequently affected by a simpler deformation controlled by subhorizontal compressive stresses which have buckled the gneiss pile into large-scale, open, upright noncylindrical folds. This style of tectonothermal evolution bears striking similarities to those described from other Archaean high-grade metamorphic terrains e.g. West Greenland (Bridgwater et al1974; James, 1975) and Scotland (Sheraton et al1973) in the North Atlantic Archaean craton, and the Limpopo belt in southern Africa (Coward et al 1976). In the Napier Complex, paragneisses* of the *
Tula series (for lithologies see Fig. 1) are structurally and compositionally distinct from the orthogneissic Raggatt series with which they are intimately associated on a large scale (Fig. 1). The origin of this association, and of the early major structures, is difficult to ascertain because of the lack of continuous outcrop, the recumbent nature of the structures, and intense parallelism caused by the earliest recognisable deformation termed here Dy. Stereoplots of the layering (S0) and layer-parallel fabric (S lv L1? F]) in Figure 1 largely reflect reorientation by later events. The resultant gneissic pile was "crumpled" by smallscale, consistently asymmetric, tight folds which appear to postdate D\ closely, but due to their uniform style and orientation have been separated into a D2 deformation. Granulite-facies conditions prevailed during but this event failed to produce pervasive, axial-plane, tectonothermal fabrics. The large-scale tectonic pattern of the Napier Complex (Fig. 1) is a typical dome-and-basin fold (interference pattern type 3 of Ramsay, 1967) of upright, non-cylindrical (doubly plunging) major folds. Refolding of the earlier structures and the cylindrical nature of D3 in individual mountainsized outcrops are shown on the map and stereograms of Figure 1. However, it is difficult to assess whether the final pattern is a product of pre-D3 non-planarity, D3 inhomogeneous deformation (i.e. anticlastic bending) or even subsequent refolding. Development of the Layering and Recumbent Gneisses (Dj and D2) One of the major problems in interpreting Archaean gneiss terrains is the process of formation of the gneissic layering, including the juxtaposition of orthogneissic and paragneissic units. Suggested mechanisms have included tectonothermal modification of supracrustal or mixed supracrustal and magmatic sequences; lit-par-lit intrusions or technically imposed parallelism of originally discordant intrusions of acid magmas into supracrustal piles; deformation and imposition of parallelism on basement (orthogneissic) and cover (paragneissic) unconformities, implying much older histories; and finally tectonic interleaving on various scales. The difficulties in providing evidence to support these processes typically result from the general paucity of continuous outcrop and the frequent flat-lying nature of structural regimes which therefore require considerable relief to reveal the diagnostic
N.B. The terms paragneiss (variably coloured, usually strongly layered gneiss) and orthogneiss (homogeneous and usually grey, pyroxene gneiss) are used here in non-genetic sense to simplify the description of the lithological types. The origin of each layer as igneous or sedimentary is discussed fully in the text after description of the field relationships.
STRUCTURAL EVOLUTION AND GEOGHRONOLOGY, ENDERBY LAND
73
NAPIER MTS. ^
TULA
MTS.
i i H q r d ^ g p i
Amundsen Bay feeaver
Khamara Bay /
Mt. McLennan SCOJT
MTS
McNaughton Ridge
RAGGATT
MTS.
STEREO MAP - P LOT
•. X
Major rock outcrops
2 0 p i p and strike of undifferentiated layering and layer parallel foliation
Xtfflffi) M r
Rock outcrops visited in this study
A
Direction and plunge of undifferentiated L ( elongation and Ft minor fold axis
Boundary of Napier complex
X
Axial plane of F, minor fold
Trend lines and Tula/ Raggatt boundaries (modified after Kamenev 1975)
4
Fl 2 minor fold axis (S or Z denotes vergence)
*
F 3 minor fold axis
•
1 Undifferentiated S3 schistosity and axial plane of F2 minor fold 1 ^ *
^
F 3 major axial surface trace, antiform or synform with plunge direction and major culmination (C) and depression (D)
Tula series —finely laminated mesoperthetic quartzo feldspathic gneisses garnet, sillimanite hypersthene and cordierife, quartz-rich gneisses, quartzites and more mafic pyroxene plagioclase gneisses Raggatt series-homogeneous hypersthene bearing granitic gneisses
Undifferentiated axial traces
Fig. 1.
Location and geological map of the western Napier Complex, East Antarctica.
P. R. JAMES & L. P. BLACK
74
2 a ,
see
2c
Pegmatite
—
Talus
FJ-P5
Tectonic slide (?) Paragneissic layers
White marker horizon
0,-0 6
Orthogneissic layers
S
Orthogneissic sheets with sinistral discordance
NV
Net vein migmatites with F| fold hinges
F'2 axial plane dislocation
F 2 axial plane dislocation
strongly layered para gneiss
SCALE =1=2400 (approx.)
Fig. 2.
Details of the NE face of Mount McLennan.
75 recumbent axial surface; and there is a noticeable dearth of recognisable parasitic folds or vergence variation. An interesting feature of this major fold is the lack of thinning of the layer P in the lower limb although the 350 m sequence above the Fi axial surface appears to be represented by less than 150 m below the axial surface, and none of the lower sequence of the lower half of the cliff (P3_5 and O4.6) is repeated. This evidence strongly suggests that the boundary between the upper and lower halves of the cliff may comprise an important semi-brittle or brittle tectonic slide as well as including an element of ductile thinning close to the contact (i.e. shown by inverted layers Pi and Oi). Conversely the Pj layer illustrates the form of the F folds, which are tight (up to 15° dihedral angle) and inclined. They refold the layering and Di flattening fabric without further transposition,^ formation of axial-plane fabric or marked hinge thickening, and show a distinct asymmetry of over-turning to the southeast with welldeveloped parasitic folds of consistent vergence. The orthogneissic units (0p0 ) of the upper half of the cliff (Fig. 2) are typically homogeneous although locally, for example in layer 0 , they exhibit a distinct but poorly developed stratification. On a large scale, they retain a regular thickness and overall continuity of individual horizons where folded, which suggests that the whole of this upper sequence must have been largely concordant, before Dj. The orthogneiss units in this face display no obviously older or younger,or more-or-less complex structures than the paragneisses, and there is thus no reason to invoke either a basement-cover relationship or intrusion of orthogneisses into already deformed paragneisses. On the basis of this evidence the orthogneiss units therefore could represent original contemporaneous acid volcanics and/or arkosic sediments of suitable composition in the overall stratigraphic succession. However, the possibility that at least part of the orthogneiss units are large-scale intrusives, or tectonic-sheets> cannot be discounted. In places, smaller-scale sheets of orthogneiss interrupt the large-scale continuity of the paragneissic layers across the face, whereas the boundaries of the paragneiss are commonly invaded, disrupted and, in places, completely transected by apparently intrusive grey orthogneiss. Significantly these intrusions, rather than having a random form, commonly appear to bear a systematic orientation relationship to the paragneiss layers. Most are in the form of concordant sill-like or slightly discordant dyke-like sheets up to 10-20 metres in thickness. Figure 2c, an enlargement of layer Pi on the southeast peak, illustrates the
STRUCTURAL EVOLUTION AND GEOCHRONOLOGY, ENDERBY LAND
information. In Enderby Land, mountains on the flanks of major glaciers commonly contain wellexposed steep cliff sections up to 1 km high and nearly normal to the recumbent layering (and often near-parallel to the profile planes of early and late folds). The NE face of Mount McLennan on the southern side of Amundsen Bay, shown in Figures 2a and b, probably provides the best such section in this area, offering an excellent illustration of geological relationships and demonstrating the various processes involved in the tectonic evolution of typical, high-grade, Archaean crust. The section illustrated in Figures 2a and b can be subdivided into upper and lower halves with individual sequences which display markedly dissimilar lithological and geometrical characteristics. The upper half, which consists essentially of Kamenev's Raggatt series (Kamenev, 1975), is dominated by homogeneous grey orthogneiss with subordinate, relatively thin, recumbently folded, white units. Conversely, the lower sequence comprises thick, well-laminated, continuous, paragneiss units with a distinctive redweathering, garnet-rich upper horizon and only minor orthogneiss, typical of the lithological character of the Tula series of Kamenev. The two series in this exposure appear to be separated by a tectonic discontinuity. The orthogneissic sequence of the upper part of the cliff (Fig. 2) can be subdivided into a number of distinct orthogneissic ( 0 0 ) and paragneissic (Pj and P2) units. These paragneissic units comprise garnet-sillimanite gneiss with fine (rammetre) internal compositional layering which is regular and parallel to the unit boundaries; a strong layer-parallel tectonite fabric also parallels the boundaries and the weak foliation in the external orthogneisses. Not only are the units Pj and P persistent across the cliff for nearly 4 km (with repetition), but single thin layers within them can be traced for considerable distances; an example is the white layer (10-15 m thick) at the upper and lower boundary of Pi. Although no conclusive sedimentary structures remain, the lateral persistence and form of this layering, together with the lack of both small-scale isoclinal folds and transposed structures, strongly suggests largely unaltered sedimentary stratification. On a larger scale, the paragneissic layers Pi and P magnificently illustrate the styles of the two early fold phases (Fi and F ). Layer P , averaging 100 m in thickness, is repeated by a Di isocline (F ) closing to the southeast, with a grossly thickened (up to 500 m) hinge; layerparallel flattening fabric is shown by pinch-andswell and boudinaged schlieren structures strongly elongated parallel to the original (pre-F ) r
3
2
2
2
2
{
3
2
2
3
2
76
P. R. JAMES & L. P. BLACK
discordance of these sheets to both the Pi layer boundary and to the internal paragneiss lamination. If these sheets are traced from the upper orthogneissic layer Oj to the lower orthogneissic layer 0 2 , then all clearly cut across the paragneissic layer from an upper left (SE) to a lower right (NW) position (i.e. in a structurally sinistral sense). Surprisingly, this relationship seems to hold when the white layer is traced around both of the major fold closures (Fig. 2b). This evidence of folding around the fold without change of vergence clearly shows that the discordant intrusive layers were emplaced before Both the Fj and F 2 fold phases. What is not clear, however, is the relationship between these orthogneissic sheets and the major orthogneiss layers ( 0 e t c . ) . In addition, an intrusive origin for the orthogneisses in general is indicated by the common disruption of the P j and P 2 paragneissic layers. This can be clearly seen for P j on the upper right hand side of the southeast peak (Fig. 2b) and also around the major F 2 fold closure, where large blocky xenoliths of white layered paragneiss are rafted off and totally surrounded by the grey orthogneiss. Although the rafts are completely separated from their parent horizons (at least in the two dimensional section shown by this exposure) their internal layering has the same orientation. This strongly suggests that the intrusion of the orthogneissic sheets took place under the influence of a deviatoric stress, rather than in a hydrostatic or low differential stress regime, allowing the continuity of layering to be retained even during a period of major mobility. It is this evidence, together with the near concordance of the sheet-like intrusives, plus their systematic asymmetrical relationships to the layering, which suggest that the intrusion took place during or slightly preceding the climax of the first major recumbent folding event (Dj). The lower half of the cliff face of Mount McLennan contrasts markedly with the upper sequence. The lower levels are dominated by finely laminated paragneisses (P3-P5) containing subordinate layer-parallel sheets of homogeneous grey pyroxene gneiss (0 4 -0 6 ). This sequence is considered to be the lithological equivalent of Kamenev's Tula series. Thus at Mount McLennan the Tula paragneisses are overlain by Raggatt series orthogneiss (the reverse relationship to that cited by Kamenev, 1972) with the contact marked by a discrete slide or fault and an intense component of ductile shearing and/or flattening. The paragneissic units display evidence of their original sedimentary nature which resembles that described for the upper paragneiss. In the lower units, however, there appears to be a marked increase in the amount of layering developed by the
extreme folding, flattening and attenuation, probably during D { , of locally derived pegmatitic quartz-feldspar ± garnet net-vein migmatite. A few F j minor fold hinges of the originally discordant veins are visible (e.g. half way down the NW limb of the fold arch of layer P 3 ) (Fig. 2b, NV). However, the folded veins now mainly comprise boudinaged limb regions parallel to the layering and to similarly oriented axial planes. The reduced migmatitic development in the upper horizons may be due to the lack of suitable compositions to generate melts, or to the difficulty in observing vein material in the more mixed upper layers. Characteristic of the lower sequence are the sharp and continuous boundaries between all lithologies on all scales, with neither obvious transposed interdigitation nor evidence of magmatic intrusion. There is no rafting or intrusion of discordant sheets as in the upper sequence. It is even more likely therefore that much of the lower orthogneiss may represent autochthonous members of the supracrustal sequence, although early (pre-D ]) intrusion of sills or tectonic interleaving cannot be completely ruled out. Thus, on the macroscopic scale shown by this section of Mount McLennan, the lithological layering comprises a number of distinct morphological types which have formed in varying ways. On the largest scale (400-500 m) there is the layering (characteristic of the upper horizons) defined by the alternation of the thick pile of homogeneous, in part, obviously igneous orthogneiss containing semi-continuous sheets and pods of subordinate paragneiss, and the thick sequence of the lower horizons comprising well-layered and locally migmatised paragneiss. This type and scale of gross layering is repeatedly exposed on other large cliff faces of the mountains and nunataks of the area around Amundsen Bay. The boundary between the two major units consistently appears sharp, continuous and generally lacking in interdigitation. Other exposures are composed of predominantly orthogneissic "Raggatt" series lithologies, whereas some are composed of very thick sequences of strongly layered paragneiss only. No field evidence is provided by Kamenev (1975) to suggest whether the Raggatt series orthogneiss is truly orthogneissic, intruding "earlier" paragneiss of the original Tula sedimentary pile, or whether the homogeneous Raggatt series rocks perhaps wholly or in part formed a basement on which the Tula sediments were deposited. The lack of geological evidence of a significant age difference between the two series tends to support the former. What is clear from the map (Fig. 1) and from the evidence of Mount
STRUCTURAL EVOLUTION AND ( EOCHRONOLOGY, ENDERBY LAND
McLennan, is that the major lithological series were intimately interleaved into parallel interdigitating sheets on an enormous scale before the formation of most of the earliest recognisable structures. There are a number of possible explanations. Firstly, the Raggatt series may have formed as a result of abundant large-scale intrusions either as discordant bodies which were later deformed into parallelism, or as concordant sheets. Secondly the orthogneiss may not be wholly intrusive in origin, but may belong to the same metasedimentary pile as the Tula sediments. As shown on Mount McLennan (Fig. 2) the orthogneiss is mostly concordant and laterally persistent. It is also internally layered in several places and therefore may have originally been made up of greywacke-like sediments and/or acidic volcanics. However, it is also associated with orthogneiss sheets which undoubtedly intrude and raft off blocks of paragneiss. The precise relationship between the intrusive orthogneiss and the main body of (Raggatt series) orthogneiss remains unclear and it is not known whether their juxtaposition resulted from a major intrusive episode before Dj or major thrusting and tectonic interleaving, before or during the early stages of Dj. Whatever the relationships within the series, evidence suggested by the Tula/ Raggatt relationship exhibited at Mount McLennan, and elsewhere in the Amundsen Bay area, indicates that the juxtaposition of the two series was essentially a tectonic process. On a smaller scale (both mesoscopic and microscopic), distinctive lithological layering commonly observed within the major subdivisions of the Napier Complex and described on Mount McLennan, is thought to reflect the primary bedding stratification of the sediments from which the paragneiss was derived, though the layers have obviously been much affected and some totally overprinted by the early intense deformations. Repetition and transposition of this layering by D isoclinal folding is often difficult to differentiate from the extreme flattening of the common net-vein migmatites. {
Early (Dj /D ) Microfabric Development Throughout the Napier Complex, Dj produced a remarkable consistency of microstructure (i.e. grain sizes and grain-boundary configurations). Microstructures are generally equigranular and granoblastic in both orthogneiss and paragneiss with grains between 2-4 mm in diameter. This uniformity of fabric is however only the final product of the whole deformation, and its longevity and complexity is evidenced by details of microstructure illustrated by and within particu2
77
t lar mineral phases, especially within garnet crystals. There are microstructural variations due either to compositional differences or to differences in the intensity of deformation. The two major lithological variants, the orthogneiss and paragneiss, are dominated by different mineralogies, often showing varying styles and intensities of Dj fabrics even in adjacent areas. The orthogneiss typically does not develop strong fabrics. It is composed of varying proportions of mesoperthitic potash and/or plagioclase feldspar, with subordinate quartz and minor orthopyroxene. Quartz and mesoperthitic feldspar crystals are interstitial, granular or amoeboid. Grain boundaries are typically low energy (viz., straight, slightly curved or rational) and inclusions are uncommon. Single quartz grains or aggregates of two to three grains show a weak elongation and dimensional preferred orientation parallel to the foliation, whereas feldspars are typically less elongate and commonly consist of equant granoblastic aggregates with straight grain boundaries and 120° triple junctions. These grains typically show no intracrystalline deformation features. Weak streaky fabrics are formed by parallelism of elliptical aggregates composed of small orthopyroxene grains separated by interstitial leucocratic mineral grains which are normally quartz but include some feldspar. The aggregates are up to 10-15 mm in length and 4-5 mm in width and appear to have developed by a process of fracture and break-up of originally larger orthopyroxene grains. After brittle rupture, the relict fragments of orthopyroxene apparently have separated by either extension across the fractures (now new grain boundaries) or sliding along them. The quartz and feldspar have filled in the intervening areas by either diffusion or grain boundary sliding and readjustment. The orthopyroxene fragments show little dimensional, and no crystallographic preferred orientation, suggesting that there was no syntectonic orthopyroxene growth during Di, and restoration of the original pre-Dj orthopyroxene grains shows them to have been up to five times coarser than the average grain size of the present D\ microstructure. The paragneiss provides the best-developed examples of Dj tectonites mainly due to its varied mineralogy which has reponded differently to the imposed deformation, resulting in complex microfabric relations. These microfabrics also reveal to a greater extent the length and complexity of the Di microstructural changes. Overall, equigranular to inequigranular granoblastic polygonal aggregates with equilibrium grain boundaries are the main microstructures. An LS •
78
P. R. JAMES & L. P. BLACK
(LjSj) tectonite fabric is variably developed, the most common individual fabric elements being elongate quartz ribbon aggregates, elongate aggregates of mafic minerals, and sillimanite laths. The quartz aggregates were the most ductile. In quartz-rich rocks and garnet-sillimanite quartzite, grain-boundary migration parallel to the foliation is not inhibited by other minerals, and large quartz-ribbon aggregates reach dimensional elongation ratios of 10:1 parallel to the elongation lineation Lj. Within the ribbons, individual quartz grains show no obvious crystallographic or dimensional preferred orientation. In spite of the obviously large strains suffered by these rocks, internal crystal deformation features such as strain-shadowing or deformation bands are rarely developed. Feldspar grains are generally only weakly dimensionally elongated parallel to the S\ fabric, although original large grains commonly recrystallize to elongate granoblastic aggregates of equant unstrained grains. Dimensional grain or grain aggregate elongation ratios, however, rarely exceed 2:1. Apart from quartz ribbons, the best fabricforming elements in most of the paragneissic lithologies are sillimanite grains. These develop to varying sizes, generally 2-4mm in length. They are lath-like, elongate, idioblastic and aligned in excellent preferred orientations parallel to Lj. The sillimanite grains parallel the axial planes of rare Fj isoclinal folds and appear to have grown syntectonically throughout Dj. Deformation features are rare and sillimanite, the only indicator of the Sj fabric in many paragneisses, appears to have continued growing after the cessation of D\ movements. Mafic minerals such as pyroxene and garnet typically show little tendency to elongate or develop a preferred orientation. The garnets, which abound in most paragneissic lithologies, are, however, excellent indicators of the style, complexity and intensity of the Dj fabric development. They are probably the most interesting crystals as they apparently grew (and survived) throughout Dj. The garnets are even grained at 1-4 mm in diameter and are the most competent minerals with respect to the deformation. They are typically subrounded-subidioblastic although in high D{ deformation areas some may be flattened and aligned parallel to S r . This shows that many of the garnets are syn- or pre-tectonic. However, many other garnets in the paragneiss are poikiloblastic to some extent, containing inclusions of variable grain size, usually of quartz or sillimanite. The patterns of these inclusions reveal that
garnet growth was complex. Many garnets show a concentric zonation of both the proportion and grain size of inclusions. These patterns, which are rarely simple, suggest that the garnets grew at all stages of D\. Some garnets have inclusion-free centres, then small zones of inclusions followed by outer inclusion-free margins; others have numerous inclusions in the centres then inclusionfree margins, whereas others, even in the same thin section, may be completely inclusion-free. Although masked by further complexities of the inclusion-trail development, there appears to be a general tendency for the inclusions closer to the centre of garnets, which are commonly quartz or sillimanite, to be very fine grained, and for the grain size to increase gradually toward the margin of individual garnets; at the margins, matrixsized material is often partially included resulting in an irregular garnet outline. This suggests that the garnets grew over a comparatively long time span, during the gradual coarsening of the paragneissic microstructure. This is of further interest when compared with the orthogneissic lithologies which, from the evidence of their streaky orthopyroxene fabrics, appear to have been reduced in grain size to reach the very uniform 2-4 mm dimension of present Dj granoblastic grains. There is a further microstructural complexity in that, though many of the inclusions within the garnets are rounded and spherical, many others are non-equidimensional and aligned to produce an internal fabric. Some garnets contain a central core with very fine elongate inclusions in straight, parallel arrangements randomly oriented with respect to adjacent garnet inclusion trails and the external Sl fabric. These central core inclusions suggest that there was an original high-grade (pre-Dj) fabric in the finer grained precusor to S v Other inclusion trails, especially those formed by larger inclusions nearer to the margins of garnets, tend to be curved, wrapping around the central parts of the garnets and paralleling the external Sj which itself wraps around the whole garnet. These are typical of irrotational syntectonic growth of garnets which is considered to have taken place during the later stages of the D} flattening. Other garnets contain poorly developed internal inclusion trails reminiscent of those in various types of syntectonic snowball garnets described by Powell & Treagus (1970). The development of the Sj fabric and microstructure comprises a long and complex process involving solid-state syntectonic recrystallization causing very large ductile strains at high temperatures and pressures and probably at very slow strain rates. This appears over a large area to have resulted in a very uniform microstructure which developed from a microstructural conver-
STRUCTURAL EVOLUTION AND GEOCHRONOLOGY, ENDERBY LAND
79
gence from coarser grained orthogneiss and pervasive axial-plane schistosities are faintly generally finer grained paragneiss. visible as streaky ill-defined foliations cutting These microstructures further indicate that across the layering; associated microstructures solid-state grain boundary readjustment con- are described below. tinued after cessation of the final Dj tectonic The few discrete shear zones which are asmovements, removing any crystal defects or signed to this event are generally thick upright deformation features. zones of anastomosing black mylonite and ultraThe Si fabric is folded by F folds without new mylonite which cut across the layering. They may S2 axial-plane development or significant intra- be parallel or conjugate to the D axial plane. crystalline deformation. This, together with the External foliations curve into the zones where earlier indicated close relationship between Dj anastomosing shears progressively develop protoand D2, suggests that final microstructural equi- mylonite, slaty and schistose mylonite, and black, unlayered and poorly-foliated ultralibrium was attained at the culmination of D . Streaky lineations on the mylonite foliSubsequent Evolution of the Gneissic Pile (D ) mylonite. ation typically plunge steeply to the northeast or The third recognisable tectonothermal event in southwest. the Napier Complex produced a distinct and uniform character resulting from a fundamental D Microfabric Development change in the style and intensity of imposed The D microfabrics show a general heterostress, as well as a change in the environmental geneity, being in part retrogressive and in part conditions. As outlined earlier and illustrated in non-retrogressive. There are two main types Figure 1, the subhorizontal compression which associated respectively with (i) the pervasive, operated during D produced the characteristic though still inhomogeneous, tectonite fabric; and dome-and-basin fold-pattern of the craton with (ii) an inhomogeneous development of discrete the style of individual folds probably again most and partly retrogressive mylonitic shear zones. clearly shown on the face of Mount McLennan The pervasive schistosity and mylonite schistosity are the two end members of a spatially grada(Figs 2a, b). D3 folds are typically large-scale, upright struc- tional sequence, the shear zones being best tures which in the north of the Napier Complex developed to the north of Amundsen Bay, and the (Fig. 1) trend NNE-SSW, swinging through NE- pervasive schistosity gradually increasing in SW further south to near E-W on the southern intensity of development toward the south. boundary of the Napier Complex. It is possible These different types of fabrics, although that cross folding associated with this swing is morphologically variable, are correlated with the responsible at least in part for the formation of D event as they all develop as planar structures the dome-and-basin pattern. The variations of at high angles to the layering and layer-parallel the D major folds defining the pattern are fabrics, and typically form parallel to the axial regular with major south-plunging areas to the planes of D folds. They are also similar in that north of Amundsen Bay and north-plunging the micro-structures are characterised by intraareas to the south. These plunge variations can be crystalline deformation features and recrystalgauged both from large-scale folds and inter- lization. This results in rocks of variable and section lineations and also from rare mesoscopic reduced grain size and contrasts markedly with the uniform Dj microstructure. F folds. F folds are uncommon on the scale of single As the D deformation on the microscopic outcrops (cliff faces or smaller), Mount McLen- scale is defined by intracrystalline deformation nan (Fig. 2) being an exception. This is a result of and syntectonic recrystallization without signifithe open nature of the typically 1-2 km wave- cant retrogression, the form of the microfabrics length F folds which do not have smaller para- developed depends not only on the intensity of sitic structures, but simply develop gently curved the deformation, but also on the resistance of the shallow-dipping limb regions and curved hinges. various minerals to recrystallization. A few measurable minor F folds however are In areas where D effects are least intensely upright with planar axial surfaces either vertical developed (i.e. north of Amundsen Bay), they or consistently, steeply, northwesterly-dipping. take the form of simple intracrystalline deformain the various minerals. Typically D was also responsible for the inhomogeneous tion features is no recrystallization of the equigranular generation of a variety of pervasive and non- there granoblastic microstructure and, overall, pervasive, high-grade, axial-planar tectonite deformation Sj are weakly developed. fabrics and discrete, retrogressive mylonite- Quartz grains.,features still in Sj, are typically developing shear zones, all of which cut discor- unstrained or show elongated strain shadowing, and the indantly across the recumbent gneiss pile. The more 2
3
2
3
3
3
3
3
3
3
3
3
3
3
3
3
3
80
P. R. JAMES & L. P. BLACK
cipient development of deformation bands is usually at high angles to S]. Plagioclase feldspars are generally untwinned or show the weak marginal development of wedge-shaped, lamellar, deformation twins, which may themselves, along with the grain as a whole, be slightly bent. Other minerals such as orthopyroxene or garnet show no internal deformation Teatures, thus illustrating their greater competence during deformation. South of Amundsen Bay, the D3 microfabrics tend to be better developed with more intense intracrystalline deformation, recovery structures and subgrain development, and minor recrystallization which is non-retrogressive but heterogeneously developed. This heterogeneity shows as small areas where intracrystalline deformation structures predominate, transected by narrow (up to 1 cm wide) zones, generally parallel to the D3 axial plane, where intense recrystallization occurs without significant retrogression. The intracrystalline deformation in quartz grains is shown by well-developed deformation bands. The bands may be planar and inclined at moderate to high angles to the Si elongation of the grains, or may develop as a network of orthogonal intersecting bands, which typically leads to the development of subgrains at the intersection of the deformation bands: This is caused by the high lattice rotation at these points, and eventually leads to the development of new grains (recrystallization). Plagioclase grains develop complex albite and pericline lamellar twins which are wedge-shaped and obviously a product of intracrystalline deformation. The twins may be bent or kinked with angles of bending varying from 2G-10°. Beyond 10° the plagioclase grains are not able to accommodate the lattice rotation by simple bending and high-angle, discrete grain boundaries appear, a precursor to eventual recrystallization. Some twinned plagioclase crystals develop new, recrystallized, fine plagioclase grains (up to .1 mm) at the end of an old grain, where the formation of the new grains appears to be controlled by marginal lattice rotation of more than 10°. Even with the high intensity of D3 intracrystalline deformation, orthopyroxene grains remain largely unaffected. Some locally show the weak development of planar kink bands with lattice rotation of only 2-3 In the areas of high intracrystalline deformation and overall weak recrystallization, thin seams of intense recrystallization typically form parallel to the D3 axial plane. In these seams the grain size is effectively reduced by recrystallization to an average of .08-. 1 mm, with the remnant grains and the new grains both elongated parallel to the seam, forming a new S3 fabric. The re-
crystallized mineral assemblage is the same inside and outside the shear zone. The formation of the shear zones is complex and most probably entails rotational strain (simple shear?), as the S3 fabrics curve in and out of elliptical pods of variably recrystallized material. Within the shear zones large, original garnet grains are deformed into elliptical shapes with long axes contained within S3 which is defined by parallelism of trains of finely recrystallized garnet, orthopyroxene, quartz and magnetite which wrap around the old grains. The most mobile minerals in the recrystallized aggregates appear to be quartz and magnetite. Magnetite forms as elongate finely recrystallized grains in the matrix, as pressure shadows around deformed orthopyroxene grains, or as an interstitial phase filling shear and extension fractures in old hypersthene or garnet grains. These fractures are typically aligned perpendicular to S3. South of Amundsen Bay, at Mount Tod, the S3 fabrics appear to be a more pervasive schistosity. In thin section, this fabric is defined by both the elongation of remnant feldspar and pyroxene grains, and by the alignment of deformation features such as kinks within these grains. It is further defined by parallel trains of new recrystallized grains, which wrap around the old grains, and which themselves although typically granoblastic, show a weak elongation and preferred orientation parallel to S3. Quartz grains typically recrystallize and the aggregates become very elongate parallel to S3, presumably by grain-boundary sliding. Feldspars show the most varied and interesting microstructures. Old deformed grains may be elongate in S\, or, where affected by deformation bands which recrystallize, they may be elongate in S3 as mentioned above. Bands and kinks may be sharp, but are usually outlined by curved lamellar deformation twins developed in most plagioclase grains. Kinks in plagioclase typically show lattice misorientations of less than 10° and may be planar where they are parallel or nearparallel to S3, or they may form as conjugate pairs with their acute bisector parallel to S3. With increasing strain, true deformation bands form parallel to S 3 , with the development of high-angle new-grain (and kink-band) boundaries showing misorientations up to 50°-60°. These deformation bands have become the locus of recrystallization of fine, equigranular, granoblastic, polygonal aggregates, where the new grains develop with straight grain boundaries and low-energy 120° triple junctions. The new grains which are very regular in grain size (approximately .1 mm), may show minor lamellar deformation twinning. Pyroxenes show a similar range of deformation
STRUCTURAL EVOLUTION AND GEOCHRONOLOGY, ENDERBY LAND
81
TABLE I
U-Pb analytical data for zircons from Mount Hardy (78285017) and Mount Sones (78285008) Samp1e
Samp1e fraction
Mag-
Mesh size
Colour
-75
n e t i c
suscep-
Concentration genie Pb
(ug/g)
206p 2 0 -
Total Pb
b /
20 7Pb/
p b
2 06Pb
measured
Atomi c Rat i o s 06 207pb/ Pb/ 2 3 5-u 2 3 8,U
pink
NMl
1
126.5
126.7
263.9
9180
.0022
16.226
.16197
.45387
-75
p i nk
Ml
2
.143.0
143.4
294.6
7555
.0052
16.295
6. 7 2 7
.16230
.45958
10.2848
-150
pink
NM1+M1
3
157-7
157.7
328.1
48780
.0007
16.218
4.456
.16209
.46328
10.3519
"
-150
p i nk
M2
4
143,2
144,1
307.5
5055
.0109
16.077
4.,836
.15940
.44980
9.8860
"
-150
brown
NM1+M1
5
1250
1251
2673 .
32020
.0021-
16.183
1 . .681
.16157
.46172
10.2861
"
-T50
brown
M2
6
1632
1634
3448
39970
.0018
16.190
1 . ,584
.16167
.46390
10.3411
-215
brown
.46521
10.3914
78285017 "
78 285008
" 1.
6. .676
NM2+M2
7
1526
1527
3240
53250
.0015
16.220
1 . .584
.16200
-45
NM
8
457.6
458.5
876.2
12080
. 0042
17.990
8..266
.17938
.48154
-75
NM
9
394.8
395.3
759.4
28760
. 0026
17.171
1 1 . ,894
.17139
.46847
1 6 . • 925
.16513
.46482
10.1363
11.9097 11.0707
-150
pi nk
NM
10
305.6
306.0
572.8
25430
. 0026,
16.546
-150
brown
NM+MO
11
1241
1242
2301
27630
.0023
19.568
3. . 4 1 9
.19540
.51005
13.7420
-250
pi nk
NM+MO
12
322.7
323.6
598.2
25950
.0029
16.250
1 9 . .481
.16215
746209
10.3308
-250
brown
NM+MO
13
1199
1200
2118
30500
.0017
20.550 •
2.• 909
.20529
.53336
15.0974
and recry stallization microstructures. Clinopyroxenes appear to have recrystallized very readily, whereas orthopyroxenes show a much greater resistance to both deformation and recrystallization. Some orthopyroxene grains with their cleavage at a sufficiently high angle to S3 are kinked, whereas others, with low angles between their cleavage and S3, deformed by internal slip and passive rotation into parallelism with S . Clinopyroxenes suffered some internal deformation such as bending, but apparently recrystallized very readily into granoblastic polygonal aggregates. The total amount of S3 recrystallization in these rocks is of the order of 40% in the mafic layers, whereas in the leucocratic, more plagioclase-rich layers, the total recrystallization is over 50%. In conclusion, the S schistosity is a clearly identifiable microfabric with a characteristic microstructure which appears to grade in intensity from north to south across the Napier Complex. In contrast to the equilibrium microstructures developed during the D /D tectonothermal period, D microstructures heterogeneously affected most lithologies by selective deformation and recrystallization. Within small areas, however, S development can be considered to represent new homogeneous equilibrium microstructures. 3
3
{
2
3
3
GEOCHRONOLOGY Isotopic dating is difficult in this region of superposed events because high metamorphic grade and generally dry conditions (Sheraton et al., 1980) have not allowed the regular production of overprinted schistosity or crenulation cleavage. The main task is to ascertain to which, if any, of the events a derived age relates. Partial
'
10.5832
resetting, in which an age has no direct reality, but lies between those of two events, is a major problem. Two sites in the Napier Complex (see Black & James, 1979) have so far yielded useful geochronological information and demonstrate the intricacies of isotopic-data interpretation in this terrain. U/Pb zircon data for these are recorded in Table I and Figure 3. Zircon analyses for leuconorite from Mount Hardy (78285017) plot near the concordia at about 2500m.y. which is considered to approximate the age of a major tectonothermal event. Derivation of a more precise age depends on which of two models is applied. Deletion of zircon fraction -150jimM2 (pink), under the likely assumption that it might represent an inaccurate analysis, produces an alignment in which concordancy is related to grainsize. The resultant age is 2 4 7 4 ^ . Alternatively, deletion of the two -75/rni fractions, which may have preferentially suffered diffusional loss of Pb since 2500m.y., yields an age of 2 4 8 8 m . y . Further analyses from this and other sites should show which of these models is more appropriate. Unlike the Mount Hardy site, isotopic composition is related to zircon morphology in paragneiss at Mount Sones (78285008). Coarse pinkish, rounded, and transparent zircons plot near, with the -250pm NM + MO fraction plotting right on the Mount Hardy trend. In contrast, two dark brown, translucent, and less rounded zircon fractions fall to the right with higher 207pb/ 206pb ages. These brown zircons have distinctly lower 208pb/206pb and hence Th/U ratios than the pinkish ones, at both Mount Sones and Mount Hardy. It is most likely that the different coloured zircons grew during separate events and the older (brown) ones were variably reset during the growth of the pinkish variety about 2500m.y.
P. R. JAMES & L. P. BLACK
82
Fig. 3 .
206
Pb 238 U - 2 0 7 Pb/ 2 3 5 U concordia diagram for zircon separates from Mount Sones (crosses) and Mount Hardy (circles).
ago. The -45/mi and - 7 5 ^ m fractions from Mount Sones, which because of their finer grainsize are difficult to categorise, plot between these groups. Extension of the Mount Sones trend cuts the concordia at about 3100 m.y. The nature of the cluster of zircon analyses near the 2500 m.y. point makes it probable that most if not all of the Hardy and Sones zircons have suffered minor Pb loss at that time. Approximate correction for this loss by an upward adjustment of the Mount Sones chord yields a probable growth or complete isotopic equilibration age for the brown zircons of about 3000 to 3050m.y. The 207pb/206pb age of 2869m.y. for 78285008 - 250 NM + MO (brown) is a minimum age for this event. Although the two sites Mount Hardy and Mount Sones are respectively assigned to the Tula and Raggatt series on Figure 1 the difference in the ages obtained in this study probably does not reflect any difference in age between these series. Rather it reflects evidence for different metamorphic episodes that have affected both series at both localities. Correlation of these ages with specific events is based on local and regional observations. The similar styles of Dj and D2 suggest that they represent closely spaced orogenies during a period of continuous granulite-facies conditions.
The approximately 3000 m.y. zircon age should then specifically represent the onset of a significant post-D 2 cooling and would probably also approximate the age of Dj. No folds were seen in the tholeiitic dykes which cut the Napier Complex. Sheraton & Black (unpublished data) have derived a Rb/Sr total-rock alignment which indicates that the oldest of these dykes were emplaced about 2400m.y. ago. Hence the 2500 m.y. event is believed to be D3 . Our study thus confirms our earlier assertion (Black & James, 1979), and that of Grew & Manton (1979), that a widespread metamorphic event affected the Napier Complex at about 2500m.y. It also clearly shows that this was merely the last episode in the formation of the Napier Complex, a process that began at least 500 million years before. It also shows yet again the value of a close association between geochronological and structural studies in complexly deformed, ancient gneiss terrains. ACKNOWLEDGMENTS We acknowledge the skilled technical assistance of M. J. Bower, T. K. Zapasnik, K. Barrett, and N. C. Hyett. R. J. Tingey, R. L. Oliver, A. D. Haldane and K. R. Walker critically read the manuscript.
STRUCTURAL EVOLUTION AND GEOCHRONOLOGY, ENDERBY LAND
The authors gratefully acknowledge the assistance and logistic support of the Antarctic Division Department of Science and the Environment while they were members of the Australian National Antarctic Research Expedition in the
83
1977-78 and 1979-80 Antarctic summer field seasons. The contribution of L. P. Black is published by permission of the Director, Bureau of Mineral Resources, Geology and Geophysics, Canberra, Australia.
REFERENCES
BLACK, L. P., & JAMES, P. R . , 1979:
Preliminary isotopic ages from Enderby Land, Antarctica [Abstr.]; in Tingey, R. J. (Compiler) Workshop on Antarctic Geology 17-18 May 1979. J. geol. Soc. Aust.,26, 266-267.
BRIDGWATER, D . , MCGREGOR, V . R . , & MYERS, J . S>,
1974: A horizontal tectonic regime in the Archaean of Greenland and its implications for early crustal thickening. Precamb. Res., 1, 179-197.
COWARD, M . P . , JAMES, P . R . , & WRIGHT, L , 1976:
Northern margin of the Limpopo mobile belt, southern Africa. Bull. geol. Soc. Am., 87, 601-611. V
ELLIS, D . J . , SHERATON, J . W . , ENGLAND, R . N . , & DALLWITZ, W . B., 1980: Osumilite-sapphirine-
quartz granulites from Enderby Land, Antarctica—mineral assemblages and reactions. Contrib. Mineral. Petrol., 72, 123-143. GREW, E. S., & MANTON, W . I., 1979: Archaean rocks in Antarctica: 2.5 billion year uranium-lead ages of pegmatites in Enderby Land. Sci., 206, 443-445. JAMES, P . R . , 1975: Field mapping of Bjefrnecfen and the adjacent coast of Nordlandet, GodthSbsfjord, southern West Greenland. Rapp. Grdnlands geol. Unders., 75, 58-62. KAMENEV, E . N . , 1972: Geological structure of Enderby Land;m Adie, R. J. (Ed.) Antarctic Geology and Geophysics, 579-583. Universitetsforlaget, Oslo.
, 1975: The geology of Enderby Land (in Russian). Acad. Sci. USSR, Comm. Ant. Res. Rep., 14. POWELL, D., &TREAGUS, J. E., 1970: Rotational fabrics in metamorphic minerals. Mineral. Mag., 37, 801-814. RAMSAY, J. G., 1967: Folding and Fracturing of Rocks. . McGraw-Hill, New York. RAVICH, M . G . , & GRIKUROV, G . E . (EDS), 1976:
Explanatory notes to the geological map of Antarctica, scale 1:5000000. Research Institute of the Geology of the Arctic, Ministry of Geology of the U.S.S.R., Leningrad.
SHERATON, J . W . , TARNEY, J . , WHEATLEY, T . J . , & WRIGHT, A. E., 1973: The structural history of the
Assynt District; in Park, R. G., & Tarney, J. (Eds) The Early Precambrian of Scotland and Related Rocks of Greenland, 31-43. University of Keele, Staffordshire.
SHERATON, J . W . , OFFE, L. A . , TINGEY, R . J . , & ELLIS,
D. J., 1980: Enderby Land, Antarctica—an unusual Precambrian high-grade metamorphic terrain. J. geol. Soc. A ust., 27, 1-18.
SOBOTOVICH, E . V . , KAMENEV, E . N . , KOMARISTYY, A. A., & RUDNIK, V. A., 1976: The oldest rocks in
Antarctica (Enderby Land). Int. geol. Rev., 18, 371-388.
ARCHAEAN ROCKS IN THE SOUTHERN PART OF THE CANADIAN SHIELD—A REVIEW P. K. Sims & Z. E. Peterman U.S. Geological Survey, Denver, Colorado, f/.S.A 50225 ABSTRACT Archaean rocks in the southern part of the Canadian Shield consist of two crustal segments of widely different ages. Greenstone-granite complexes (greenstone terrain), 2750-2600Ma, which compose the southern margin of the Superior structural province, are bounded on the south by older Archaean migmatitic gneisses (gneiss terrain), in part 3500 Ma. The type area for the greenstone terrain in the Lake Superior region—the Vermilion district of northern Minnesota—consists of a thick, steeply dipping, intertonguing accumulation of low-grade volcanic and sedimentary rocks intruded marginally by elongate calc-alkalic granitoid batholiths. The sequence is made up of an older (Ely Greenstone) and younger (Newton Lake Formation) mafic volcanic succession separated by dacitic volcaniclastic and derivative sedimentary rocks. The lower member of the Ely is composed mainly of amygdaloidal basaltic and andesitic lavas belonging to a calc-alkaline association; the upper member is mainly tholeiitic basalt. The Newton Lake Formation contains tholeiitic and komatiitic basalts, which grade laterally into calc-alkalic dacitic rocks. The rocks compose a regional anticline between the fringing batholiths, resulting from deformation caused by the diapiric rise and expansion of the granitoid rocks. The granitoid rocks have initial 8 7 Sr/ 86 Sr of less than 0.701, and appear to have been derived either by fractional melting of basic rocks or partial melting of greywacke with a short crustal residence time. The volcanic-sedimentary pile and intrusive granitoid rocks were developed in 50-100 Ma about 2700 Ma ago. The gneiss terrain is characterized by migmatitic gneisses and amphibolite 3000 Ma or more old, amphibolite- and granulite-facies metamorphism, and generally moderately open folding. At three localities, 3500 Ma ages have been determined. The Minnesota River Valley, the type area for the gneiss terrain, contains a grossly interlayered sequence of migmatitic gneisses of probable igneous and volcanic derivation and an associated succession of pelitic rocks. Detailed radiometric studies indicate that the older, dominantly tonalitic gneisses and the pelitic gneiss are about 3500 Ma old; they were modified by addition of a granitic neosome about 3000 Ma ago, which followed high-grade metamorphism arid folding, and by intrusion of younger granitic rocks about 2600 Ma ago, after a second episode of tectonism. The two Archaean terrains largely evolved separately in different geographic regions, but there was some overlap of greenstone deposition onto the gneiss terrain. The virtual contemporaneity of granitic plutonism in both crustal segments possibly reflects a major subcrustal thermal surge of wide extent. Isotopic data, however, clearly distinguish late Archaean granites in the gneiss terrain—derived from older crust—from those in the? greenstone terrain, which mainly represent juvenile additions to the crust. INTRODUCTION T h e a r c h i t e c t u r e a n d e v o l u t i o n of the A r c h a e a n rocks o n the s o u t h e r n m a r g i n of the C a n a d i a n Shield, in t h e L a k e S u p e r i o r region, p r o v i d e s o m e insight i n t o early stages of d e v e l o p m e n t of the North American craton and perhaps apply to other shield a r e a s . In this r e g i o n , t h e t w o terrains t h a t c h a r a c t e r i z e A r c h a e a n crust, g r e e n s t o n e a n d gneiss ( W i n d l e y & B r i d g w a t e r , 1971; W i n d l e y , 1977), a r e p r e s e n t , a n d their age a n d spatial relationships a r e r a t h e r well k n o w n . G r e e n s t o n e granite c o m p l e x e s (2750-2600 M a ) typical of m u c h of t h e S u p e r i o r s t r u c t u r a l p r o v i n c e (Stockwell etal., 1970), occur in the n o r t h e r n p a r t of the region; a n d m i g m a t i t i c gneiss a n d a m p h i b o l i t e , in
Spec. Pubis geol. Soc. Aust., 7 (1981)
p a r t 3500 M a or older, b o r d e r this terrain o n the s o u t h . T h e t w o crustal segments a p p e a r t o constitute nearly discrete geographic a n d geologic entities t h a t evolved at widely d i f f e r e n t times, a l t h o u g h late A r c h a e a n volcanic-sedimentary rocks were locally deposited with u n c o n f o r m i t y o n the gneiss terrain a p p r o x i m a t e l y c o n t e m p o r aneously with d e v e l o p m e n t of the greenstone terrain. T h e principal p u r p o s e s of this r e p o r t a r e t o describe the stratigraphic-tectonic evolution of the t w o A r c h a e a n terrains in the L a k e Superior region a n d t o discuss their t e m p o r a l a n d spatial relationships. T h e t e r m i n o l o g y used in t h e L a k e S u p e r i o r region d i f f e r s f r o m t h a t generally used in C a n a d a
86
P. K. SIMS & Z. E. PETERMAN
between the two recognized terrains are for rocks of the Superior structural province. The boundary on Figure 2. Except for the Proterozoic greenstone terrain constitutes a part of both the shown (Keweenawan) rocks of the Midcontinent rift Shebandowan (or Wawa) volcanic-plutonic belt and the Quetico gneiss belt, as defined by Good- system, the younger supracrustal and plutonic win (1978). The older gneiss terrain has no rocks are not shown. However, areas where Prodefinite counterpart in the Superior province, terozoic rocks cover the basement are distinalthough some tonalitic gneisses and associated guished separately from those in which the baserocks in the English River subprovince (Wooden, ment rocks crop out. The boundary between the Archaean green1978), some 100 km north of Lake of the Woods, possibly are in part of the same age. The older stone and gneiss, terrains has been located moderately accurately on both sides of the Midcontigneisses in the Lake Superior region may be coeval with those in the Nutak province (Doug- nent rift system (Morey & Sims, 1976; Sims, las, 1973) of eastern Canada (Barton, 1975; 19766; Sims, 1980). Although shown on Figure 2 Bridgwater & Collerson, 1976) and those in as a narrow zone, it is marked by a wide (25-40 southwestern Greenland (Moorbath et al., 1972). km) zone of distinctive tectonism. In the disturbed zone, late Archaean rocks of the greenstone terrain were folded and metamorphosed GENERAL GEOLOGY during a late Archaean (ca 2650 Ma) deformation Rocks of both Archaean and Proterozoic age and during an early Proterozoic (Penoare exposed in the Lake Superior region (Fig. 1). kean)again, deformation. Structures formed. during Archaean rocks constitute almost the entire bed- both tectonic are oriented subrock in northern Minnesota and adjacent Ontario parallel to the disturbances boundary, and were developed within the Superior province, and crop out dis- under low to intermediate Because of continuously as inliers through younger rocks in its unique structural fabricpressures. and its great length central and southern Minnesota and in northern (over 1200 km), the zone has been the Wisconsin and Michigan, within the Southern Great Lakes tectonic zone (Sims et al,, named 1980). structural province as defined by Stockwell et al. (1970). Early Proterozoic (Proterozoic X, see Fig. 1) supracrustal rocks, some 2100-1900Ma old, Greenstone Terrain which are deformed and metamorphosed vol- Late Archaean greenstone-granite complexes canic and sedimentary rocks, compose the sur- exposed in northern Minnesota and adjacent face bedrock in much of eastern Minnesota and Ontario and, on a smaller scale, in northern northern Wisconsin and Michigan. In the Wisconsin and Michigan (Fig. 1) are grossly southern part of the region they are intruded by similar in lithology and age. The Vermilion granite-tonalite (ca 1850 Ma) plutons (Sims & district in northeastern Minnesota has been Peterman, 1980; Van Schmus, 1980). South of designated as the type area of greenstone belts in the erosional edge of these deformed rocks, sub- the greenstone terrain (Sims, 1980) of the Lake aerial rhyolite and coeval epizonal granite (1760 Superior region because it is well exposed and Ma; Van Schmus, 1978) locally crop out. Except preserved, and has been studied intensively; it is for the large anorogenic Wolf Creek Batholith in described below. central Wisconsin (Van Schmus et al., 1975), Pro- Geology. The Vermilion district—so named terozoic Y (Keweenawan) rocks are confined to a because it has been an important past producer of narrow, elongate belt that transects the older high-grade hematite iron ore—is a curvilinear rocks in and adjacent to Lake Superior. The belt of low-grade, steeply dipping, metavolcanic rocks in this belt consist mainly of thick succes- and metasedimentary rocks 10 to 30 km. wide, sions of flood basalts, cogenetic gabbroic intrusive rocks, and associated red-bed sedimentary and about 150 km long (Fig. 3). The bedded sequence is bounded on the north by the Verrocks. They are reflected in the Midcontinent gravity high, which extends from Lake Superior milion Granitic Complex, on the south by the south westward into Kansas, and are interpreted Giants Range Batholith, and on the east by the as having been deposited in a major intraconti- Saganaga Tonalite. The plutonic rocks intrude nental rift system called the Midcontinent rift the bedded sequence and are virtually coeval with it (Goldich, 1972). The western limit is estabsystem. lished arbitrarily, for the bedded rocks extend for BASEMENT ROCKS several kilometres to the west ofthe district. The eastern part of the district is truncated by the ProExtent terozoic Y (Keweenawan) Duluth Complex. The geographic extent of the Archaean base- The volcanic-sedimentary sequence is a comment rocks in the Lake Superior region and the plex pile consisting essentially of a core of mafic
EXPLANATION PHANEROZOIC Platform s e d i m e n t a r y r o c k s P R O T E R O Z O I C Y- 1 , 6 0 0 - 8 0 0 Ma Basalt, gabbro, and s e d i m e n t a r y r o c k s ( 1 , 1 0 0 Ma) A n o r o g e n i c felsic intrusions ( 1 , 5 0 0 Ma) PROTEROZOIC X - 2,500 - 1,600
Ma.
Quartzite Marquette
Rhyolite and coeval granite ( 1 , 7 7 0 Ma) Granite - tonalite plutons ( 1 , 8 5 0 Ma and 1 , 7 7 0 Ma) M e t a s e d i m e n t a r y and metavolcanic r o c k s A R C H E A N - 2 , 5 0 0 Ma and older f+
•
Granitoid r o c k s ( 2 , 7 5 0
+
:.
,| | LL LLIL
- 2 , 6 0 0
Ma)
Metavolcanic and metasedimentary r o c k s Migmatitic ( 3 , 5 0 0
gneiss and amphibolite Ma)
- 3 , 0 0 0
SOURCES: Morey, 1 9 7 8 ; Sims, Cannon, and M u d r e y , 1 9 7 8
INDEX
Fig. 1.
Generalized geologic map of Lake Superior region. Compiled by Sims.
MAP
88
P. K. SIMS & Z. E. PETERMAN
volcanic rocks and associated iron-formations that interfingers along the longitudinal axis in both directions with felsic volcanic rocks and. greywacke-slate. Detailed mapping in the western part of the district, mainly at a scale of 1:24000, has revealed a well-defined stratigraphic sequence (Sims, 1976a). It consists of an older and a younger mafic volcanic succession separated by felsic volcaniclastic and sedimentary rocks. The older mafic volcanic succession, the Ely Greenstone, is separated into a lower and an upper member by a persistent stratigraphic marker, the Soudan Iron-formation Member. In the west, the Ely Greenstone is overlain by, and inter tongues with, felsic volcaniclastic rocks of the Lake Vermilion Formation. In the central part of the district, it is overlain by similar volcaniclastic rocks assigned to the Knife Lake Group. The Newton Lake Formation overlies the Knife Lake Group in the central part of the district. It consists of a mafic member in the western part of its outcrop area and a felsic member in the east (Green, 1970). The volcanic-sedimentary accumulation differs markedly in chemical composition from the model greenstone belt of Anhaeusser (1971) and from the typical belt in the Superior province, as defined by Goodwin (1968). The lowermost exposed rocks (lower member of Ely Greenstone) consist mainly of amygdaloidal basaltic and andesitic lavas of calc-alkalic composition (Schulz, 1980) interpreted as having been deposited in relatively shallow water; they have low Y and high Ni and Cr contents. Following a relatively quiescent volcanic interval, during which iron-formation was deposited with mafic extrusive and tuffaceous rocks (Soudan Ironformation Member), contemporaneous calcalkalic (Lake Vermilion Formation and Knife Lake Group) and tholeiitic (upper member, Ely Greenstone) volcanism occurred. The tholeiitic volcanism yielded massive and pillowed, in part, variolitic and amygdaloidal lavas, whereas the calc-alkalic volcanism produced dacitic pyroclastic rocks and epiclastic rocks derived from their reworking. Schulz has interpreted the tholeiitic volcanism as having built a submarine basaltic shield volcano. The calc-alkalic volcanism took place on the flanks of the shield Volcano, with resulting coalescence of the magmatic products. Apparently, the calc-alkalic volcanism continued after tholeiitic eruptions had largely ceased. Rapid erosion of the felsic volcanic centres supplied clastic detritus, which was deposited by turbidity currents generated on the unstable slopes of the. eruptive centres. This epiclastic, felsic material buried the inactive, probably subsiding tholeiitic shield volcano. Volcan-
ism in deep water commenced again abruptly, with deposition of tholeiitic and komatiitic basalts in the western part of the Newton Lake Formation (Green & Schulz, 1977) and calcalkalic andesitic and dacitic lavas, tuffs, and breccias in the eastern part of the formation. Within the mafic succession large quantities of magma were intruded as sills, some of which differentiated to form layered intrusions. Archaean volcanic activity ceased, so far as known, after deposition of the Newton Lake Formation. The volcanic and sedimentary rocks of the sequence mainly contain greenschist-facies mineral assemblages (Morey, 1978a) developed during regional folding. The metamorphic grade increases toward the granitoid masses, and in some places a biotite isograd can be defined (Sims, 1976a). In the northwestern and northcentral parts of the district, greenschist-facies rocks of the sequences are separated abruptly from lower amphibolite-facies rocks of the Vermilion Granitic Complex by a fault system, the Haley and Burntside Lake dip-slip faults (Fig. 3). The granitoid masses that surround the greenstone belt differ in composition, structure, mode of emplacement, and relative age. The Saganaga Tonalite (Hanson, 1972), at the east end of the district, was emplaced and partly eroded during a later stage of Knife Lake deposition, so that granitic boulders were incorporated in the Knife Lake Group on the west side of the pluton (Ojakangas, 1972; McLimans, 1972). Apparently it was emplaced during a period of northwesttrending folding in the Knife Lake Group. Younger folding deformed the tonalite-bearing conglomeratic beds, which truncate the older succession having northwest trends. The Giants Range Batholith, on the south side of the greenstone belt (Fig. 3); is composite and was emplaced during two major magmatic episodes, after volcanism ceased. The western part, consisting of several distinct plutons ranging in composition from granite to tonalite (Sims & Viswanathan, 1972) was emplaced contemporaneously with regional folding. These rocks have a steep foliation conformable to the internal structure of the adjacent bedded rocks. The eastern part of the batholith, a large body of coeval monzonite and quartz monzonite that truncates the regional structure and cuts out the lower part of the volcanic pile, has a steep primary foliation and lineation adjacent to its margins that are discordant to the regional structure. The Vermilion Granitic Complex, on the north side of the district, differs from the other large granitoid masses in being dominantly a migmatitic body of major dimensions. It contains large quantities of biotite schist (metagreywacke) and amphibolite
96°
t :
\ 48°-
46°
+ +
+ +
+
—
+
+, + + — + + v + +
+
+
^
+ +
+
+ + + + + + + + + + + + + +• + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + -1- + + + + + + + + + +
+
+
+ -»-
+
+ +
+
+
+ +
+
+
+ +
+
+
+ +
+
+
+ +
+
+
+ +
+
+
+ +
+ +
+
+
+ +
+ + +
+
+ J i ri + \ ni r t ^ +
+
^
ASA
EXPLANATION
Lava and gabbro of Midcontinent rift system
>
GREENSTONE TERRANE
X
Q
Greenstone-granite complexes of late Archean age
E m
Greenstone-granite complexes concealed by Proterozoic rocks GNEISS TERRANE Migmatitic gneiss and amphibolite of early Archean age Gneiss concealed by Proterozoic rocks Boundary between Archean terranes
44°
Fault, rejuvinated during Proterozoic (Keweenawan) time
Fig. 2. Map showing known and inferred distribution of Archaean basement terrains in Lake Superior region. Bayfield Penninsula is underlain by Proterozoic Y sedimentary rocks. Numbers refer to localities in text.
O
>
m > z
GO O
c H
X W 7*
Z O >
>a >z cN
r5
a
90
92°30
92°00'
i •
91 °00
91°30
r
7-
CANADA
Syenitic
plutons,
GIANTS RANGE BATHOLITH Monzonite and quartz monzonite Granite-tonalite Vermilion Granitic Complex Saganaga tonalite 47°45
Newton Lake Formation Knife Lake Group Lake Vermilion Formation
Contact High-angle strike-slip fault
Fig. 3.
ELY GREENSTONE Upper member
High-angle dip-slip fault
Soudan iron-formation member
Structural form lines
Lower member
G e n e r a l i z e d g e o l o g i c m a p o f t h e V e r m i l i o n d i s t r i c t , M i n n e s o t a . C o m p i l e d f r o m S i m s (1973) a n d G r u n e r (1941).
P. K. SIMS & Z. E. PETERMAN
EXPLANATION
48°00
ARCHAEAN IN SOUTHERN CANADIAN SHIELD
(metabasalt) of lower amphibolite grade, and grades from the schist through migmatite to massive leucogranite (Southwick & Sims, 1979). The migmatite grades northward (outside the map area, Fig. 3) into a thick succession of biotite schist of mainly greywacke affinity; the schist and the Vermilion Granitic Complex together compose the Quetico gneiss belt as defined in Canada. In addition to the large bodies of plutonic rocks, small, post-tectonic bodies of general syenitic composition intrude the greenstone locally (Sims & Mudrey, 1972) as shown on Figure 3. Structurally, the Archaean bedded rocks compose a complex anticline having a predominantly eastward trend (Fig. 3). Fold axes trend northwestward to northeastward, and reflect at least two generations of folding (Hooper & Ojakangas, 1971; Sims, 1972a, 1976a). The steep dips and general easterly trend resulted from a regional deformation (Fj) along the length of the belt attributed to dominant north-south compression. A local, slightly younger deformation (F j superposed on the steeply dipping rocks in the area south of Lake Vermilion, expressed mainly by small-scale Z-folds, is attributed to shortening resulting from a northward-convex bending of Fj fold axes (see Fig. 3). .Faults that are younger than the folds are mainly of two types—longitudinal and transverse strike-slip faults. The major longitudinal structure, the Vermilion fault system, has an apparent maximum right-lateral displacement of about 17 km (Sims, 1976a); it transects the central part of the district and together with other rightlateral faults accounts for part of the pronounced east-northeast elongation of the greenstone belt. A major dip-slip fault, segments of which have been named the Haley and Burntside Lake Faults, separates much of the greenstone belt from the Vermilion Granitic Complex. It is estimated from contrasting mineral assemblages on either side to have a vertical component of movement of about 2000 m. Tectonic-igneous history. The greenstonegranite complexes represent a major rockforming event during which thousands of cubic kilometres of new crust formed. Volcanism, sedimentation, folding, metamorphism,- and igneous intrusion took place during a time span of 50 to 100 Ma, about 2700 Ma ago (see Goldich, 1972, for discussion). Folding and metamorphism of the greenstone were contemporaneous with emplacement of the granitic magmas that now constitute the Vermilion Granitic Complex and the western part of the Giants Range Batholith, and are interpreted as resulting from diapiric emplacement of these 2
91 granitic bodies. The mechanics of folding of this type have been discussed by Schwerdtner et al. (1979). Folding was accompanied by greenschist regional metamorphism and, adjoining the plutons, by metamorphism of low-amphibolite grade. This crustal segment stabilized by the end of the Archaean and has remained stable; it now constitutes the southern part of the stable inner part of the North American craton. The longterm stability of the terrain is indicated by widespread epizonal igneous rocks and by the local preservation of miarolitic cavities in these intrusions, which must have formed at depths no greater than 2km. It also is indicated by Rb-Sr and K-Ar biotite ages, commonly between 2400 and 2600 Ma in the Vermilion district (summary by Peterman, 1979&), which indicate that the region has not been subjected to any major thermal event since latest Archaean or earliest Proterozoic time. Apparently the terrain has remained at or near existing sea level since early Proterozoic time, as indicated by the approximately concordant altitudes in the Mesabi Range area in northern Minnesota of the bases of the early Proterozoic Pokegama Quartzite (Animikie Group), the Cretaceous platform sedimentary rocks, and Pleistocene glaciation.
Some implications to evolution of the Superior
province. Some results of the studies in the Vermilion district are pertinent to understanding the evolution of the Archaean Superior province. 1. Physical evidence of the crust underlying the accumulation of volcanic and sedimentary rocks is lacking because the lower part of the volcanic pile has been cut out by post-tectonic monzonite of the Giants Range Batholith (Fig. 3). Also, conglomerates containing plutonic rock clasts increase in abundance upward in the sequence, and in fact are mainly present in beds of the Knife Lake Group, where it has been clearly shown (Gruner, 1941; Ojakangas, 1972) that the clasts were derived from the Saganaga Tonalite. Conglomeratic lenses in the Ely Greenstone, the Knife Lake Group, and the Newton Lake Formation contain small granitic clasts having Rb-Sr systems consistent with an age of 2630m.y., which is indistinguishable from ages of the volcanic units (Jahn & Murthy, 1975). The low initial Sr-isotope ratio (0.7008) of these rocks suggests probable derivation from contemporaneous plutonic rocks. Geochemical and isotopic data place important constraints on the environment in which the greenstone-granite complexes formed and the sources from which the constituent rocks were derived. First, low initial S r / S r ratios of both 87
86
P. K. SIMS & Z. E. PETERMAN
92
-
(Rb/Sr) Q = 0.03 ~ ( R b / S r ) o = 0.02 O 1
1
2.8
1
2.7
20@6 O O 7 0 4 G^ 5 I
-
8
I
I 2.6
2.5
AGE (1 0°YRS)
Fig. 4. Initial S r / S r ratios for some units in the Vermilion district plotted against Rb-Sr isochron ages. Data are (1) Northern Light Gneiss; (2) Saganaga Tonalite; (3) Icarus pluton (Hanson et al., 1971); (4) Vermilion Granitic Complex (Peterman et al1972; Jahn & Murthy, 1975); (5) Giants Range Granite (Prince & Hanson, 1972); (6) granitic clasts from conglomerate lenses in the Ely Greenstone, Knife Lake Group, and Newton Lake Formation; (7) Ely Greenstone, and (8) Newton Lake Formation (Jahn & Murthy, 1975). Reported uncertainties (2 sigma) of these values range from 3 x 10~ to 19 x 10 ~ . The average value weighted inversely by the square of the uncertainties is 0.70065. The solid lines depict closed-system mantle evolution from an initial ratio of 0.6990 at 4.55 Ga with starting Rb/Sr ratios of 0.03 and 0.02. 87
86
4
4
plutonic and volcanic rocks have led workers to conclude that these were not derived from older continental crust (Hanson et al., 1971; Hanson & Goldich, 1972; Jahn & Murthy, 1975). The plutonic and volcanic rocks have initial Sr-isotope ratios in the range of 0.7002 to 0.7009 as shown on Figure 4. A mean value of 0.70065 would characterize the rocks of this terrain. For comparison, two closed-system mantle growth curves are shown for initial Rb/Sr ratios of 0.02 and 0.03, respectively evolving from a primordial value of 87Sr/86Sr of 0.6990. The present-day intercepts of these curves, 0.7029 and 0.7048, would encompass many Sr-isotope values for mantle-derived oceanic basalts and calc-alkalic island-arc rocks. As emphasized by the investigators cited in Figure 4, these low initial Srisotope ratios are consistent with a derivation from a mantle source either directly or indirectly from short-lived intermediate stage materials. Consequently, derivation of the late Archaean granitic rocks from older (ca 3500 Ma) gneisses such as those exposed in the Minnesota River Valley or in northern Michigan is definitely precluded. Granitic rocks derived from such older gneisses in the late Archaean would have had significantly higher initial Sr-isotope ratios.
Pb-isotope data also support derivation of the late Archaean granitoid rocks in the Vermilion district from mantle or short-lived intermediatestage rocks. Recent high-precision Pb-isotope analyses by Maryse H. Delevaux of the U.S. Geological Survey on microcline from two widely separated cores of quartz monzonite in the Vermilion Granitic Complex are virtually identical with 206pb/207pb model ages of 2690 Ma (B. R. Doe, written commun., 1979). This age agrees with the Rb-Sr isochron age of 2640 ± 50 Ma (Peterman et al., 1972; Jahn & Murthy, 1975). Doe & Delevaux (1980) have emphasized that the Pb isotopic composition of these late Archaean granites differs markedly from that of the Sacred Heart Granite (Goldich et a/., 1970) of similar age in the gneiss terrain of the Minnesota River Valley. They reported an age of 2605 Ma derived from Pb-isotope systematics for the Sacred Heart Granite, which agrees with independent dating. The data, however, indicate that this rock was derived from crustal material that had undergone a complex pre-2800 Ma history, including granulite-facies metamorphism; this contrasts with the late Archaean granites of northern Minnesota which have isotopic data indicating a juvenile source at 2700Ma. 2. Geological and geochemical data on the late Archaean granitoid rocks from northern Minnesota indicate the complexity of their genesis, even though they are virtually coeval. Arth & Hanson (1975) have concluded that the syntectonic granite (quartz monzonite) that characterizes the Vermilion Granitic Complex and composes much of the western, older part of the Giants Range Batholith was probably derived by partial melting of short-lived greywacke at crustal depths. These bodies are linear, subparallel to the long dimensions of the Vermilion greenstone-belt, and abut metagreywacke. This greywacke possibly was down-buckled into the crust to depths at which partial melting took place; alternatively, an older cycle of greywacke, for which we have no evidence now at the surface, could have been the source of these magmas. At any rate, the availability of large quantities of greywacke in the Archaean crust may have been required for the generation of large granitic (quartz monzonite) magma bodies. Gneiss Terrain The gneiss terrain, as defined in the Lake Superior region, is that segment of Archaean crust composed mainly of basement gneisses that are older than 3000 Ma, have an ubiquitous amphibolite facies and, locally, granulite-facies metamorphic grade, and a distinctive structural style characterized in many areas by moderately
93 on banded gneiss and migmatite indicate a minimum age of about 2800 Ma (Van Schmus & Anderson, 1977). The Rb-Sr data indicate elevated initial Sr/ Sr ratios, however, suggesting that the migmatite and gneiss were derived from crustal rocks that were formed more than 3000 Ma ago. Similar migmatitic gneisses in the Lake Arbutus area (loc. 4, Fig. 2), about 70 km west of locality 3, have a U-Pb zircon age of 2900 Ma and a Rb-Sr isochron age of about 2800Ma (Dubois & Van Schmus, 1978). These dates could be primary ages or represent the time of upper amphibolite- to granulite-facies metamorphism. In east-central Minnesota, gneisses of definite Archaean age (Morey, 1978Z?) occur near Mille Lacs Lake in the cores of gneiss domes. East of the lake (loc. 5, Fig. 2), a migmatitic granitic gneiss (McGrath Gneiss) has a minimum Rb-Sr whole-rock isochron age of 2700 Ma (Stuckless & Goldich, 1972). West of the lake (loc. 6), gneisses that resemble those in the MRV but have not been dated radiometrically occur over a wide area (Morey, 1978&). In northern Michigan, south of Marquette (Fig. 2), granitic gneisses are exposed in the cores of domes and uplifted fault-bounded blocks at several localities, but their isotopic systems are highly disturbed and interpretation of existing data is equivocal. The southern complex of the Marquette district (loc. 7, Fig. 2) contains granitic gneisses of tonalite to granite composition (Cannon & Simmons, 1973) for which Rb-Sr wholerock analyses plot between 2500 Ma and 2800 Ma reference isochrons, and one sample yielded a model age of about 3200 Ma (Van Schmus & Woolsey, 1975). Van Schmus & Woolsey concluded, however, that the apparent old age resulted from Sr and Rb redistribution during an event about 2100 to 2200 Ma ago, and that 2800Ma is approximately the age of these rocks. Similar gneisses form the cores of three domes just west of the southern complex.
ARCHAEAN IN SOUTHERN CANADIAN SHIELD
dipping foliation (25-60°) and low to moderately plunging fold axes. The terrain contains some apparently younger (2800-3000 Ma) migmatitic gneisses of Archaean age (Van Schmus & Anderson, 1977), but their extent, abundance, and stratigraphic relationship to the older gneisses have not been determined fully because of poor outcrops and highly disturbed isotopic systems. In addition, late Archaean supracrustal rocks, which are intruded by late Archaean granite, locally unconformably overlie the gneisses, as at Watersmeet, Michigan (Sims, 1980) and in Dickinson County, northern Michigan (loc. 8, Fig. 2; James et al. 1961). In contrast to the greenstone terrain, the gneiss terrain records a long history of rock-forming events and tectonism that continued into early Proterozoic time at least until 1800Ma (Goldich & Wooden, 1978; Sims, 19766). }
Geographic extent and radiometric ages. Exposures of Archaean rocks in the gneiss terrain are limited by an extensive cover of early Proterozoic and younger rocks (Fig. 1); accordingly, knowledge of the extent of this crustal segment is based on sporadic exposures and extrapolation in the subsurface. The gneiss terrain is inferred to continue eastward into the Lake Huron region of Ontario (Sims et at., in press) and westward an unknown distance beneath Phanerozoic cover. It extends southward for at least 200 km from its juncture with the greenstone terrain (see Fig. 2), but whether it composes the basement in the central United States is equivocal. In the Lake Superior region, gneisses that are definitely older than 3000 Ma have been delineated in the Minnesota River Valley (MRV) in southwestern Minnesota (locality 1, Fig. 2) and near Watersmeet, Michigan (loc. 2). In the MRV, an inlier of Archaean rocks within Cretaceous strata, the oldest dated rocks are tonalitic and granodioritic gneisses; these rocks have U-Pb zircon and Rb-Sr isochron ages of about 3500Ma and possibly are as old as 3 800 Ma (Goldich & Wooden, 1978). A comparable age has been determined for tonalitic gneiss in the core of a gneiss dome (Sims & Peter man, 1976) in the Watersmeet area (Peterman et al, 1980). This gneiss has a minimum age of 3410 Ma, as indicated by 207pb/ 206pb data on zircon; a primary age of 3500 Ma or greater is suggested by a Concordia plot of the zircon data. McCulloch & Wasserburg (1980) obtained a Sm-Nd model age of 3600 Ma (3620 ± 30Ma and 3570 ± 50Ma) on a whole-rock sample of the same gneiss. Younger Archaean ages have been obtained on gneisses from other areas. In central Wisconsin (loc. 3 Fig. 2), Rb-Sr and U-Pb zircon analyses y
87
86
Composition and structure. The Minnesota River Valley provides the only coherent succession of well-dated Archaean gneisses exposed in the Lake Superior region, and it has been designated the type area for this terrain (Sims, 1980). The area has been studied intensively since the late 1950s by both geologists and geochemists. The studies have been concentrated in two rather well-exposed areas, the Morton area (loc. 1A, Fig, 2) and the Granite Falls area (loc. IB). The rocks in the MRV consist of a grossly conformable sequence of interlayered migmatitic gneisses, apparently a few thousand metres thick (Grant, 1972), that is intruded by younger granitic and pegmatitic rocks. The gneisses are
94
P. K. SIMS & Z. E. PETERMAN
folded on east-trending, gently plunging axes, and have mineral assemblages characteristic of upper amphibolite- and granulite-facies metamorphism (Himmelberg & Phinney, 1967; Grant, 1972). In the Morton area, four gross stratigraphic units have been delineated (Grant, 1972). The three older units recognized by Grant are quartzofeldspathic gneisses which, respectively, have abundant, common, and rare rafts of amphibolite. The middle unit in this succession has been formally named the Morton Gneiss (Goldich et al., 1970) where exposed in quarries near Morton. Where not migmatized, the quartzofeldspathic gneisses are tonalitic or granodioritic in composition and generally are compositionally layered. Where migmatized, the neosome is granite and, locally, pegmatite. A probable youngest stratigraphic unit is composed of biotitic gneisses and amphibolite. The biotite gneisses are of two types, one containing the mineral association biotite-cordierite-garnetanthophyllite in addition to quartz and plagioclase, and the other containing sillimanite and Kfeldspar and, locally, garnet and cordierite. The associated amphibolite occurs as discontinuous layers, lenses and boudins in the biotite gneisses. The younger Sacred Heart Granite is a medium-grained, generally homogeneous to weakly foliated rock that partly crosscuts older structure in the gneisses, indicating that it is latetectonic or possibly post-tectonic. The rocks in the Granite Falls area also compose a layered succession of mafic and felsic gneisses (Himmelberg, 1968), but detailed correlation with the succession in the Morton area is uncertain. The principal exposed gneiss, earlier called the Montevideo Gneiss (Goldich et al., 1970) is a layered rock commonly consisting of two distinct lithotypes—grey, foliated granodiorite and red massive granite; the latter forms both concordant and discordant layers and veins in the granodiorite. This composite unit is interlayered with hornblende-pyroxene gneiss, a dark, medium-grained, equigranular rock varying from uniform amphibolite to a layered gneiss. A third unit in the succession is composed of garnetbiotite gneiss, a dark grey rock containing both biotite and orthopyroxene. Post-metamorphic mafic dykes and a small body of granite about 1800 Ma old (Goldich et al., 1970) intrude the gneisses. Combined field and radiometric studies of the migmatitic gneisses in both the Morton and Granite Falls areas indicate three rock-forming episodes and two periods of regional metamorphism. Detailed studies of the Morton Gneiss, summarized by Goldich & Wooden (1978), indicate
that the tonalitic and granodioritic gneiss and associated amphibolite (paleosome) are 3500 Ma or more old, and that the granitic neosome was formed during two later episodes. An older, deformed granite and pegmatite were emplaced about 3000 Ma ago and a younger, largely undeformed granite neosome was emplaced about 2600Ma ago. The latter is comparable in age to the Sacred Heart Granite, but for the most part probably is not cogenetic with it. A similar, but somewhat simpler, succession of rock-forming events has been determined for the Granite Falls area (Goldich & Wooden, 1978). The grey granodioritic phase of the so-called Montevideo Gneiss is 3500 Ma or more old and the red granite neosome is about 3000 Ma. Presumably the hornblende-pyroxene gneiss and garnet-biotite gneiss that are interlayered with the Montevideo Gneiss also are mainly about 3500 Ma old. The protoliths of the gneisses are poorly known for the most part because of the high-grade metamorphism. The possible oldest rocks in the sequence, the quartzofeldspathic gneisses, probably are metamorphosed plutonic and volcanic rocks, as favoured by Goldich & Wooden (1978). The intercalated amphibolites and hornblendepyroxene gneiss probably are metamorphosed volcanic rocks and sills; the amphibolite in the Morton Gneiss is tholeiitic to basaltic komatiite in composition. The possible youngest unit of pelitic rocks in the Morton area clearly is metasedimentary in origin; these rocks have a primary layering suggestive of bedding, and a chemical composition indicative of clay-rich sediments such as argillite (Goldich & Wooden, 1978). The garnet-biotite gneiss in the Granite Falls area probably also is metasedimentary, perhaps a metagreywacke (Grant, 1972). A metagabbro in one of the bodies of hornblende-pyroxene gneiss at Granite Falls has a composition of anorthositic gabbro (Goldich & Wooden, 1978). Tectonic evolution. The geologic and radiometric studies in the MRV indicate that these rocks had a protracted Archaean history of nearly a billion years, and that the migmatitic gneisses were formed over an interval of at least 500 Ma. The scattered data from other areas of exposed gneisses are consistent with this time span, but in themselves are fragmentary. In the MRV, the older (ca 3500 Ma) gneisses were folded and metamorphosed before emplacement of the 3000 Ma old pegmatite and granite, and presumably attained amphibolite or granulite metamorphic grade at that time. A younger tectonic episode occurred before 2600 Ma, before emplacement of the Sacred Heart Granite. A later
ARCHAEAN IN SOUTHERN CANADIAN SHIELD
thermal metamorphism approximately contemporaneous with emplacement of 1800 Ma plutons in the area, produced retrogressive effects on the rocks and reset K-Ar and Rb-Sr ages on biotite in the Granite Falls area (Goldich et al., 1970). These authors proposed that the biotite ages reflected the effects of differential vertical uplift and stabilization of two major blocks. The Morton block stabilized in the latest Archaean (ca 2500 Ma), whereas the Granite Falls block was uplifted and cooled at about 1800Ma. A major, east-trending fault, shown in Figure 1, separates the two blocks. The style of deformation of the gneisses, as well as the grade of metamorphism, differs markedly from that in the greenstone terrain. Throughout the MRV, the gneisses are deformed into east-northeast-trending, open, buckle folds that plunge gently eastward (Himmelberg, 1968; Grant, 1972). Except for local, small recumbent folds, axial surfaces dip steeply to the south or north; lineations are variably developed in separate rock units and are subparallel to major and minor fold axes. Two episodes of younger folding on a small scale have been distinguished (Bauer, 1980) in the Granite Falls area (loc. IB, Fig. 2). The dominant folds, which control the distribution of rock units, resulted from subhorizontal compression oriented about northsouth. The history of the apparently younger Archaean gneisses in central Wisconsin and Michigan is too uncertain now for meaningful discussion. The distribution and significance of late Archaean supracrustal successions and granitoid rocks in the gneiss terrain is gradually becoming known through combined geologic mapping and radiometric studies. In addition to that in the Watersmeet, Michigan area, noted above, a supracrustal succession that has an exposed areal extent of 400 km 2 has been delineated in Dickinson County, northern Michigan (James et al., 1961; loc. 8, Fig. 2). It consists of a thick basal unit of metaconglomerate and meta-arkose overlain by amphibolite and mica schist, and is intruded by granitic gneiss and granite. It lies unconformably on a granitic gneiss informally called gneiss at Norway Lake. The intrusive granite is about 2700 Ma old (W. R. Van Schmus, pers. comm., 1979). Relationships of the Two Archaean Terrains We contend that the two Archaean crustal segments of widely different ages in the Lake Superior region largely evolved separately in adjacent parts of the crust, and that they do not
95
represent two exposed crustal levels resulting from the vagaries of erosion. As was noted previously, late Archaean volcanic-sedimentary rocks were deposited unconformably at least locally on the older gneiss crustal segment, but by comparison with the vast accumulations in the greenstone terrain, as defined in the region, these are relatively insignificant volumetrically. Several indirect lines of evidence support our contention that early Archaean gneisses do not underlie the greenstone terrain. As discussed in an early part of this report, there is no field evidence for pre-2750Ma sialic material in the greenstone-granite terrain of the Lake Superior region, although search for possibly older rocks has been extensive (Sims, \912b). Also, initial 87 Sr/ 86Sr ratios and Pb-isotope data for granitic rocks in the greenstone terrain negate the possible occurrence of basement such as that exposed in the Minnesota River Valley. If a sialic basement existed, it must have been composed solely of tonalites with mantle-like Rb/Sr ratios. Although such rocks occur (Peterman, 1979a), their host terrains contain abundant granitic material with much higher Rb/Sr ratios. Another constraint on the nature of the crust beneath the greenstone belts is the strongly contrasting tectonic styles characteristic of the two terrains. If a continuous basement of MRV-type gneisses had underlain the greenstone belts, the structure of the overlying rocks should have been fundamentally similar to that in the gneiss domains. Instead, the deformational style in the greenstone belts is dominated by vertical tectonism, with folding caused mainly by the diapiric rise of granitoid magma bodies, in a manner similar to the gravitational experiments by Ramberg (1967). The tectonic environment that existed, together with the vast geographic scale of the greenstone belts, indicates that the crust at this time was weak and probably thin. By contrast, the small accumulations of late Archaean supracrustal rocks that demonstrably overlie a gneiss basement, as in Dickinson County, Michigan (loc. 8, Fig. 2), are folded on east-trending, nearly horizontal axes that are subparallel to the structure in the underlying gneiss basement. The fold pattern indicates dominant subhorizontal compression, a style characteristic of earlier Archaean tectonic episodes in the gneiss terrain. Finally, the nature of the boundary supports a dichotomy of the two Archaean crustal segments. In the Watersmeet, Michigan area (loc. 2, Fig. 2) granite dykes (2600 Ma old) that apparently are offshoots from a large magma body in the greenstone terrain that crystallized as the late Archaean Puritan batholith (Sims, 1980), were emplaced across the boundary, intruding the 3500 Ma tonalite gneiss in the
P. K. SIMS & Z. E.
96
gneiss terrain. The granite in the dykes has textures indicative of crystallization at shallow or moderate depths, suggesting that the crustal levels on both sides of the boundary at the time of intrusion were comparable. In the same way, the 2600 Ma Sacred Heart Granite in the MRV has textures indicative of emplacement at a relatively high level in the crust (Goldich & Wooden, 1978); this indicates that at 2600 Ma the gneiss terrain was not deeply buried; it could not then have been blanketed by a thick cover of supracrustal rocks. The specific tectonic environment in which the late Archaean greenstone-granite complexes evolved is uncertain, but the available data are compatible with a rifting mechanism and accumulation on a simatic crust. Segmentation of the early Archaean gneissic crust by rifting, with upwelling of mantle-derived material, is a plausible model that deserves further testing.
PETERMAN
ACKNOWLEDGMENTS The data summarized here are based on geologic mapping and laboratory studies carried out during the past two decades, mainly by personnel of the Minnesota Geological Survey and the U.S. Geological Survey, who are too numerous to mention individually. The key element in this work has been refinement of radiometric dating techniques, which has made it possible to obtain reliable ages of geologic events. S. S. Goldich and colleagues pioneered these studies in the region, and more recently, significant studies also have been made by G. N. Hanson and W. R. Van Schmus. We are indebted to G. B. Morey, J. L. Wooden, and A. Y. Glikson for critical reviews of the manuscript. Ages have been recalculated, where necessary, using isotopic and decay constants recommended by the IUGS Subcommission on geochronology (Steiger & Jaeger, 1977).
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BARTON, J. M. JR., 1975: Rb-Sr isotopic characteristics and chemistry of the 3.6b.y. Hebron Gneiss, Labrador. Earth planet. Sci. Lett., 27, 4 2 7 - 4 3 5 . BAUER, R. L., 1980: Multiphase deformation in the Granite Falls-Montevideo area, Minnesota River
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Michigan. U.S. geol. Surv. J. Res., 1, 165-172. DOE, B. R., & DELEVAUX, M. H., 1980: Lead isotope investigations in the Minnesota River Valley— Late-tectonic and post-tectonic granites. Spec. Pap. geol. Soc. Am., 182 (Goldich Vol.), 1 0 5 - 1 1 2 . DOUGLAS, R. J. W., 1973: Geological provinces; in The
National Atlas of Canada, 27-28. Can. Geol. Surv., Ottawa. DUBOIS, J . F . , & VAN SCHMUS, W . R . , 1 9 7 8 : P e t r o l o g y
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Proc. 24th Inst. Lake Superior Geol., Milwaukee, Wisconsin, 11. GOLDICH, S. S., 1972: Geochronology in Minnesota; in Sims, P: K., & Morey, G. B. (Eds) Geology of
Minnesota—A Centennial Volume, 27-37. Geol. Surv. Minn,, Minn.
geochem. Conf, 38-106. Univ. Toronto Press, Toronto. GRANT, J. A., 1972: Minnesota River Valley, southwestern Minnesota; in Sims, P . K., & Morey, G. B.
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Volume, 177-196. Geol. Surv. Minn., Minn. GREEN, J. C., 1970: Lower Precambrian rocks of the G a b b r o Lake quadrangle, northeastern Minnesota.
Spec. Pubis geol. Surv. Minn., SP-13. GREEN, J. C., & SCHULZ, K. J., 1977: Iron-rich komatiites in the early Precambrian Vermilion district, Minnesota. Can. J. Earth Sci., 14, 2181-2192. GRUNER, J. W., 1941: Structural geology of the Knife Lake area of northeastern Minnesota, Minnesota.
Bull. geol. Soc. Am., 52, 1577-1642. HANSON, G. N., 1972: Saganaga batholith; in Sims, P.
K., & Morey, G. B. (Eds) Geology of Minnesota—A Centennial Volume, 102-107. Geol. Surv. Minn., Minn. HANSON, G . N . , & GOLDICH, S . S . , 1972: E a r l y Pre-
cambrian rocks in the Saganaga Lake—Northern Light Lake area, Minnesota—Ontario: Part II. Petrogenesis. Mem. geol. Soc. Am., 135, 179-192.
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HANSON, G . N . , GOLDICH, S. S., ARTH, J. G . , & YARDLEY, D . H . , 1971: Age of the early Pre-
PETERMAN, Z. E., 19796: Geochronology and the Ar-
Sci., 8, .1110-1124. HIMMELBERG, G. R . , 1968: Geology of Precambrian
PETERMAN, Z . E . , GOLDICH, S. S., HEDGE, C . E . , & YARDLEY, D . H . , 1972: Geochronology of the
cambrian rocks of the Saganaga Lake—Northern Light Lake area, Ontario-Minnesota. Can. J. Earth
rocks, Granite Falls-Montevideo area, southwestern Minnesota. Spec. Pubis geol. Surv. Minn., SP-5. HIMMELBERG, G. R . , & PHINNEY, W. C., 1967: Granulite-facies metamorphism, Granite Falls-Montevideo area, Minnesota. J. Petrol., 8, 325-348. HOOPER, P . R . , & OJAKANGAS, R . W . , 1971: Multiple deformation in Archaean rocks of the Vermilion district, northeastern Minnesota. Can. J.. Earth Sci., 8, 423-434. JAHN, BOR-MING, & MURTHY, V. RAMA, 1975: Rb-Sr
ages of the Archaean rocks from the Vermilion district, northeastern Minnesota. Geochim. cosmochim. Acta, 39, 1679-1689.
JAMES, H . L . , CLARK, L. D . , LAMEY, C . A . , & PETTIJOHN, F. J . , 1961: Geology of central Dickinson
County, Michigan. Prof. Pap. U.S. geol. Surv., 310.
MCCULLOCH, M . T . , & WASSERBURG, G . J . , 1980: Sm-
Nd model ages from an early Archaean tonalitic gneiss, northern Michigan. Spec. Pap. geol. Soc. Am., 182 (Goldich Vol.), 135-138. MCLIMANS, R . K., 1972: Granite-bearing conglomerates in the Knife Lake Group, Vermilion district; in Sims, P. K., & Morey, G. B. (Eds) Geology of Minnesota—A Centennial Volume, 91-97. Geol. Surv. M inn., Minn.
MOORBATH, S., O'NIONS, R . K., PANKHURST, R . J., & GALE, N . H . , 1972: Further rubidium-strontium
age determinations on the very early Precambrian rocks of the Godthaab district, west Greenland. Nature, Phys. Sci., 240, 78-82. MOREY, G. B., 1978a: Metamorphism in the Lake Superior region, U.S.A., and its relation to crustal evolution; in Fraser, J. A., & Heywood, W. W. (Eds) Metamorphism in the Canadian Shield. Pap. Can. geol, Surv., 78-10, 283-314. , 1978hi Lower and middle Precambrian stratigraphic nomenclature for east-central Minnesota. Rep. Invest, geol. Surv. Minn., 21. MOREY, G . B., & SIMS, P. K., 1976: Boundary between - two Precambrian W terranes in Minnesota and its geologic significance. Bull. geol. Soc. Am., 87,
chaean of the United States. Econ. Geol., 74, 1544-1562. Rainy Lake region, Minnesota-Ontario. Mem.
geol. Soc. Am., 135, 193-215. PETERMAN, Z . E . , ZARTMAN, R. E . , & SIMS, P . K., 1980:
Tonalitic gneiss of early Archaean age from northern Michigan, U.S.A. Spec. Pap. geol. Soc. Am., 752 (Goldich Vol.), 125-134. PRINCE, L. A., & HANSON, G. N., 1972: Rb-Sr isochron ages for the Giants Range granite, northeastern Minnesota. Mem. geol. Soc. Am., 135, 217-224. RAMBERG, HANS, 1967: Gravity, Deformation and the Earth's Crust. Academic Press, London. SCHULZ, K. J,, 1980: The magmatic evolution of the Vermilion greenstone belt, northeastern Minnesota. Precamb. Res., 11, 215-245.
SCHWERDTNER, W . M . , STONE, D . , OSADETZ, K., MORGAN, J., & STOTT, G. M . , 1979: Granitoid
complexes and the Archaean tectonic record in the southern part of northwestern Ontario. Can. J. Earth Sci., 16, 1965-1977. SIMS, P. K., 1972a: Vermilion district and adjacent areas; in Sims, P. K., & Morey, G. B. (Eds) Geology of Minnesota—A Centennial Volume, 49-62. Geol. Surv. Minn., Minn. _, 19726: Northern Minnesota general geologic features; in Sims, P. K., & Morey G. B. (Eds)Geology of Minnesota—A Centennial Volume, 41-48. Geol. Surv. Minn., Minn. ______ 1973: Geologic map of western part of Vermilion district, northeastern Minnesota. Geol. Surv. Minn., misc. Map, M13. , 1976a: Early Precambrian tectonic-igneous evolution in the Vermilion district, northeastern Minnesota. Bull. geol. Soc. Am., 87, 379-389. , 19766: Precambrian tectonics and mineral deposits, Lake Superior region. Econ. Geol., 71, 1092-1127. , 1980: Boundary between Archaean greenstone and gneiss terranes in northern Wisconsin and Michigan. Spec. Pap. geol. Soc. Am., 182 (Goldich Vol.), 113-124. SIMS, P . K., CANNON, W . F., & MUDREY, M . G . , JR.,
1978: Preliminary geologic map of Precambrian rocks in part of northern Wisconsin. U.S. Geol. Surv. Open-file Rep. 78-318 (Denver, Colorado). 141-152. SIMS, P. K., CARD, K. D., & LUMBERS, S. B., in press: Evolution of the early Proterozoic basins of the OJAKANGAS, R. W . , 1972: Graywackes and related rocks Great Lakes region. Spec. Pap. geol. Surv. Can. of Knife Lake Group and Lake Vermilion Formation, Vermilion district;.in Sims, P. K., & Morey, SIMS, P . K., CARD, K. D . , MOREY, G . B., & PETERMAN, Z. E., 1980: The Great Lakes tectonic zone—a maG. B. (Eds) Geology of Minnesota-A Centennial jor Precambrian crustal structure in central North Volume, 82-90. Geol. Surv. Minn., Minn. America. Bull. geol. Soc. Am., 91, 690-698. PETERMAN, Z. E., 1979a: Strontium isotope geochemistry of late Archaean to late Cretaceous tona- SIMS, P . K., & MUDREY, M . G., JR., 1972: Syenitic plutons and associated lamprophyres; in Sims, lites and trondhjemites; in Barker, F. (Ed) P. K., & Morey, G. B. (Eds) Geology of MinneDevelopments in Petrology, 6, Trondhjemites, sota—A Centennial Volume, 140-152. Geol. Surv. Dacites, and Related Rocks, 133-147. Elsevier Sci. Minn., Minn. Publ. Co., Amsterdam.
P. K. SIMS & Z. E. PETERMAN VAN SCHMUS, W. R., 1978: Geochronology of the southern Wisconsin rhyolites and granites. Geosci. Sr ages of reactivated Precambrian gneisses and Wise., 2, 19-24. granite in the Marenisco-Watersmeet area, nor,1980: Chronology of igneous rocks associated thern Michigan. U.S. geol. Surv. J. Res., 4, (4), with Penokean orogeny in Wisconsin. Spec. Pap. 405-414. geol. Soc. Am., 182 (Goldich Vol.), 159-168. , 1980: Geology and Rb-Sr age of lower Proterozoic granitic rocks, northern Wisconsin. Spec. Pap. VAN SCHMUS, W. R., & ANDERSON, J. L., 1977: Gneiss geol. Soc. Am., 182 (Goldich Vol.), 139-146. and migmatite of Archaean age in the Precambrian SIMS, P. K., & VISWANATHAN, S., 1972: Giants Range basement of central Wisconsin. Geology, 5, 45-48. batholith; in Sims, P. K., & Morey, G. B. (Eds) Geology of Minnesota—A Centennial Volume, VAN SCHMUS, W . R . , MEDARIS, L . G . , & BANKS, P. 0 . , 120-139. Geol. Surv. Minn., Minn. 1975: Geology and age of the Wolf River Batholith, SOUTHWICK, D. L., & SIMS, P. K., 1979: The Vermilion Wisconsin. Bull. geoL Soc. Am., 86, 907-914. Granitic Complex—a new name for old rocks in northern Minnesota. Prof. Pap. U.S. geol. Surv., VAN SCHMUS, W . R., & WOOLSEY, L . L . , 1975: Rb-Sr geochronology of the Republic area, Marquette 1124-A. County, Michigan. Can. J. Earth Sci., 12, STEIGER, R . H . , & JAEGER, E . (COMPILERS), 1977: S u b 1723-1733. commission on geochronology: Convention and use of decay constants in geo- and cosmochrono- WINDLEY, B. F . , 1977: The Evolving Continents. Wileylogy. Earth planet. Sci. Lett., 36, 359-362. Interscience, New York. STOCKWELL, C . H . , MCGLYNN, J . C . , EMSLIE, R. F . , SANFORD, B. V . , NORRIS, A . W . , DONALDSON, WINDLEY, B. F . , & BRIDGWATER, D., 1971: The evoluJ. A . , FAHRIG, W . F . , & CURRIE, K. L . , 1970: G e o tion of Archaean low- and high-grade terrains. logy of the Canadian Shield; in Douglas, R. J. W. Spec. Pubis geol. Soc. Aust., 3, 33-46. (Ed.) Geology and economic minerals of Canada. Econ. Geol. Rep. Can. geol. Surv., 1, 44-150. WOODEN, J. L., 1978: Rb-Sr isotopic studies of the Archaean rocks of the eastern Lac Seul and Kenora STUCKLESS, J. S., & GOLDICH, S. S., 1972: Ages of some Precambrian rocks in east-central Minnesota-Pt. I area, English River subprovince, Ontario; in (Abstr.). Inst. Lake Superior Geol., 18th Ann. Smith, I. E. M., & Williams, J. G. (Eds) Proc. 1978 Abstr. and Field Guides, Houghton, Michigan, Archaean geochem. Conf, 131-150. Univ. Toronto Pap. 31. Press, Toronto.
98
SIMS, P. K. & PETERMAN, Z. E., 1976: Geology and Rb-
ARCHAEAN CRUSTAL EVOLUTION—A GLIMPSE FROM THE SLAVE PROVINCE W. A. Padgham
Geology Office, D.I.A.N.D., P.O. Box 1500> Yellowknife, N.W.T. XIA 2R3, Canada ABSTRACT The Slave Structural Province, which lies on the northwest side of the Canadian Shield, comprises 6 large supracrustal basins and remnants of a number of smaller ones. These basins are filled with Yellowknife Supergroup supracrustals and are set in a sea of plutonic rocks. There are 22 more or less distinct volcanic belts marginal to the basins which are filled with greywacke and mudstone. Volcanic belts in the west, south and northeast parts of the Slave Province are dominantly basaltic and elsewhere are dominantly felsic. Ultramafic lavas have not been found and mafic rocks are uncommon. The Yellowknife Supergroup contains numerous artifacts of pre-existent granitic rocks that formed a widespread crust on which the supracrustals accumulated by upwelling of volcanic material along fractures formed through local thinning and block faulting. These fractures became the margins of gently subsiding basins that filled with turbidites eroded from the basinmargin volcanics and from uplifted blocks of granitic basement. The sediments comprise 80% of the supracrustal rocks by area. Exposures of pre-Yellowknife Supergroup granitic rocks, the lack of ultramafic lavas, the Calc-alkaline nature of the bulk of the volcanics and the presence of felsic volcanics and sediments beneath some of the basaltic volcanic belts suggest that Yellowknife Supergroup ' volcanism was ensialic and that the volcanic belts formed on continental rather than oceanic - crust. _ ' The supracrustals were intruded by granitic rocks and metamorphosed to greenschist and middle-amphibolite facies around 2500 Ma. Since then the Slave Province has been a stable area only peripherally affected by Proterozoic metamorphic and igneous events.
INTRODUCTION The Slave Structural Province, an elliptical area 510km wide and 710km long, is the second largest Archaean structural block of the Canadian Shield. The major part of the Slave Province is 172 500 km in area. Two smaller segments of the Province—the Bathurst Block comprising 16250 km , and the Tree River Segment, 1250 km , are separated from the main part of the province by continental clastic sediments (Campbell & Cecile, 1976; Hoffman et al., 1978). Low-pressure metamorphism and magmatism at approximately 2.6 Ga, equivalent to the "Kenoran" Orogeny (Frith, 1978), converted extensive supracrustal basin fillings to metamorphic rocks, remobilized sialic basement, and emplaced numerous granitic plutons that together produced the Slave Craton. About 34 of the Slave Province is underlain by Yellowknife Supergroup or its migmatized equivalents and no more than 20% of these supracrustals are of volcanic origin. Recent mapping has identified numerous areas where granitic basement to the Yellowknife supracrustals is sus2
2
2
Spec. Pubis geol. Soc. Aust., 7 (1981)
pected. Age determinations, based on Rb-Sr and Pb-U, have clearly identified prexisting sialic crust that was metamorphosed during a "PreSlave" event about 2.9 to 3.1 Ga (Frith, 1978; Krogh & Gibbins, 1978; Nikic etal., 1975). Various aspects of the geology, metamorphism and tectonics of Slave Province are discussed by Thompson (1978), Frith (1978), Nielsen (1978) and Fraser (1978) in a publication of the Canadian Geological Survey entitled 'Metamorphism in the Canadian Shield". , Numerous maps and reports in various Geological Survey of Canada series give important details and descriptions of the geology of the Yellowknife, Indin Lake, Hackett River, Back River-Beechey Lake and Acadia Bay areas. Baragar & McGlynn (1976) reviewed the status of knowledge about the Archaean basement in the Slave and in the other structural provinces of the Canadian Shield. Lambert (1977, 1978) presented detailed field studies of four of the volcanic belts shown in Figure 1. Henderson (1981) has reviewed the Yellowknife Supergroup and concludes from the 4
W. A. PADGHAM 100 The summary description of the main groups limited available zircon ages that it was deposited between 2650 and 2680 Ma ago, and may have in the Southwestern Slave Province (from Frith, formed during an even shorter period—15 to 20 1978) which follows, can be applied with minor additions throughout the province. million years. Summaries of mineral developments and "(1) A granitoid basement consisting mainly of tonalitic to trondhjemitic gneiss with some exploration are available in Geological Survey (to granite and granodiorite is about 3000 Ma 1968) and Indian and Northern Affairs (1969 onold. The rocks are commonly migmatized wards) yearly summaries of the activities of the or metasomatized and intruded by an eastmineral industry in the North West Territories. west trending set of gabbro dykes which are locally truncated by volcanic belts about GEOLOGY 2700Ma old. Rocks near the volcanic belts Geological Setting may be extensively mylonitized. The Slave Province comprises 6 large supra(2) Yellowknife Supergroup volcanic rocks crustal basins and remnants of a number of which presumably overlie the basement are smaller ones set in a sea of plutonic metamorphic about 2700Ma old. Pillow tops face away and igneous rocks. There are numerous, more or from the basement-volcanic contacts. The less distinct, volcanic belts, 22 of which have been volcanic belts are variable in thickness and identified in Figure 1. Most belts in the west, distribution, due principally to the effects south and northeast parts of the province are of tectonism. On a macroscopic scale the dominantly basaltic; a few in the north and belts, as much as 10000m thick,'present a central-east parts of the province are dominantly pinch and swell appearance. They are made felsic. Where extensive areas of supracrustal up mainly of pillowed basalts capped with rocks are preserved, the volcanic belts lie along about 10 per cent (by area) rhyolitic brecthe margins of broad basins filled with turbidite cias and tuffs. of greywacke and mudstone. Numerous granitic (3) Yellowknife Supergroup metasediments, diapirs have intruded the supracrustal basins and about 2700 Ma old, for the most 'part far more extensive granitic batholiths lie between overlie the volcanic rocks. Locally, thin them. sequences of volcanic rocks are interRemnants of a sialic basement have now been fingered with the sediments. At the contact, identified in 6 localities based either on isotopic volcanic pebble conglomerate and volcanicages of approximately 3.0Ga (Frith et al., 1977; granitoid-pebble paraconglomerate may be Krogh & Gibbins, 1978; Nikic et al., 1975) or on present. The true thickness of the metasediidentification of gneissic fabrics that predate the mentary pile is unknown, due in part to Slave Province volcanic sequence. Locally intrutight isoclinal folding. sive relations (Hauer, 1979; Davidson, 1972, p. (4) Granitic and migmatitic rocks, about 114) suggest metamorphism and intrusion in the 2500 Ma old, intrude as plutons, as migmabasement before the onset of Yellowknife Supertitic-supracrustal-plutonic complexes, and group volcanism. as remobilized 2700 or 3000 Ma old The Yellowknife Supracrustal Basin (Padggranitic-migmatitic complexes. The emham, 1981a), which contains type sections of the placement of these rocks was the last Yellowknife Supergroup, is the largest, most Archaean event to have taken place. Most completely preserved and most studied basin in contain a fabric that can be related to their the Slave Province. It has been used as a model uplift, particularly at their margins. for other Slave Province Basins, but as know- (5) Minor intrusive diabase dykes, granoledge increases it has become apparent that most dioritic stocks and granitic pegmatites, basins display numerous singular attributes. about 1960 Ma old, make up less than 5 per cent of the total surface exposures. These Description of Rock Types rocks were formed as a result of ProteroThe distribution of six main rock types which zoic tectonic events related in time, if not compose most of the Slave Province is shown in origin, to activity within the Coronation Figure 1, which was drawn from McGlynn's Geosyncline." (Frith, 1978, p. 105.) 1:1000000 summary map (McGlynn, 1977). Felsic volcanics predominate locally. Though Basement and Pre(?) Yellowknife Supergroup less common, carbonates (mainly dolostones), Rocks iron formations and conglomerates are extremely Basement to the Slave Province is dominantly important in the interpretation of Slave Province granitoid but migmatized supracrustals are evolution. present in the Point Lake Area (Henderson,
CRUSTAL EVOLUTION IN SLAVE PROVINCE
1977). Krogh & Gibbins (1978) obtained a minimum age of 2730 Ma for first-generation zircons from a migmatite collected west of the granitic basement at Point Lake. This date suggests that the rocks were metamorphosed before accumulation of the Yellowknife Supergroup, provided deposition of that supergroup took place in ±20 million years as suggested by Henderson (1981). Diatrernes with gneissic basement fragments in the Con Mine (Fig. 2,A) and with granodiorite clasts at Berry Hill, 12 km north of Yellowknife, attest to granitic rocks beneath the Yellowknife Supergroup. Recent mapping of the south end of the Yellowknife Volcanic Belt (Helmstaedt et aL, 1979, 1981) defined a unit of supracrustal rocks comprising 750 m of metamorphosed basic tuffs and flows capped by 75 m of greywacke, siltstone, conglomerate, and felsic volcanic rocks. These lie between the basalts of the Kam Formation and the western granodiorite. Mapping by Goff at the north end of the Yellowknife Belt identified a sequence of felsic volcanics with thin beds of iron formation beneath the Kam Formation (Hauer, 1979). These felsic rocks appear to be folded into a tonalitic granitic complex that is cut by metadolerite dykes which do not cut the supracrustal rocks of the volcanic belt. Felsic Sediments Polymictic conglomerates with granitic cobbles have now been reported from 8 volcanic belts in the Slave Province (locations 1, 3, 4, 6, 9, 10, 14, 16 and 18, Fig. 1). Most of these polymictic conglomerates contain a mixture of granitic, volcanic and sedimentary clasts. Granitic cobbles that have been dated (Green & Baadsgaard, 1971; Krogh, personal communication) give ages around 2.7 to 2.6Ga, which combined with the field relations of the conglomerates to the enclosing supracrustals suggest that the granite cobbles were eroded from sub-volcanic intrusions which were uplifted more or less synchronously with the later phases of volcanism (Henderson, 1975a, p. 45). Tirrul & Bell (1980, p. 160) have interpreted the strongly discordant conglomerate in the Anialuk Volcanic Belt as being later than all volcanism. Felsic sediments occur beneath the volcanics of the Hackett River Volcanic Belt (Frith & Hill, 1975; Jefferson et al., 1976; Percival, 1979) and are covered by mafic, locally pillowed basalt and carbonate-rich layers (Fig. 2B) that are locally dolos tones. Near Takijuq Lake (14, Fig. 1) the 'Tak' granite has a broad sheared zone against strongly deformed sediments which include graniteboulder conglomerate (Fig. 2C, D). Cobbles of
101
quartz-monzonite porphyry as much as 23 by 60 cm are present in the conglomerate which, where the large granitic cobbles abound, seems to have escaped or resisted the intense polyphase folding seen only 100 m to the north (compare Fig. 2C, D). Zircons from one of the large quartz-monzonite boulders gives a lead-uranium age around 2.65 Ga (Krogh, pers. comm.). The petrography of the boulders and the "zircon" date suggest they were eroded from a synvolcanic pluton. The sediments are interbedded with felsic volcaniclastics, which include quartz-eye rhyolites, and are overlain by pillowed basalts. In the easternmost lens of conglomerate the cobbles are predominantly fine-grained cherty or silicified material, and iron sulphides abound in the matrix at the top of the conglomerate. A black graphitic slate overlies and is interbedded with the top part of the conglomerate unit which has been traced south for 6.5 km. Here and elsewhere in the Slave Province, basic volcanism is later than the felsic volcanism which is associated with clastics containing granitoid boulders. The graphitic slate and abundant iron sulphide indicate a distinct break in clastic accumulation prior to basaltic volcanism. Felsic Volcanic Rocks In most Slave Province volcanic belts, basalt and andesitic basalt comprise 90% or more of the volcanic rocks. Belts containing a considerably higher percentage of felsics (20 to 40%) include (on Fig. 1) 4 (Henderson, 1976), 6 (Heywood & Davidson, 1969), 11 (Bau et al., 1979), 14 (Hyde et al., 1976), and 16 (Tirrul & Bell, 1980). Others, composed predominantly of felsic volcanics include 18 (Frith & Hill, 1975), 19 (Lambert, 1978, p. 153), and 15 (Padgham et al., 1974). It has been argued that the High Lake and Hackett River Belts are anomalously felsic because the underlying basalts have been destroyed by assimilation, uplift and erosion or some other event capable of selectively disposing of vast sections of basalt. Tirrul & Bell (1980, p. 161) suggest that the Anialuk Belt is a remnant of the mafic part of the High Lake Volcanic Belt. Though this is possible, the Anialuk Belt is only 1/10 as large as the High Lake Belt and is about 20% felsic. Were it ten times larger, felsic volcanics would still constitute 50% or more of the combined belts. On the north and northeast sides of the Slave Province, felsic volcanic rocks make up a major part of some volcanic belts. The High Lake Belt (15, Fig. 1) is more than 50% dacite to rhyolite and has an extensive unit of coarse rhyolitic breccias and (?) agglomerate at its top in the immediate High Lake area. Hackett River Belt is similar
102
W. A. PADGHAM
U2b00'W
63°00 N
SLAVE
»l2o00' \N
C R U S T A L E V O L U T I O N IN SLAVE P R O V I N C E L E G E IO
P R O T E R O Z Q I C
'
Supracrystal -rocks^,dominant ly conti nental clastlcs; main 1 y "Epworth and Sou1 bur n Groups" and Great Slave Supergroup. Bear and Churchill Prdvin.ee Plutonic rocks
ARC
tr £ A Ji
103
16. Anialik River, gold-quartz veins in plutonics 17. H o p e Bay, epigenetic native, silver and gold vein deposits, volcanogenic base metals ! 18. Hackett River volcanogenic silver-zinccopper-lead deposits 19. Back River Volcanic Complex 20., Musk ox Lake gold deposits 21. Clinton-Colden Lake volcanogenic zinc showings 22. Walmsley Lake
-Intrusive .granitic rocks of Slave Province 'Kenoran'
m a 11Orpgeny (2.7 to 2.6 GA). .:, • Migmatized'YelTqwIfnife Supergroup with granite gneiss and undifferentiated intrusions.
i d ]
Belts of dominantly felsic and intermediate volcanics. Mainly greywacke-mudstone turbidites ,. Belts of dominantly basaltic volcanics - Plutonic rocks; granite gneisst migmatite undifferentiated intrusive igneous rocks. -In part basement to Yellowknife Supergroup.
but has a thin unit of basaltic flows and breccias at or near the top of the succession for much of its 100km length. Back River Volcanic Complex (Lambert, 1978) contains only minor volumes of basalt, and roughly equal proportions of rhyolite, dacite and andesite. Felsic volcanics are equally abundant at the base of these belts as they are at the top. In the Hackett Belt, most of the basalt lies near or at the top of the belt.
. Slave Province BoundaryMajor faults
Fig, 1.
M a p of the Slave Structural Provin.ee showing distribution of m a j o r rock types. Inset m a p , Slave Province and C a n a d i a n Shield. N u m bered volcanic belts, most of which are referred to in the text, and the main economic mineral deposits, are: 1. Yellowknife with shear - zone gold . deposits ( G i a n t a n d Con Mines) 2. Thislewait Lake, gold in quartz veins (Discovery Mine) 3. C a m e r o n River,, gold quartz veins in near by sed iirren t s (Ca m lar en Mine) 4. Turnback . Lake, silver-lead sulphide deposits 5. Beaulieu River 6. Indian M o u n t a i n Lake, zinc-lead, silver and copper volcanogenic sulphide 7. Cam sell Lake 8. McKay L a k e , gold in quartz veins (Tundra and Balmita Mines) also yolcanpgenip zinc deposits 9. Indin Lake, gold in quartz veins and dykes 10. Point Lake, volcanogenic massive sulphides 11. Itchen Lake, volcanogenic zrnc-silverleacp-cppper sulphide deposits including Izpk Lake cjeposit 12. Redrock La^g 13. C o n t w o y t o y | k e with syngenetic amphibolite dep0§?ts in nearby sediments 14. Takijuk Lake, volcanogenic zinc-coppersilver deposits 15. High Lake with High Lake volcanogenic copper-gold deposit
Carbonate and Iron Formation Thin but persistent dolomitic carbonate units and more discontinuous iron formations are common- along the volcanic belt-greywacke transitions and are most conspicuous in the more felsic belts. Lambert (1978, p. 155) shows photos of the oolitic carbonate from the Back River Complex where iron formation is well developed. Lambert (p. 156) notes "Carbonate that fills fractures and impregnates breccias yet forms massive to layered pods and lenses, may be related to tufa and sinter deposited by circulating groundwater in highly porous terrain during the waning stages of volcanism." Carbonate beds, prominently developed at Snofield Lake at the top of the High Lake volcanic belt (Padgham et aL, 1974), contain stromatolites (Fig. 2E) (Henderson, 19756). During a collecting trip to this locality in 1979 columnar stromatolites were identified by H. J. Hoffman. Iron formations are commonly thin and discontinuous; typically a metre or twp thick several hundred metres long. The extensive Algoma-type iron formations that range up to 30 m or more in thickness (Gross, 1970) are not represented in the Slave Province, The predominance of clastic sediment over volcanic rocks in this province and the much smaller size of the supracrustal basins compared to some in the Superior Province suggests.(1) that iron was not separated effectively from clastic debris and possibly (2) that iron was not as abundant in the volcanic emanations because of their more sialic provenance.
W. A. P A D G H A M
104
Fig. 2.
A
B
C
D
E
F
A: Diatreme containing angular and subrounded clasts of tonalitic gneiss, Con Mine. Clasts from this dyke gave Pb-U ages for zircons of 3.03 Ga (Nikic et al., 1975). Each segment of rule is 15 cm long. B: Carbonate-rich layers in strongly folded siliceous volcaniclastic near the top of the Hackett River Volcanic Belt. C: Quartz monzonite porphyry cobbles and boulders in the western phase of the 'Tak' Lake conglomerate. Largest cobble is 20 cm in diameter. D: Deformed cobbles in the west lens of the 'Tak' Lake conglomerate. E: Stromatolite at Snofield Lake (Henderson, 19756). An oncolite is visible near the top of the photograph. Scale is 17 cm long. F: Cleavage in cordierite knots at a distinct angle to cleavage in the surrounding metaturbidite, Hackett River area. Larger cordierite porphyroblasts are 2 cm in diameter.
105 CRUSTAL EVOLUTION IN SLAVE PROVINCE Some of the most laterally extensive magnetite gravitational collapse before deposition of the iron formation in the Slave lies along or near the youngest facies of the Yellowknife Supergroup. top of the Hackett volcanics where it forms a Shearing of the borders of marginal plutons and fairly continuous but thin unit, rarely more than the basal coarse clastic sediments of the Yellowa metre thick, that extends for many kilometres. knife Supergroup, as at 'Tak' Lake, suggests that It is no doubt a clastic sedimentary unit as it is such shearing took place after basin-margin thinly bedded and lies within mudstone and grey- granitic intrusions had been emplaced, uplifted wacke a few hundred metres above the volcanics. and eroded to produce granite-boulder conToward the north part of the belt the magnetite glomerate. iron formation appears to be replaced by a Fyson (1978) has discussed the fold and pyrrhotite-rich, sulphide-facies iron formation cleavage patterns in parts of the Yellowknife that occurs as thin lenses in black muddy siltstone Supergroup and concludes that the patterns disjust above the volcanics. played are not related to diapiric uprising as argued by Drury (1977). Fyson has identified a Structure northwest-trending cleavage of regional extent The structural framework of the Slave Pro- related to a third deformation phase that took vince is well established but details are poorly place after some plutons had reached their understood as few areas have been the object of present level because earlier structures are warped around the early plutons. A later phase of granite modern detailed structural analysis. Formation of shallow basins by extensional plutons cuts across the structures with little thinning of a + 3.0Ga continental crust with deflectipn of them. block faulting (Henderson, 1981; Lambert, 1978) Structures in the remnants of pre-Yellowknife that initiated volcanism in linear belts along the Supergroup basement are even more complex, basin margins was followed by continued margin and little work has been done on them. uplift, Toward the end of this cycle, gravity Tectonism ceased about 2.6 Ga with the intrusliding of parts of the volcanic belts from the ris- sion of the post-tectonic granitic rocks that cut ing margins into the basins initiated folding of across all pre-existing elements. In the Yellowthe turbidite basin-fill. knife Basin these are accompanied by a group of As the basins were meridionally elongate, the rare-element-bearing pegmatites. Widespread trend of fold axial planes was also meridional. fracturing took place and dolerite dykes were Continued basin-margin uplift, to permit unroof- intruded in swarms trending east-northeasterly ing of subvolcanic plutons, was accompanied by (2.3-2.4 Ga) mainly in the south part of the widespread intrusion of granodioritic magmas province, northwesterly and northeasterly and the basins were compressed laterally to (2.0-2.1 Ga) mainly in the west, and north-northtighten folds initiated by gravity sliding. westerly (1.2Ga) throughout the province Marginal volcanic belts were turned on their sides (McGlynn & Henderson, 1972). to form the typical homoclines in which they are Post-Yellowknife Supergroup deformations now found, and small plutons invaded the supra- have strongly affected much of the Bear-Slave crustals mainly by diapiric upwelling which boundary (Frith, 1978) and have brought domes of Archaean rocks with pervasive Aphebian added to the compression of the sediments. The resulting structures in the Yellowknife deformation structural overprints into the border Supergroup tend to be extremely complex with zone of the Bear Province (Nielsen, 1978; Frith et al., 1977). multi-episodic deformation evident (Fig. 2F). In most belts the volcanic rocks form marginal Rifting along the Bathurst and MacDonald homoclines, but the metaturbidites display com- (East Arm) fault zones was probably caused by plex fold and cleavage patterns resulting from compression of the main body of the Slave block superposition of at least two additional deforma- against the Churchill Province by continental coltions. The effects of repeated deformation are lision to the west of the Slave Province during more difficult to recognize and harder to inter- the development of the Wopmay Orogen (Hoffpret in most volcanic belts because of a paucity of man, pers. comm.). Extensive dolerite sheets, fed marker features. There are no obvious later folds mainly from faults developed along the edge of in most volcanic belts to ascribe to the later defor- the province, intruded along or near the unconmations, but the Con and Giant shear systems formity between the Archaean and the Aphebian may be examples of shear folds with translation cover. parallel to bedding as described by O'Driscoll Early Aphebian intrusion related to the forma(1962, 1964). Henderson (1978) has argued that tion of the Coronation Geosyncline took place the shearing that formed mineralized 'schist along the south and southwest sides of the prozones" at Yellowknife was at least initiated by vince and large parts of the province became 4
W. A. PADGHAM
106
areas of accumulation of continental and shallow marine sediments. Uplift of the area took place in the Helikian (approximately 1.0 Ga) and the Slave Province assumed its present form. It has been a positive area undergoing erosion ever since.
Metamorphism
The supracrustal rocks of the Yellowknife Supergroup display mainly low-pressure metamorphism of low to medium metamorphic grade. Thompson (1978) reviewed in considerable detail the metamorphism of the Province concluding Pyrenean-type regional metamorphism reached a maximum during the final stages of the 2.6 Ga tectonism. He notes (p. 88) "Textural relations and the fact that isograds transect S and the axial surface of major folds indicate that the peak of metamorphism occurred after the main phases of folding . . Cordierite porphyr©blasts which are typical of middle amphibolite-grade knotted schists appear spatially related to, but are slightly earlier than, late granitic intrusions in the supracrustal basins. Typical assemblages in these rocks include quartz, plagioclase, biotite and cordierite with or without andalusite, gedrite, garnet and muscovite. Large, randomly oriented, euhedral crystals of staurolite and andalusite are present in the aureoles of middle to upper amphibolite-facies metamorphism in the Hackett River Supracrustal Basin. In lower-grade rocks quartz, plagioclase, biotite, muscovite, chlorite, and in some rocks chloritoid, garnet, andalusite, or staurolite, may be present. With increasing grade, sillimanite appears locally. Biotite, quartz and plagioclase remain in abundance, whereas chlorite, staurolite and gedrite disappear. Muscovite, cordierite, andalusite and garnet may remain and K-feldspar may appear. Kyanite is uncommon, being found only as relicts in sillimanite-bearing rocks along the east side of the Slave Province where it is related to local basement uplift (Percival, 1979) or broad uplift along the Churchill Province boundary (Fraser, 1968, 1972, 1978; Henderson & Thompson, 1980). Low-grade metamorphic zones within areas of medium grade that are surrounded by high-grade rocks are distributed randomly across the main structural trends of the Slave Province because the metamorphism outlasted the main phases of deformation. Metamorphism of the basement enclaves is extremely complex as all such areas have been partly remobilized or at least re-metamorphosed during the 2.6 Ga cratonization of the province. 2
Few of the younger migmatite areas have had their metamorphism or structure adequately studied.
METALLOGENY OF THE SLAVE PROVINCE The Slave Province is abnormally rich in gold, as are many Archaean terrains, and has produced just over 11.2 million ounces, of which approximately 88% was from mines in the Yellowknife Volcanic Belt (Padgham, 19816). Massive and volcanogenic base-metal deposits have been found in at least 8 of the Slave Province Volcanic Belts. One of these, the High Lake deposit, is a Cu-Au,( + Zn) deposit of the Noranda type. Most of the others contain considerable silver (2 to 50 oz/ton), abundant zinc with copper and relatively high percentages of lead, but little gold. Silver is not only abundant in most of the volcanogenic polymetallic sulphide deposits of the Slave Province, but it is an important constituent of most gold ores as well. Boyle (1961, p. 164) reports gold/silver ratios for the Yellowknife Volcanic Belt deposits as 5:1 and for the quartz veins in the sediments (Yellowknife Supergroup greywacke-mudstone-turbidites) 3.5:1. Not only is the Yellowknife Supracrustal Basin an area with a high clarke of concentration for gold, but it contains numerous pegmatites with lithium, beryllium, tin, tungsten and tantalum. Silver is also found abundantly in vein-type deposits in the younger rocks surrounding the Slave Province, mainly in the Camsell RiverEcho Bay Districts on Great Bear Lake. Similar, rich, native-silver deposits have been superimposed on Slave Province basalt of the Hope Bay Volcanic Belt (17, Fig. 1). Native silver and nickel-cobalt arsenides which commonly accompany the Great Bear silver deposits, form veins in the Aphebian rocks of the Athapuscow Aulacogen or East Arm Fold Belt (Davidson, 1972; Badham, 1979) and are found also in late veins cutting the Slave Province rocks along the north side of the East Arm of the Great Slave Lake. The most productive gold deposits in the Slave Province are epigenetic, developed in quartzcarbonate-chlorite-sericite shear zones in the basaltic volcanics of the Yellowknife Volcanic Belt. Epigenetic gold-quartz veins in greywacketurbidites are the commonest and most widespread type of gold deposit in the Slave Province. There are nearly 80 of these in the Yellowknife Supracrustal Basin alone (Padgham, 1981a). Syngenetic deposits in amphibolitic iron formation abound in the Contwoyto Lake area (25 showings) (13, Fig. 1). Gold-quartz veins in metamorphic-plutonic rocks are common (23 + show-
CRUSTAL EVOLUTION IN SLAVE PROVINCE
ings) in the Acadia Bay area, east of the Tree River Segment on the coast of Coronation Gulf. Numerous researchers (Allison & Kerrich, 1981; Kerrich, 1981; Myers, 1981) contend that the Yellowknife deposits are a product of metamorphic-hydrothermal deposition with C0 -rich gold-bearing solutions derived by metamorphic differentiation of basement rocks or the lower parts of the volcanic belts. If all the gold-bearing solutions are of similar derivation, the epigenetic deposits in the north and south sides of the Slave Province could have originated later than the syngenetic deposits in the more central Contwoyto Lake area. Perhaps the rocks in the southern and northern parts of the province are older so that when metamorphism generated the gold-bearing solutions, structurally prepared sites were available to trap the gold. However, in the Contwoyto Lake area, structurally prepared sites were not available and the gold-bearing fluids reached the surface to mineralize sediments in areas where the rate of sedimentation was low (Bostock, 1981). The earliest dolerite dyke swarm (intruded 2.3 to 2.4Ga, McGlynn & Henderson, 1972, Fig. 3, p. 517) is confined to the southern third of the Slave Province, additional evidence that orogeny was more advanced here than in the central part of the province. The paucity of oxide and carbonate iron formations in general, and the complete lack of economic iron deposits, may mean that conditions for separation of iron did not occur. Perhaps the Slave Province supracrustal basins (epicontinental seas?) were too small to permit separation of iron from clastic material and the deposition of extensive iron formations. Alkaline complexes mineralized with elements such as rare earths, uranium, thorium, fluorine and copper intruded the margins of the Slave Province at Big Spruce Lake (Frith, 1978; Martineau, 1970) and at Blachford Lake (Davidson, 1972) during the lower Aphebian (2.3-2.6 Ga). Diatreme breccias mineralized with traces of copper and cobalt cut the margins of the Slave Province along the north shore of the East Arm of Great Slave Lake (Badham, 1979). These later mineralizations, related to evolution of the Coronation Geosyncline, cut the Slave Province apparently only along zones of transition from the 33.9km thick Archaean crust to the 37.9km thick crust beneath the Proterozoic East Arm Fold Belt (thicknesses quoted by McGlynn & Henderson, 1972, p. 508). 2
107
by stretching of a +3.0Ga granitic crust. Rifting formed graben-like basins that were filled with greywacke-mudstone turbidites (Henderson, 1981) eroded in part from uplifted basement blocks and in part from the early felsic outpourings that initiated volcanism in most areas. Extensive basic volcanism followed. Mafic magmas generated at the base of the crust were contaminated by crustal melts before they could be extruded, resulting in strong calc-alkaline trends superimposed on the basaltic volcanic piles (Baragar, 1966). Sedimentation took place continuously as mafic cones were built up to form linear volcanic belts confined mainly to the basin margins. Subsidence of the volcanic belts concomitant with melting of large portions of the pre-existing sialic crust generated granitic magmas tnai intruded the supracrustal basins, and the intervening higher levels of granitic crust. Continued uplift of the intervening, dominantly granitic, areas permitted erosion of most of the less reworked sialic basement. Compression, effectively directed toward the east, folded the Yellowknife Supergroup on northerly-trending axial planes. A late phase of compressive deformation imposed a regional cleavage across large areas of the province and warped earlier structures around plutons that had reached their present position before the end of this regional deformation (Fyson, 1978). Additional plutons were intruded during this deformation but they cut across structures. A metamorphic climax followed the last regional compression and caused the growth of randomly oriented crystals. Province-wide uplift coincided with and continued after the cratonization around 2.6 Ga. Periods of extensive fracturing tapped mafic and ultramafic magmas and dolerite dykes were intruded. Following nearly a billion years of erosion (2.6 to 1.7 Ga) broad down-warping initiated continental-type sediment accumulation in broad shallow basins now preserved as. the Kilohigok Basin (Campbell & Cecile, 1976). Since the Helikian, the Slave Province has been a stable positive area undergoing slow but continuous erosion down to the present. This billion years of erosion has not planed much more than a kilometre from the Slave Province. Gossans preserved above a number of massive sulphides suggest that erosion since the Pleistocene has been negligible (Padgham, 1977).
CONCLUSIONS GEOLOGICAL SYNTHESIS The Archaean geological environment of the Yellowknife Supergroup rocks accumulated in Slave Structural Province was vastly different from that of the high-level and low-level terrains relatively small ensialic basins apparently formed
108
W.
A.
distinguished by Windley & Bridgwater (1971). The development of such areas as the Superior Province and Yilgarn Block is dominated by ultramafic to mafic lava generation in elongate basins lying above linear fractures that tapped mantle-derived magmas. Sedimentary rocks are a minor component of such areas because there was rarely subareal exposure. Development of granulite-facies gneiss terrains is fundamentally different also because these represent deeply eroded sites of ultrametamorphism within thick sialic crust. The main characteristics of the Slave Structural Province that make it different are: (1) supracrustal basins filled dominantly with sediments formed by regional extension of a (2) thick sialic crust which permitted formation of an (3) extensive block-faulted terrain and (4) initial felsic volcanism followed by (5) mafic volcanism along the faults that bounded the basins. Extension was not sufficient to permit (6) ultramafic-magma sources to be tapped to provide volcanic outpourings nor (7) to form oceanic crust beneath any of the basins. The volcanics vary from (8) tholeiitie with a strong calc-alkaline character superimposed by sialic crust contamination to (9) dominantly calc-alkaline probably derived f r o m ensialic magmas. Metallogenically the province has a (10) high clarke of concentration for gold. It contains abundant epigenetic auriferous quartz veins and (11) unique quartz-carbonate-sericitechlorite shear zones that host large gold orebodies and has, as well, syngenetic amphibolite-ironarsenic-gold deposits in metaturbidites, Volcanogenic massive sulphide deposits are dominated by (12) high silver, lead and zinc and low gold content.
PADGHAM
Metamorphism was low pressure, low temperature in the supracrustal basins with highertemperature regimes in areas of intrusion and basement uplift. High pressures are indicated only where basement upwelling has brought once deeply buried material to the surface. The Slave Province is one type of Archaean terrain that has been ignored until recently in conjecture on Archaean crustal evolution. However, it is not the only type of Archaean terrain that has not been considered, for Schau (1977) has described a sequence of komatiites and quartzites deposited in a technically stable regime. Thus it can be expected that additional types of Archaean terrain will be described in the future. As pointed out by Henderson (1981), " I t is not useful and probably counterproductive to think in terms of a " t y p e e x a m p l e " of Archaean terrain . . . as much can be gained f r o m understanding . the differences . . . " a s can be gained from understanding the similarities. ACKNOWLEDGMENTS The writer is indebted to J . C. McGlynn, R. A. Frith, M. B. Lambert, J. B. Henderson and other members of the Geological Survey of Canada Bear-Slave project team for their willingness to share their knowledge of, and ideas on, the Slave Province. J. B. Henderson kindly provided a manuscript copy of his paper on Archaean Basin Evolution in the Slave Province. Denis Valiquette drafted the figures and Valerie H o m e typed the manuscript. Linda Padgham reviewed the manuscript and gave valuable advice on organization.
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_, 1981: Archaean basin evolution in the Slave Province, Canada; in Kroner (Ed.) Plate Tectonics in the Precambrian. Elsevier, Amsterdam.
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of Mackenzie. Geol. Surv. Can., Open File, 353. , 1977: Geology of Keskarrah Bay, Point Lake, Northwest Territories. Geol. Surv. Can., Open
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Benjamin Lake map-area, District of Mackenzie, 75M/2. Mem. geol. Surv. Can., 361. DE BIE, I., 1978: Geology of the Coronation geosyncline (Aphebian), Hepburn Lake sheet (86J), Bear Province, District of Mackenzie. Pap. geol. Surv. Can., 78-1A, 147-151. HYDE,
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Geology Hackett River, 7 6 - F - 1 6 . EGS Map 1 9 7 6 - 4 —preliminary edition, DIAND, Yellowknife. 1981: Archaean lode gold deposits: A synthesis of data on metal distribution, fluid inclusions and stable isotopes, with special reference to Yellowknife; in Morton, R. D. (Ed.) Gold Workshop Volume. Yellowknife Geo-workshop Committee, Yellowknife.
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KROGH, T . E . , & GIBBINS, W . , 1 9 7 8 : U-Pb isotopic ages
of basement and supracrustal rocks in the Point Lake area of the Slave Structural Province,
Canada. Abstracts with Programs, Geol. Ass. Can., Min. Ass. Can., 3, 4 3 8 . LAMBERT, M. B., 1977: Anatomy of a greenstone belt— Slave Province, N.W.T; in Baragar, W. R. A., Coleman, L. C., & Hall', J. M. (Eds) Volcanic regimes in Canada. Spec. Pap. geol. Ass. Can., 16, 331-340. _, 1978: The Back River volcanic complex—a cauldron subsidence structure of Archaean age; in Current Research, Part A. Pap. geol. Surv. Can., 78-1A, 153-157. MARTINEAU, M .
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chemistry of the Bigspruce Lake syenite complex, N.W.T., Canada. Ph.D. Thesis, Oxford University [unpublished]. MCGLYNN, J : C., 1977: Geology of the Bear-Slave Structural Provinces. Geol. Surv. Can., Open File,
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Province; in Price, R. A., & Douglas, R. J. W. (Eds) Variations in tectonic styles. Spec. Pap. geol. Ass. Can., 11, 506-526. 1981: Geochemistry of Con Mine, Yellowknife, N.W.T; in Morton, R. D. (Ed.) Gold Workshop Volume. Yellowknife Geo-workshop Committee, Yellowknife.
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L E E C H , A. P . , 1 9 7 5 : Diatreme containing boulders of 3 0 3 0 m.y. old tonalite gneiss, Con Mine, Yellowknife, Slave Craton; in Abstracts with programs, geol. Soc. Am., 7, 1 2 1 3 - 1 2 1 4 .
SHEGELSKI, R. J., 1974: Geology, High Lake, District of Mackenzie. Geol: Surv. Can., Open File, 208. PERCIVAL, J. A., 1979: Kyanite-bearing rocks from the Hackett River area, N.W.T.: Implications for Archaean geothermal gradients. Contr. Min. Petrol., 69, 177-184. S C H A U , M I K K E L , 1977: "Komatiites" and quartzites in the Prince Albert Group; in Baragar, W. R. A., Coleman, L. C., & Hall, J. M. (Eds) Volcanic regimes in Canada. Spec. Pap. geol. Ass. Can., 16 342-354.
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comment. Geology, 5, 395. , 1981a: Gold Deposits of the Northwest Territories; in Morton, R. D. (Ed.) Gold Workshop Volume. Yellowknife Geo-workshop Committee, Yellowknife. , 198 lb: Geology of the Yellowknife Volcanic Belt, a review; in Morton, R. D. (Ed.) Gold Workshop Volume. Yellowknife Geo-workshop Committee, Yellowknife.
P. H., 1978: Archaean regional metamorphism in the Slave Structural Province—A new perspective on some old rocks; in Fraser, J. A., & Heywood, W. W. (Eds) Metamorphism in the Canadian Shield. Pap. geol. Surv. Can., 78-10, 85-102. T I R R U L , R., & BELL, I., 1980: Geology of the Anialik River greenstone belt, Hepburn-Island map-area, District of Mackenzie; in Current Research, Part A. Pap. geol. Surv. Can., 80-1 A. W I N D L E Y , B . F . , & BRIDGWATER, D . , 1 9 7 1 : The evolution of Archaean low- and high-grade terrains. Spec. Pubis geol. Soc. Aust., 3, 33-46 . THOMPSON,
ARCHAEAN GEOLOGY AND METALLOGENY OF THE ALDAN SHIELD, USSR V. I. Kazansky & V. M. Moralev 1
1
2
Institute of Geology of Ore Deposits, Petrography, Mineralogy and Geochemistry of the USSR Academy of Sciences, Moscow Institute of the Lithosphere of the USSR Academy of Sciences, Moscow 2
ABSTRACT The Early Precambrian metamorphic complexes of the Aldan Shield can be subdivided into two major tectonic units, an Early Archaean gneiss-granulite basement, and a Late ArchaeanEarly Proterozoic supracrustal trough complex (greenstone and schist belts). The oldest basement-rock ages are considered to be more than 3500 m.y. old (Katarchaean). Basic schists (mainly orthopyroxene and two-pyroxene granulites) are the lowermost exposed horizons of the basement metamorphics. They are overlain by quartzite, high-alumina schist, carbonate rock, basic schist and biotite-amphibole-garnet gneiss. Magnesium-carbonate rocks transformed by metasomatic processes into skarns abound within some subunits of basement metamorphics. Supracrustal trough structures developed in the interval from 2700 to 2200m.y. At present more than 30 isolated troughs are recognized in the Aldan Shield and Stanovoi Ridge. The trough sequences are composed of metasediments transformed into metasandstone, quartzite, iron-rich quartzite, mica and graphite schist, marble and metaconglomerate with subordinate basic metavolcanics (amphibolite). Archaean complexes have been affected by several phases of younger endogenous processes. The most intensive granitization and metasomatic alteration of Archaean rocks occurred in the _ period 1800-2100m.y. The progressive amphibolite or greenschist metamorphism of trough sequences, as well as regressive amphibolite metamorphism of granulite basement complexes, has resulted from these tectonic-magmatic processes, which also formed ore deposits of different genetic types and ages. Archaean phlogopite deposits belong to the skarn group and are confined to peripheral parts of a vast dome-like structure being located within magnesiumcarbonate horizons. The distribution of phlogopite within distinctive lithological units is characteristic. The most prominent Early Precambrian iron deposits are recently-discovered iron-rich quartzites of the Chara-Tokko region. They are confined to a volcano-sedimentary trough sequence of N-S trend. This trough series has undergone zonal amphibolite and epidote-amphibolite metamorphism. Geological data available for the Aldan Shield and for the Archaean cratons of Australia, India, South Africa and Canada reveal major similarities of Early Precambrian structural complexes and their development in time. On the other hand, certain distinctions of Archaean volcano-sedimentary sequences with respect to primary composition, degree of metamorphism, metasomatic transformation and, consequently, in metamorphogenous oreformation, can be established.
REGIONAL SETTING The Aldan Shield, taken together with the Stanovoi Ridge fold belt adjoining it in the south, is the largest metamorphic terrain in the USSR and is of great interest from the viewpoint of Archaean geology and metallogeny (Fig. 1). The Early Precambrian metamorphic complexes of this region can be subdivided into two major units, an Early Archaean gneiss-granulite basement and a Late Archaean-Early Proterozoic trough complex (greenstone and schist belts). Upper Proterozoic and Phanerozoic rocks form a sedimentary cover within isolated depressions in
Spec. Pubis geol. Soc. Aust., 7 (1981)
the metamorphic complex. Jurassic-Cretaceous granitic batholiths compose up to 60% of the exposed Stanovoi Ridge Province, while the dioritemonzonite intrusions of .the same age within the Aldan Shield may be classified as anorogenic, sharply discordant, small bodies emplaced at shallow depths. Cretaceous basalts are known at places in the Stanovoi Ridge area. GNEISS-GRANULITE BASEMENT The metamorphic Yengra and TimptonDzheltula Series define the gneiss-granulite basement of the Aldan Shield, and the Kurulta-
V. I. KAZANSKY & V. M. MORALEV
112
cz]< EU
&
m*
nn] a . 5
czi. G3*
Fig. 1. Outline geological map of the Aldan Shield and the Stanovoi Ridge: 1—Cretaceous volcanics; 2—Jurassic-Cretaceous granitic batholiths; 3—Upper Proterozoic and Phanerozoic sedimentary cover; 4—Lower Proterozoic sedimentary cover (Udokan Series); 5—Mongol-Okhotsk Phanerozoic foldbelt; 6—Upper Archaean-Lower Proterozoic greenstone and schist belts (trough complexes); 7—Lower Archaean-Katarchaean gneiss-granulite basement; 8—major faults; 9—elements of ring structures as seen on satellite imagery.
Gonam and Cupura Series of the Stanovoi Ridge are considered to be their analogues. A wholerock Pb-Pb isochron age of 3300 ± 200m.y. represents the minimum age of the rocks of the gneiss-granulite basement (Rudnik & Sobotovich, 1969). Older, but not so reliable, Rb-Sr and K-Ar whole-rock ages (3960±35, 4200 ± 500m.y.) have also been obtained (Brandt et al., 1978; Polovinkina, 1973). It is considered that the gneiss-granulite rocks of the basement are Katarchaean, and perhaps as old as 3500m.y. (Table I).
The lowermost exposed horizons of the Yengra and Kurulta-Gonam Series of the basement are represented by orthopyroxene and two-pyroxene basic schists with a measured thickness of up to 5 km. They are overlain by quartzite and highalumina (garnet, sillimanite) schist and gneiss with horizons of carbonate or diopside rocks and corundum lenses. The total thickness of the unit
varies from 2 to 6 km (the Upper Aldan and Fedorov Suites). The Timpton-Dzheltula Series consists of biotite, biotite-garnet and biotiteamphibole gneiss and basic schist with subordinate beds of marble and calciphyre. The total thickness of the series, which is considered to rest unconformably on the Yengra Series, amounts to more than 10km. The gneiss-granulite basement rocks are metasedimentary and metavolcanic in origin. Inclusions of Iherzolites and lenticular bodies of melanocratic sapphirine-bearing rocks occur rarely in basic metavolcanics. Metamorphic conditions of the basement rock units are P = 10-11 kb and T =• 1000-1100 °C for lower horizons and P - 9.0-9.5 kb and T = 820-900°C for upper horizons (Kastrykina & Karsakov, 1977). Petrochemical studies of the basic schists (metavolcanics) show that most have close affinities with oceanic tholeiite (Glukhovsky et al.,
GEOLOGY A N D METALLOGENY OF ALDAN SHIELD
113
TABLE I
Early Precambrian rock units Aldan Shield
Stanovoi
Ridge
Proterozoi c
:
Lowe r
Udokan series
Dzheltulak series
Tasmi eli n seri es vn o 0 ^ 01
Ust-Gilyui
series
Borsala series
<L> Q. CL
Timpton-Dzheltula s
series
Gupura series Stanovoi
3000 Ma
nj -C o <
Complex Idzhak suite <u 3 O
'
Vengra series 3500 Ma
Fedorov suite
KurultaGonam series
Upper Aldan suite Gorbylyakh suite
Katarchaean
1977). However high-magnesian basalts and ultramafics with basaltic and peridotitic komatiite chemistry as well as calc-alkaline basalts are also present (Figs 2, 3). These rocks are characterized by low 87Sr/86Sr initial ratios (0.70090.7040). Where biotite, phlogopite and hornblende appear in significant amounts in metavolcanics subjected to alkaline metasomatism and granitization, S r / S r ratios are considerably higher with values up to 0.7152-0.7972 (Brandt et al. \ 1978). The gneiss-granulite basement of the Aldan Shield is characterized by large domal or oval structures, separated by zones of compressed o,9Fe.0+FeO 87
86
I >
CaO
AlA
Fig. 3. C a 0 - M g 0 - A l 0 diagram for basic schists (metavolcanics) of the gneiss-granulite basement: 1—peridotitic komatiite; 2—basaltic' komatiite of the "Geluk type"; 3—basaltic komatiite of "Barberton and Badplaas type"; 4—tholeiitic basalts. 2
3
linear folds. Broad synclinoria are observed locally. These structures are recognized to be the remnants of older structural forms ( Glukhovsky & Pavlovsky, 1973). AFM plot for basic schists (metavolcanics) of the gneiss-granulite basement. Solid lines outline volcanic series of H. Kuno: I—tholeiitic, II—alkali olivine (calc-alkaline) basalts; the broken lines outline the field of mid-ocean ridge basalts.
SUPRACRUSTAL ROCKS Trough structures have only recently been recognized on the Aldan Shield and in the Stanovoi area (Moralev, 1978). The character of their volcano-sedimentary rocks and their tectonic setting indicates affinities with the late
114
V. I. KAZANSKY & V. M. MORALEV FeO 0,9 Fe 0, +
z
Fig. 4. AFM plot for basic metavolcanics of various greenstone belts. Symbols as defined in Figure 2.
stones of the Udokan Series is 2100m.y., whereas that of pegmatites from pebbles in conglomerates of the Dzheltulak Series is 2300 m.y. It is suggested that the development of the trough sequences took place during the period from 3000-2800 to 2300m.y., i.e. in the Late Archaean and Early Proterozoic (Polovinkina, 1973). The supracrustal rocks within the troughs are metamorphosed under amphibolite- and greenschist-facies conditions, commonly with decrease in grade of metamorphism up the section. Thermodynamic conditions of metamorphism vary over a wide range with P = 2.5-4.0 kb and T = 350-650 °C (Beryozkin, 1977), The trough structures which have lengths ranging from 30 to 150 km and widths ranging from 3-5 to 20-25 km, are confined between N-S, NWSE and E-W major regional faults. Extended systems of schist or greenstone belts or inliers are also observed. The trough sequences differ considerably from each other, suggesting that they were not connected during sedimentation and development of the basins which formed at an early stage as rift (trough) depressions similar to minor oceans. The volcano-sedimentary sequences are now folded into narrow elongated synclinoria, partly complicated by thrust faults and transformed into steeply-dipping lens-like monoclinal inliers. Folding of the trough sequences was accompanied by influx of alkalis and by the additional growth of dome structures in the basement (Zonenshain etal., 1976).
Archaean-Early Proterozoic supracrustal (greenstone and schist) belts of the Canadian, Indian and other shields. At present about 30 trough structures comprising metamorphosed volcanosedimentary sequences with measured thicknesses of 5 to 12 km are known within the Aldan Shield and the Stanovoi area. In most troughs, sedimentary rocks predominate although some are predominantly volcanogenic. Metamorphosed sedimentary rocks are represented by metasandstone, quartzite, micaamphibolite, cordierite and garnet schist and microgneiss, and iron-rich quartzite, and more rarely by graphite schist, marble and conglomerate. Metavolcanics (amphibolite and amphibole schist) are mainly basic rocks which occur predominantly in the lower parts of the volcanosedimentary series. AFM and C a 0 - M g 0 - A l 0 diagrams (Figs 4, 5) show that the greenstone metavolcanics mostly plot as tholeiitic and calcalkaline basalts. However some metavolcanics have close affinities with basaltic and peridotitic komatiites. In some troughs, a pile of psammite, mica and graphite schist, pebbly sandstone and conglomerate with intercalations of intermediate and basic volcanics (the Dzheltulak Series) rests unconformably on the greenstone and schist series. This series is similar in lithology to the Lower Proterozoic Udokan Series of the western part of the Aldan Shield. K-Ar whole-rock ages of amphibolites from a0-Mg0-Al 0 diagram for basic the troughs fall in the range 2600-2700m.y. The Fig. 5. Cmetavolcanics of various greenstone belts. determined age of pegmatites intruded into sandSymbols as defined in Figure 3. 2
3
2
3
GEOLOGY AND METALLOGENY OF ALDAN SHIELD
s m Fig. 6.
115
» o tzii« iz]» en,
Geological map of the central part of the Aldan Shield showing the distribution of phlogopite deposits (alter Shvetsov & Mekhanoshin, 1962): 1— Upper Proterozoic and Phanerozoic sedimentary cover; 2— Archaean Timpton-Dzheltula Series; 3—Archaean Yengra Series; 4—magnesian-carbonate and magnesian-silicate rocks of the Fedorov Suite; 5—Archaean and Early Proterozoic granites; 6—faults; 7—boundaries of phlogopite belt in the outer zone of the Central Aldan ring structure; 8—phlogopite deposits.
GRANITIC ROCKS The Archaean and Early Proterozoic granites of the Aldan Shield are subdivided into two groups, namely plagiogranite-charnockite and migmatite-granite (Bilibina et al., 1976). Plagiogranite, enderbite, hypersthene granite (char-
nockite) and their pegmatites comprise the first group. They occur mostly as intercalated but widely dispersed bodies. The migmatite-granite of the second group is represented by leuco- or biotite-microcline granite as well as by biotiteamphibole and diopside granite. Pegmatitic
116
V. I. KAZANSKY & V. M. MORALEV
leuco-granite and pink pegmatite also belong to this group. The granite and migmatite are confined to the ring structures of the gneiss-granulite basement, and are therefore mainly concentrated inside the Central-Aldan dome, whose inner zone is composed of the Upper Aldan Suite rocks and outer zone of the Fedorov Suite rocks. The distribution of the Archaean granite is also controlled by zones of high strain or m a j o r regional faults. These faults perhaps penetrated to deep zones where partial melting of the basement rocks occurred. Tightly compressed folds and belts of migmatite, as well as cataclastic rocks, blastomylonite and basic, intermediate or acid dyke swarms, are observed within these zones (Kazansky, 1978).
METALLOGENY Mineral deposits in the Archaean terrains of the Aldan Shield originated by processes of sedimentation, metamorphism, granitization and hydrothermal activity, with deposits of iron, phlogopite and crystalline quartz being the most important (Sidorenko, 1976). Skarn deposits of magnetite and phlogopite as well as hydrothermal veins with quartz crystals are associated with gneiss-granulite complexes, whereas iron-rich quartzite occurs mainly in the trough sequences. The skarn deposits of phlogopite and magnetite occur in diopside rock, dolomite marble and calciphyre of the Fedorov Suite in the outer zone of the Central Aldan ring structure (Fig. 6). According to the models of Korzhinsky (1953), these deposits are contact-metasomatic and have
0
m4|H5EI36
t
20p M
•ivCZH8
[ZD 9 Fig. 7.
Geological map of the Fedorov phlogopite deposit (after Ronnenson, 1975): 1-diopside-hornblende schists; 2—hypersthene-biotite-hornblende schists; 3-biotite-hornblende-pyroxene schists; 4-hypersthene-biotite schists; 5-magnesian-silicate (diopside) rocks; 6-granites and migmatite*; 7-phlo|opite lodes; 8—basic dykes; 9—faults.
GEOLOGY AND METALLOGENY OF ALDAN SHIELD 117 formed by progressive alteration of magnesiancarbonate rocks by granitic magma and hydrothermal solutions. The temperature of skarn formation in the early stages corresponded to the conditions of granulite-facies metamorphism and was about 800-700 °C, while in the later stages it approximated amphibolite-facies conditions and was about 700-550 °C. Country rocks are as old U \\\o o • H®/ as 3000-3500m.y., but the magnesian skarns have Rb-Sr ages of 2400-2200m.y. (Brandt et al., M o < ESP 9 I [f « 1978). Thus, phlogopite and magnetite deposits CZ> were formed during the final stages of evolution of the gneiss-granulite complex, when it was subjected to intensive granitization and superimposed metamorphism of the amphibolite facies EZk l ! f / I + +/> (Duk et al., 1975). The complicated structure of rv ^ the ore deposits results from the en echelon arrangement of several folded ore horizons and S3' zones of migmatization (Fig. 7). The Fedorov Suite in the central part of the Aldan Shield is 2300-3100 m in thickness and includes 10-15% of magnesian-carbonate and magnesian-silicate rocks. Beds and lenses of these rocks, together with the ore-bearing skarns, occur + + in four productive horizons of 100 to 500m thick. I ^ m m Within the productive horizons, the lens-like ore bodies inevitably occur in the vicinity of f Z 4hm / * magnesian-carbonate rocks with intercalations of gneiss and basic schist (Shvetsov & Mekhanoshin, 1962; Ronnenson, 1975). Lithology is also of great importance in localization of quartz crystal deposits, as they all occur in quartzites of the Upper Aldan Suite; i.e. in BU'X the lower part of the Early Archaean Yengra Series (Arkhipov & Frumkin, 1979). The thickness of the Aldan Suite amounts to 2500-2000 m position of banded iron formation with the dominant quartzites containing thin Fig. 8. inGeological the western part of the Aldan Shield (after intercalations of amphibolite, and high-alumina Glagolev et a/., 1977): 1—Cambrian sedimenand basic schists. The deposits are confined to tary cover; 2—Upper Proterozoic sedimentary small folds complicating the major structures. cover; 3^Borsala Series; 4—Olekma Series; 5—Archaean and Early Proterozoic granites Such folds are cut by numerous fault zones in and migmatites; 6—Jurassic-Cretaceous which crystal-bearing quartz veins and stock work diorite-monzonite intrusions; 7—banded iron zones are localized. They are considered to be formation; 8—BIF under sedimentary cover; generated in the Middle Proterozoic due to 9—geological boundaries; 10—faults; 11 — remobilization of silica by high-temperature shear zones; 12—iron deposits. metamorphogenous solutions. Open cavities with quartz Crystals occur within the veins and rarely The ore-bearing Borsala Series is about 3.5 km in enclosing recrystallized quartzites. thick. The lower part consists of intermediate Late Archaean ore deposits of iron-rich quartz- metavolcanics, high-alumina and mica schists ite are known from the western part of the Aldan and magnetite quartzite, whereas the upper part Shield near the Baikal-Amur railway, being local- comprises alumina and mica schists and metaized within a system of submeridional troughs andesite with intercalations of hematite quartzite. filled with volcano-sedimentary rocks of the Bor- Beds of iron-rich quartzite up to 80 m thick, altersala Series (Fig. 8). They represent a new, large, nating with quartz-muscovite, garnet-amphibole iron-ore basin in the eastern Soviet Union, with and garnet-staurolite schists and gneisses account reserves of iron ore in the Chara-Tokko region for up to 70-80% of the ore-bearing horizons of estimated at 8 billion tonnes (Arkhipov & Frum- the Borsala Series. Hornblende-magnetite quartzite is the dominant rock type. kin, 1979; Glagolev etaL, 1977). "
4
\\\
°
°
mim
7
i'tyy.+
+
+
+
+
118
Fig. 9.
V. I. KAZANSKY & V. M. MORALEV
Geological map (A) and cross-section (B) of the Imalyk iron deposit in the Chara-Tokko region (after Kudryavtsev et al., 1980): 1—Quaternary deposits; 2—Upper Proterozoic sedimentary cover; 3-6— Borsala Series (mica schists, quartzites, metavolcanics); 3—low grade (greenschist) domains, 4—mediumgrade (epidote-amphibolite) domains, 5—high grade (amphibolite) domains, 6—magnetite and hornblende-magnetite quartzites; 7—Jurassic-Cretaceous diorite-monzonite intrusions; 8—Archaean granites; 9—Archaean basic intrusions; 10—cataclasite and diaphthorite; 11—porphyroblastite.
GEOLOGY AND METALLOGENY OF ALDAN SHIELD
The structure of the Chara-Tokko region is very complicated because of its position in the marginal part of the Aldan Shield (Kudryavtsev etal., 1980). Several tectonic blocks are separated by large N-S and NW-SE faults, and the folded structures of iron-ore horizons within such blocks have not yet been interpreted (Fig. 9). Prograde metamorphism of the ferruginous quartzite and enclosing rocks corresponds to the middle and high grade of the amphibolite fades and markedly affects the metallurgical properties of the iron ore. Regressive dislocational metamorphism in fracture zones was accompanied by formation of chlorite- and talc-magnetite quartzites.
COMPARISON WITH OTHER TERRAINS Geological data available for the Aldan Shield and the Archaean cratons of India, Canada, Australia and South Africa reveal considerable similarity in Early Precambrian lithology and structure as well as some important regional distinctions. The Archaean volcano-sedimentary sequences of domes, synclinoria, trough structures and fault zones, high-grade metamorphism and intensive granitization shown within the Aldan Shield also characterize parts of the other old shields. Its most distinctive feature, however, is a considerable thickness of magnesian-
119
carbonate rock and quartzite which served as favourable horizons for development of skarn deposits of phlogopite, magnetite, and for distribution of hydrothermal quartz veins. The Archaean trough structures of the Aldan Shield have been developed within rigid microplates of continental crust, in trough (rift) basins. Their volcano-sedimentary sequences contain volcanic rocks of basaltic and komatiite chemistry but these trough structures differ from many such structures described from other continents in their relatively small size, predominantly terrigenous lithology, their large volume of ferruginous quartzite with subordinate sulphide mineralization, and the intensive metamorphism and granitization of the component supracrustal rocks. They more closely resemble the supracrustal rocks of high-grade gneiss terrains than those of most granite-greenstone terrains (e.g. Windley, 1977). ACKNOWLEDGMENTS This paper is a contribution to IGCP Projects No. 91 (Metallogeny of The Precambrian) and No. 92 (Archaean Geochemistry). We thank M. Z. Glukhovsky, A. A. Glagolev, M. I. Kuzmin and many other colleagues working in these projects for stimulating discussions on the problems relevant to tectonics, geochemistry and metallogeny of the Aldan Shield.
REFERENCES
ARKHIPOV, YU. V . , & FRUMKIN, I. M . (EDS), 1979: Geo-
logy of the USSR. XVIII. The Yakut ASSR. Mineral Resources. Nedra, Moscow (in Russian). BERYOZKIN, V . I., 1977: Metamorphism of Lower Proterozoic Complexes of the Aldan Shield. Nauka, Novosibirsk (in Russian).
BILIBINA, T . V . , DASHKOVA, A . D . , DONAKOV, V . I., - ERMOLAEV, B. A . , TITOV, V. K . , & BALASHOVA,
Z. N., 1976: Geological formations and metallogeny of the Aldan Shield. Tr. Vsesoyuzn. Nauchn. Issled. Geol. Inst. Min. Geo. SSSR, 276. Nedra, Leningrad (in Russian).
BRANDT, S. B . , KISELEV, YU. V . , PERMINOV, A . V . , GERASIMOV, N . S . , GODVINSKY, G . P . , EGOROVA, T . v., VOLKQVA, N . v., & PETROVA, Z . J . , 1978:
Rb-Sr and K-Ar absolute age determinations of the Aldan Shield schists. Soviet Geol., 9, 70-79 (in Russian).
DUK, V . L . , SALYE, M . E . , & BAIKOVA, V . S . , 1975:
Structural and Metamorphic Evolution of the Aldan Phlogopite-Bearing Granulites. Nauka, Leningrad (in Russian).
GLAGOLEV, A . A . , KISELEV, G . N . , KRAVCHENKO, V. M . , & BORONIKHIN, V . A., 1977: Mineral types
and.metamorphism of ferruginous quartzites from the Chara-Tokko region, the western part of the Aldan Shield. Geol. Ore Depos., 2, 51-62 (in Russian).
GLUKHOVSKY, M . Z . , & PAVLOVSKY, E . V . , 1973: T o t h e
problem of early stages in the evolution of the Earth. Geotectonics, 2, 3-7 (in Russian) [Geotectonics, 7(2), 61-63 (Eng. Transl.)].
GLUKHOVSKY, M . Z . , MORALEV, V . M . , & KUZMIN,
M. I., 1977: Tectonics and pedogenesis of the Katarchaean complex of the Aldanian shield relative to the proto-ophiolitic problem. Geotectonics, 6, 103-117 (in Russian). [Geotectonics, II (6) 467-475 (Eng. Transl.)]. KASTRYKINA, V . M . , & KARSAKOV, L . P., 1977: Petrochemical analysis and geochemical peculiarities of the basic schists of katazone of the Aldan shield. Bull. Moscow Soc. Naturalists Geol. Ser. 52(2), 5-21 (in Russian). KAZANSKY, V. I. (ED.), 1978: Endogenous Mineralization of Ancient Shields (Evolution, Structural and Petrological Conditions of Ore-formation). Nauka, Moscow (in Russian). KORZHINSKY, D . S., 1953: On metasomatic processes; in Betekhtin, A. G. (Ed.) Basic Problems in the Study of Magma tic Ore Deposits, 332-452. Publ. House Acad. Nauk SSSR, Moscow (in Russian).
KUDRYAVTSEV, V . A . , AKHMETOV, R . N . , TEREKHOVA, R. V., & KRYLOV, V. V., 1980: On the structure of
iron ore deposits belonging to the Imalyk group. Geol. Ore Depos., 2, 3-17 (in Russian).
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MORALEV, V. M., 1978: Characteristic features of metallogeny and tectonics of greenstone belts; in Bilibina, T. V., & Sokolov, Yu. M. (Eds) The Problem of Metallogeny of the Precambrian, 205-211. Nauka, Leningrad (in Russian). POLOVINKINA, Y u . IR. ( E D . ) , 1973: Geochronology
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RUDNIK, V . A . , & SOBOTOVICH, E . V . , 1969: O n t h e a g e
of polymetamorphic complexes of the Yengrian series in the Aldanian shield. Dokl. Akad. Nauk SSSR, 189(4), 834-837 (in Russian) [Dokl. Akad. Nauk SSSR, 189, 82-85 (Eng. Transl.)].
SHVETSOV, E . S . , & MEKHANOSHIN, A . P . , 1962: Regu-
larities of localization of phlogopite deposits in the Aldan mica-bearing region; in Petrov, V. P. (Ed.) Regularities of Distribution of Mineral Deposits. Publ. House Acad. Nauk SSSR, Moscow, 6, 373384 (in Russian). SIDORENKO, A. V. (ED.), 1976: Geology and Mineralogy of Old Shields Within the USSR. Nedra, Leningrad (in Russian). WINDLEY, B. F., 1977: The Evolving Continents. Wiley* New York. ZONENSHAIN, L . P . , KUZMIN, M . I . , & MORALEV, V. M.,
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and Metallo-
THE PRINCIPAL GEOLOGICAL AND GEOCHEMICAL CHARACTERISTICS OF THE ARCHAEAN GREENSTONE-GNEISS SEQUENCES IN NORTH CHINA Sun Dazhong & Wu Changhua Tianjin Institute of Geology and Mineral Resources; Chinese Academy of Geological Sciences
ABSTRACT North China was a mature portion of the "continental" crust from the early Archaean. It was technically stable until the later part of the era, when the rocks were intensely folded and consolidated to form a sialic basement for early Proterozoic volcano-sedimentary basins. The Archaean rocks are quite distinctive. No undoubted komatiite or anorthosite have been reported, and tonalite and trondhjemite are rare. Marble is intercalated in many places, so that parts of the sequence are similar to the Madras gneiss succession. Chemically, the Archaean rocks are potassium-rich, resembling neighbouring terrains in the USSR, but contrasting with the sodium-rich Archaean terrains of the Canadian Shield, Gondwana and the North Atlantic Craton. The crust formed in the Archaean of North China has profoundly affected the subsequent tectonic development of the region.
INTRODUCTION The Precambrian in China is customarily divided by Chinese geologists into two units, i.e. the Sinian and Presinian. The Presinian includes Archaean and Lower Proterozoic rocks and is exposed in the area between latitudes 30° and 45° North (Fig. 1). Between these latitudes the area to the east of longitude 105 °E is the best-known area of the North China Platform. The stratigraphic sequences and geological structures of the Archaean and Proterozoic in North China have been summarized in a series of papers by Cheng Yu-chi et al. (1973, 1979) and Ma Xingyuan etal. (1979). In addition to the principal geological characteristics of the Archaean of North China, the present paper deals with the chemical characteristics of the rocks, particularly the potassiumrich Archaean crust of the Yanshan and WutaiTaihang regions, and the sedimentary environments of the area. Chemical averages are calculated from 650 major-element analyses of Archaean and Proterozoic rocks. Except for the uppermost rocks of the Lower Proterozoic, which contain a large proportion of carbonate sediments, the Archaean and Lower Proterozoic of North China consist of four rock types: gneiss and migmatite (covering about 48% of the total area of the outcrop); granulite (4%); meta-volcanic and meta-sedimentary rocks (9%); and granite and granodiorite older than 2000 m.y. (39%). The Archaean is composed of wide-
Spec. Pubis geol. Soc. Aust., 7 (1981)
spread gneiss and migmatite and scattered granulite, forming a gneiss-granulite region. The early part of the Lower Proterozoic is represented by metamorphosed volcanic and sedimentary rocks, or greenstones, and these are overlain by welldifferentiated and slightly metamorphosed sedimentary rocks containing a great quantity of carbonate. PRESINIAN STRATIGRAPHY AND ITS CHRONOLOGICAL SUBDIVISION In North China the Wutai-Taihang Region and Eastern Yanshan are typified by well-developed and distinctive Archaean and lower Proterozoic sequences, each comprising four units. Ma Xingyuan et al. (1979), in the light of the Rb/Sr isochron age of 3670m.y. for the Qianxi Group (Geol. Inst. Academia Sinica, 1978), and the Pb/Pb isochron age of 3186m.y. for the Anshan Group, have divided the tectonic evolution of the Archaean and lower Proterozoic into five stages with the boundaries at 3500m.y., 3000m,y., 2500m.y., 2000m.y. and 1700m.y. Alternatively, Wang Hongzhen (1980) has divided them into the Qianxi, Dantazi, Wutai and Huto Groups, the upper limit of each being 3100m.y., 2600m.y., 2200m.y., and 1800m.y. As there are reservations about the age of 3670m.y. for the Qianxi Group, and the other dates older than 3000m.y. need to be confirmed, Cheng Yu-chi et al. (1979) considered 2500 ± 100m.y. (corresponding to the Fuping Movement) as the boun-
122
SUN DAZHONG & WU C H A N G H U A 110°
-1— 0
100 200
*
•
300 4 0 0
•
I
120°
500km
l
-4d
Huang
Hai
Archaean Archaean and lower Proterozoic ^ ^
Lower part of lower Proterozoic
'Nanjing Shanghai
Upper part of lower Proterozoic ^ ^
Fault
^ ^
Regional boundary
30b_
Fig. 1.
3d1 Dong Mai
110°
120*
Sketch map showing the distribution of the Archaean and lower Proterozoic of North China. I. Southern part of Northeastern China (Liaoning and Jilin) II. Eastern part of peninsula Shandong and southeastern Shandong III. Central and western Shandong IV. Huaiyang region V. Yinshan and Yanshan region VI. Taihangshan-Wutaishan-Luliangshan region VII. The northern slope of eastern Qinling Range.
dary between the Archaean and Proterozoic. Stratigraphic correlation indicates that the Archaean is represented by the Qianxi and Badaohe Groups in eastern Hebei, and the Lower Proterozoic by the Wutai and Huto Groups in the Taihang-Wutai Region. MAIN GEOLOGICAL CHARACTERISTICS OF THE ARCHAEAN Geological characteristics of the Archaean in North China are summarized below. (1) Archaean metamorphic rocks in North China are composed of amphibolite-facies rocks—biotite-plagioclase gneiss or amphiboleplagioclase gneiss, biotite granulitite* and plagio-
clase amphibolite. In some places, chiefly in the north of the area, there are also early Archaean granulite-facies rocks—biotite-pyroxene-plagioclase gneiss and pyroxene granulite. In other places there are intercalations of magnesium-rich marble, mica schist or sillimanite gneiss, metaarkose, banded iron formation, and graphite schist or graphite gneiss, which probably reflect differences in the conditions and environments of sedimentation. The lower part of the Lower Proterozoic consists mainly of low-grade, metamorphosed semi-argillaceous rocks, clastics and basic to intermediate to acid volcanics, with carbonate rocks or banded iron formation intercalated in places. Locally, there are sodium-rich, basic to acid, volcanic rocks.
of * S l C h ° r r d S ? the<;<peppra,nd:salt M o i n e " ^ quartz-feldspar-granulites of Scotland. In order ^ "granuhtes" rom the granulite-facies "granulite" of higher grade metamorphism, ture f e x T e f ' P ™ ™ ^ P r o P o s e d f o r the former and for other rocks showing similar structure, texture and degree of metamorphism (Cheng Yu-chi et al., 1973).
GREENSTONE-GNEISS SEQUENCES IN NORTH CHINA
(2) The Archaean and Lower Proterozoic rocks in North China were extensively folded, and the Archaean structural lines to the west and east of the Tan-Lu fault occupy one fold system in the shape of an arcuate structural line (Fig. 2). In the Archaean terrain, superimposition of two or more major folds can usually be recognized. For instance, there are two stages of folding (the early one with a nearly N-S trending axis and the later one with an E-W trending axis) in east Hebei, in the north of North China, and in the Songshan region in the south (Ma Xingyuan et al., 1979). In the lower Proterozoic terrain folding is also widespread, and in places it apparently follows Archaean trends. Judging from the appearance of the structures, the fold system in the basement of the North China Platform has had a strong influence on later tectonics. It also appears that the tectonic environment, and the conditions under which tectonic movement occurred,, are similar everywhere within the North China region. (3) The metamorphism of the Archaean and lower Proterozoic strata in North China was controiled by depth, the oldest strata occurring in the katazone. Amphibolite-facies rocks of the mesozone are most widespread. Granulite-facies metamorphism in the Yanshan-Yinshan region is more
123
extensive than in the southern part of North-East China, at Taihangshan, and on the northern slope of the Qinling Range, where it is only sporadic. The outcrops of the granulite-facies rocks may indicate deep erosion, or they may be the upwarped parts of folds. Greenschist-facies rocks are chiefly in the lower part of the Lower Proterozoic and are rare in the Upper Archaean. Thus the regional metamorphism within the basement of the North China Platform was relatively uniform, and is believed to have been controlled by a layered heat flow, with generally rising temperature, slightly different in different parts of the region. (4) The Archaean rocks in the basement of the North-China Platform are quite commonly migmatized. In most areas the migmatization is predominantly potassic, though it is somewhat sodic in the early stage. Migmatization was accompanied by regional metamorphism. Three stages of migmatization can be distinguished: 2500-2600 m.y., 2200-2300m.y., and 1800-2000 m.y; Migmatization of more than 2800m.y. in age has been reported from the southern part of northeastern China. In spite of the multistage formation of migmatites and migmatitic granites, no large intrusive bodies of tonalite and trondhjemite, or gregarious batholiths, are present.
Fig. 2. The Archaean structural trends in North China (adapted from Ma Xingyuan et al1979).
124
SUN D A Z H O N G & W U C H A N G H U A
Migmatites and migmatitic granites are prominent in the Archaean rocks of the region. (5) There was widespread deposition of banded iron formation during the Archaean Era in North China. The most favourable oreforming conditions for the banded iron formation apparently existed mainly in the northern and southern zones. The banded iron formation is mostly associated with volcanics but is typical of neither the Algoma nor the Superior type, and very often seems to be transitional. This suggests that there was no marked eugeosynclinal environment, such as that for the Algoma type, even in a rather active region. A summary of these geological characteristics indicates that the Archaean geological environment in Northern China was essentially uniform, and that the region probably represented a relatively, stable portion of the Earth's crust. In the late Archaean the strata were intensely folded and consolidated, giving rise to a vast basement of sialic crust, on which volcanosedimentary basins of early Proterozoic age were subsequently formed. CHEMICAL CHARACTERISTICS OF THE ARCHAEAN ROCKS General Lithology Within the Archaean metamorphic sequences of the North China region, the melanocratic rocks consist mainly of plagioclase amphibolite with rare hornblendite, whereas the leucocratic rocks are composed mainly of biotite-plagioclase gneiss and biotite granulitite, with leuco-granuli-
tite and amphibolite-plagioclase gneiss of secondary importance. In addition there is significant migmatite and granite. Mafic Rocks Amphibolites are common in the Archaean sequence. Most are artho-amphibolites of tholeiitic affinity but there are also some amphibolites of sedimentary origin. Similar rocks of early Proterozoic age are also found in North China. Archaean amphibolite from different areas normally has the following ranges: 6-7.5% MgO, 2.27-3.20% Na 0 and 0.77-0.94% K 0. The mean value of K 0 for the whole region is 0.89%, which is twice that of oceanic tholeiite and close to the average K 0 content (-1%) of continental tholeiite. Plots of compositions of Archaean and early Proterozoic amphibolite from the North China region are clustered both on the AFM diagram (Fig. 3a) and on the K 0Na 0 CaO diagram (Fig. 3b). Thus, throughout the entire Archaean Era, the mafic volcanic rocks show little variation in composition. Moreover, it can be seen on the Cr vs FeOVMgO diagram and the Ni vs FeOVMgO diagram (Fig. 4), which are based on data from the southern part of Northeastern China, the Yanshan Range, the northern slope of the Qinling Range, and the west-central part of Shandong, that most rocks plot in the field of basalts from stable continental and oceanic areas. This is particularly true for the west-central part of Shandong where basic volcanics of alternating marine and continental facies have been confirmed. Another characteristic feature of the diagrams is 2
2
2
2
2
2
FeO+Fe 03 2
Archaean ® amphibolite Proterozoic amphibolite ^ Archaean hornblendite
Alk
MgO Na 2 0
CaO
Fig. 3. Alk-F-M (a) and K 0 - N a 0 - C a 0 (b) plots for amphibolites and hornblendites from North China. 2
2
125
G R E E N S T O N E - G N E I S S S E Q U E N C E S IN N O R T H C H I N A TABLE I
Mean chemical composition of amphibolites and hornblendites from Archaean and Lower Proterozoic sequences in the various regions of North China. The regions are represented by Roman numerals whose position is shown in Figure 1. Amph i bo 1 i te
Hornblend l i t e Lower Proterozoic
Archaean V
Ti O2 AI2O3
VII
H3.sk 4 9 . 2 7 0.84 1.11 ]k.kk 14.03
Si O2 ,
4. 39
Fe 2 0 3
7.81
FeO MnO MgO
.
CaO
VI
III
IV
VI
IV
49-38
48.91
48.68
51.85
48.60
aver.
Archaean
49.90
VII
I
VI
47..16
50 .26
46.07
49.04
0 • 95 ' •• 1 .00
0.68
1.48
1.40
1.42
1.07
0,.56
0 .36
0.55
0.43
14 • 59
14.49
14.41
15.21
14.49
16.20
14.55
8.• 56
11 .40
10.13
10.58
3 .31
4.00
2.86
3.84
4.34
3.78
3.77
8,•72
3 .27
3-9-1-
4.63
8.46
1 1 . .79
3.65
9.05
8 .69
8.64
10.15
8.53
7.00
7.44
0.18
0.26
0 .18
0.17
0.21
0.19
0.15
0.19
7.50
6.46
7 .04
7-29
7.58
6.46
6.00
8.87
10.45
9,23 .
'
50..20
aver.
8 .68
9.31
5.33
7.46
9-31
9.58
9 .88
6.42
9.92
0.19
0.,16
0 .19
0.20
0.18
6.83
8,.36
1.1 ., 1 1
15.60
10.99
8.86
8..35
7 .12
12. 17
8.00
2.90
2.50
2 •73
2.56
2.72
2.70
2.27
3.20
2.59
1,,80
1,.77
1 .40
1.73
K20
0.94
0.92
0.94
0.94
0.77
0.88
0.84
0.62
0.89
1.•26
0 , .75
0.70
0.86
p2o5
0.17
0.17
0.18
0.20
0.08
0.22
0.13
0.36
0.17
0.. 1 4
0 .. 1 4
0.11
0.14
10
29
33
36
14
6
28
5
4
11
2
Na 2 0
Samples
_
the general overlap of the Cr and Ni contents of rocks from all four geographic areas, which further indicates the small compositional range of the Archaean basic volcanic rocks and the stability of crustal development in the whole North China region. Mean compositions of hornblendites from Archaean and Lower Proterozoic areas are given in Table I. Samples taken from the southern part of Northeastern China (I, Table I) differ significantly in composition from basaltic komatiite; samples from the northern slope of the Qinling Range (VII, Table I) have a markedly higher K 0 content; two samples of hornblendite from the Wutai-Taihang area (VI, Table I) also show
higher K2O, though the composition of the rocks is close to that of basaltic komatiite. However, no genetic relationships with komatiite are yet known. Consequently, in our discussion elsewhere of basaltic komatiite in the eastern part of Hebei province, we have been unable to present definite evidence for the existence of komatiite in the Archaean of North China, but propose that at least some of the hornblendite should be considered as Mg-rich basalt. Felsic Rocks Biotite-plagioclase gneiss and biotite-plagioclase granulitite are the most abundant leucocratic rocks in the Archaean of North China.
2
TABLE IIA
Mean chemical composition of gneisses and granulitites from Archaean and Lower Proterozoic sequences in various regions of North Ch ina. For locality of samples see Figure 1. The regions are indicated by Roman numerals. Granu 1 i t i te ( l e p t y n i t e )
B rot i t e - p 1 a g i i o c l a s e gne i ss | Lower Protero| zoii c
Archaean V Si o 2
60.57
T i 02
:•0.73
VII
I
VI
III
IV
IV
58.46
62 .94
66.36
61 .80
70.15
68..21
aver.
65.43
| Lower Protero| zoi c
Archaean V
VII
I
VI
III
IV
I
65.41
62.89
64. 93
56.75
63.45
70.82
65. 04
a ve r .
63.61
0 .48
0.34
0.94
0.46
0,.52
0.56
0.39
0.50
0. 44
0.66
0.47
0.32
0. 55
0.49
14 .54
16 .61
15.49
15 .70 '
13.93
14.,28
14.67
15.48
15.41
14. 89
16.40
16.06
14.66
14. 99
15.30
Fe 2 0 3
1,• 73
3 .29
1.05
3 .42
. 1.29
1.
1.90
1.78
2.02
1 . 65
3.35
1.72
0.75
2. 73
2.03
~
3.23
•
co vO
O.83
A1 2 0 3 ; 14.98 FeO
3.72
7 • 73
2 .81
2.91
3.• 72
2.50
2.•54
3.02
2.65
3.39
4. 21
4.41
4.28
1.67
1. 86
3.70
MnO
0.09
. 0,• 17
0 .09
0.07
0,. 1 2
0.08
0., 1 2
0.10
0.06
0.09
0. 09
0.12
0.10
0.04
0. 07
0.09
MgO
3.05
5.• 09
2 • 52
2.08
2,. 1 4
1.24
1. 64
2.31
1.90
2.12
2. 40
3.78
2.7?
0.57
2. 60
2.50
5..40
•2.74
2 . 42
CaO
3.85
2,.28
2.83
3.41
4.89
3.94
1.43
2. 01
3.02
3.95
2. •75
3..40
4.44
3.•32
3.98
2. 43 4. 56
3.43
Na 2 0
3-96
4.19
4.49
4. 04
3.80
4.03
4.00
4.87
4.20
K20
3.29
1. 86
2.,84
1.62
2. 28
2.67
2. 62
2.55
2.92
2.24
2 . 79
2.27
1.82
5.00
3. 14
2,61
P20s
0.23
0. 16
0.. 1 7
0.08
0. 93
0.11
0. 15
0.16
0.14
0.21
0. 09
0.28
0.08
0.08
0.,16
0.14
13
10
1
13
Samples
14
,
3.,80
6
2
10
1
19
>
7
3.12
8
23
51
SUN DAZHONG & WU C H A N G H U A
126
Fig. 4.
Cr vs FeO*/MgO and Ni vs FeOVMgO plots (c/. Miyashiro & Shido, 1975) of Yanshan Range (a) (b) and central Western Shandong, southern part of Northeastern China and the northern slope of Eastern Qinling Range (c) (d). 1 field of volcanic, rocks from island arcs and active continental margins; • 2 field of volcanic rocks from stable continental and oceanic regions; 3 field of abyssal tholeiites. TABLE
I IB
Mean chemical composition of gneisses and granulitites from Archaean and Lower Proterozoic in various regions of North China. For locality of samples see Figure 1. The regions are indicated by Roman numerals. i Leuco-•granul itite Archaean V Si 02 64.72 T i 0 2 0.37 A1203 15.52 Fe203 2,33 FeO 2.10 :0.12 MnO 0.66 MgO CaO 2.38 NazO 5-77 KjO 3.97 0.10 P2O5 Samples 41
VII I 70- 12 74.16 0.20 0.07 14.92 12.78 0.65 1.25 ,1.22 1.80 0.08 0.05 0.52 0.33 1.72 1.19 4.00 4.39 5-76 1.87 0.04 0.03 2,
8
Hornblende--plagioclase gneii ss
Lower Proterozoic VI 74.08 0.31 12.61 1.65 1.50 0.03 0.34 0.88 3.14 3.86 0.05
IV 82.05 0.12 8.77 0.39 1.32 0.04 0.29 0.57 2.61 3.36 0.03
9
2
I VI 71 .64 67-58 0.50 0.72 12.51 14.03 3.82 1.30 0.89 3.89 0.04 0.12 0.75 1.55 0.30 2.60 3-70 5.32 4.73 0.53 0.14 0.05 4
sequences
2
IV 76.71 0.15 11.52 1.81 0.62 0.34 0.51 0.50 3.79 4.38 0.03 3
aver. 68.84 0.34 14.25. 2.05 1.86 0.10 0.61 1.84 4.90 3.72 0.08
aver.
Archaean V V 44.21 55.46 0.67 2.35 10.01 16.52 9.31 3-17 1 1.30 4.62 0.26 0.17 4.28 7.47 10.86 6.62 2.12 4.13 1.56 0.59 0.32 0.33 2 18
6.74 7.77 3.05 1.58 0.11
IV I VI 56.24 55.71 59.58 54.78 0.66 0.58 0.85 0.76 15.54 16.46 1-5.35 15.38 1 .98 ; 3.64 2.97 3.39 5.3^. 5.23 - 5.71 4,71 0.14 0..14 0.12 0.19 3.25 4.63 2.73 3.97 6.54 5.72 5.70 6.90 | . 60 3.86 3.85 3.69 2.40 1 .50 •2.76 2.02 0.29 0.24 0.23 ,0.21
17
26
VII 49.63 0.78 14.51 5.O8
5.97 0.19
1
12 .
127
GREENSTONE-GNEISS SEQUENCES IN NORTH CHINA KoO
Fe0 + F e 2 0 3
Archaean H b - Plag - g n e i s s Archaean Bio - P l a g - g n e i s s Proterozoic Bio- Plag-gneiss
Alk
MgO
CaO
Na20
Fig. 5. Alk-F-M (a) and K 0-Na 0-Ca0 (b) plots for gneisses from North China. 2
2
Hornblende-plagioclase gneiss, migmatite and granite are less widespread, and there are minor areas of granulite and aluminous metasediments. The biotite-plagioclase gneiss and granulitite show compositional differences (Table IIA), but differ mainly in texture and fabric. Compared with average dacite (Nockolds, 1954), they show relatively low CaO and higher K 0. Compared with average greywacke (Pettijohn, 1949), they have a slightly lower MgO and CaO content and higher amounts of K 0 and Na 0. On AFM plots (Figs 5a, 6a) and K 0-Na 0-Ca0 diagrams (Figs 5b, 6b) the data for each type have a distinct scatter, which may reflect mixed origins for the rocks. Petrographic data suggest that these rocks are mainly of sedimentary origin and the range in 2
2
2
2
2
chemistry indicates a variety of sedimentary sources. In some leucogranulitites which retain original sedimentary features, clastic grains of potash feldspar, mainly microcline, are preserved. A high K 0 content characterizes these rocks and consequently most tend in composition toward the granite field, but they show marked variations in the Na 0/K 0 ratio. The composition of hornblende-plagioclase gneiss (Table I IB) lies between the amphibolites and the biotite-plagioclase gneiss (Figs 5a, b). The source material of the original sediments of this rock type probably included mafic rocks corresponding in composition to amphibolites, such as immature clastic sediments containing volcanic lava and tuff fragments. 2
2
2
FeO + Fe 03 2
Archaean migmatite and granite Proterozoic migmatite and granite O O
Alk
MgO
Na20
Ar. l e u c o - g r a n u l i t i t e Pt. l e u c o - g r a n u l i t i t e
9
Ar. granulitite
©
Pt. granulitite
CaO
Fig. 6. Alk-F-M (a) and K 0 - N a 0 - C a 0 (b) plots for migmatites, granulites and granulitites from North China. 2
2
128
SUN DAZHONG & WU CHANGHUA TABLE III
Mean chemical composition of migmatite-granitesfrom the Archaean and Lower Proterozoic of North China. See Figure 1 for location of samples, indicated by Roman numerals. M i gma ti it e - g r a n i te Lower P r o t e r o z o i c
Archaean
Si02
V
VII
I
VI
III
IV
I
IV
67.78
71.50
72.66
72.49
7 1 . • 59
73.71
70..88
74.42
aver.
71.73
Ti02
0.50
0.26
0.18
0. 16
0.. 1 9
0.18
0.38
0.20
0.25
A1203
14.64
14. 15
14.17
13-42
14.36
13.71
12. .92
12.78
13.89
1 .04
0..46
0.83
2..46
1.08
Fe203
2.41
1.48
2.15
1.45
FeO
3.29
0.55
1.41
1.23
1..69
1.01
2..31
1.22
1-73
MnO
0. 12
0.06
0.06
0.08
0.,04
0.03
0..08
0.08
0.07
MgO
1.72
0.61
0.52
0.65
0.•~71
0.35
0. 30
0.43
0.76
0.73
0.95
'
1 .68 .
1 .00
1 . 78
Na 2 0
3.35
3.42
3.77
3.49
4.. 1 9
3.85
4. 41
4.1.1
3.74
K20
4.45
4.36
4.36
4.19
4. 25
4.85
3. 35
4.18
4.29
P20s
0.20
0.06
0.05
0.10
0.. 1 0
0.03
0. 07
0.05
0.09
Samples
21
3
30
23
111
8
6
11
CaO
2.04
1.86
Migmatite and granite are significantly developed in Archaean rocks in North China and are commonly associated with biotite-plagioclase gneiss, biotite granulitite and leuco-granulitite. They coincide with the leuco-granulitite and part of the biotite granulitites on the AFM plot (Fig. 5a) and the K 0-Na 0-Ca0 plot (Fig. 5b) and are also characterized by high K 0 (Tables IIA, B). Although not all these granites and migmatites are Archaean, granite and migmatite with high K 0 content do exist in the more precisely dated Archaean rocks in the southern part of Northeastern China, the Yanshan Range, and the westcentral Shandong province (Table III). Granulite and aluminous metasediments are present only in minor amounts in the sequence, 2
2
2
2
"
0.74
and their compositions vary widely in different areas (Table IV, Fig. 7), suggesting a sedimentary origin for most of them. THE MEAN CHEMICAL COMPOSITION OF ARCHAEAN ROCKS FROM YANSHAN AND WUTAI-TAIHANG REGIONS From the foregoing discussion of the principal rock types it is obvious that the Archaean of North China is characterized by a high potassium content. The Archaean rocks of the eastern part of the Yanshan area and the Wutai-Taihang district are representative and have been chosen to calculate the mean composition of this segment of the crust (Table V).
Archaean ranulite
Alk
CaO MgO N a 0 Fig. 7. Alk-F-M (a) and K 0-Na 0-Ca0 (b) plots for granulitites and khondalites, North China. 2
2
2
1 • 19
G R E E N S T O N E - G N E I S S S E Q U E N C E S IN N O R T H C H I N A TABLE IV
Mean chemical composition of granulites and khondalites from the Archaean of North China. See Figure 1 for location of samples, shown by Roman numerals. Granuli te Archaean
Khondali te aver.
V
VII
VI
Si 0 2
55.90
69.67
49.08
60.67
Archaean
aver.
V
I
VI
60.81
70.84
79.13
65.38
T i O2
? 1.13
0.55
1.10
0.91
0.65
0.42
0.53
0.58
AI2O3
16.92
12.86
18.97
15.52
17.25
13.31
12.52
15.72
Fe 2 0 3
4.23
2.13
4.25
3.24
2.35
0.84
1.65
1.88
6,91
4.69
4.70
4.20
0.64
4.14
FeO
6.25
2.05
MnO
! 0.14
0.16
0.17
0.15
0.15
0.06
0.01
0.11
MgO
3-56
0.86
4.84
2.62
2.73
2.52
0.48
2.44
CaO
5.11
2.40
9.26
4.39
3.22
1.17
0.39
2.38
Na 2 0
3.25
2.67
3.65
3-06
3.05
2.16
0.25
K20
2.43
4.31
1.00
3.04
3.54
1 -73
4.08
P 2 0s
0.47
0.12
-
0.32
0.14
0.06
7
5
1
12
5
Samples !
129
basic eruptive rocks represented by spilite-keratophyre. The salient features of Archaean stratigraphy in other areas of North China appear to be intermediate between these two typical areas. It is quite possible, therefore, to estimate the mean chemical composition of all the Archaean rocks of North China on the basis of the average value of the Archaean rocks from the Yanshan and Wutai-Taihang areas. The result of the calculations (Tables V, VI) indicates that the Archaean of the Yanshan and Wutai-Taihang regions have a K 0 / N a 0 ratio of 0.98 and 0.86 respectively, which exceeds the K 0 / N a 0 ratio of Archaean rocks found in Canada (0.65), and approximates that of Archaean rocks of the basement of the Baltic, Ukrainian and Russian shields (0.85, 0.82, 0.85) (Ronov & Migdisov, 1970). The relationship can also be seen clearly from the K 0-Na 0-Ca0 diagram (Fig. 8), and from Figure 9. The Lower Proterozoic rocks of the Yanshan and Wutai areas (Table V, C, D) differ chemically from the Archaean rocks and from each other. The Wutai group has a low K 0 content and consequently a low K 0 / N a 0 ratio of 0.51, whereas the Yanshan rocks have a high K 0 / N a 0 ratio of 0.98, similar to the Archaean rocks, but a relatively low content of both K 0 and N a 0 . However, the Lower Proterozoic rocks have so small an area of exposure that they do not significantly affect the overall composition of the ancient crust of North China. 2
2
2.52 3.12 0.12
2
2
2
2
The Archaean rocks in eastern Yanshan (eastern Hebei) consist of two megacycles of basic volcanics, greywackes and other sedimentary rocks, with calc-alkaline volcanic extrusives in the overlying Lower Proterozoic. The Archaean in Taihang is composed of greywacke, intercalated with calci-magnesian sediments, or rhythmically layered sediments derived from basic volcanics. The Lower Proterozoic in the Wutai and Taihang areas contains sodium-rich
2
2
2
2
2
2
2
2
TABLE V
Average composition of Archaean crust and Lower Proterozoic rocks from China compared with other eratons
Samples SI0 2
9983 m
9139 m.
7866 m
90
95
92
43
251
18
171
180
65.28
65.71
62.28
64.15
64.81
62.25
65.72
65. 10
0.42
no2 AI2O3 Fe203
Ik. 56 1-71
3375 m
0.33
0.59
0.59
0.47
0.46
0.46
0.50
13-69
14.74
15.59
15.72
15.11
16.13
16.00
1.97
2.82
1.61
1.84
2.33
2.14
3.49
3.24
3.14
3.00
0.09
0.04
0.08
FeO
3.70
2.43
4.7.1
3-57
MnO
0.07
0.07
0.08
0.07 2.49
2.35
2.56
3.73
CaO
3.^8
3-56
3.84
3.26
Na 2 0
3-58
3.43
3.41
2.98
MgO
0.08 2.66
2.00
1.50
1.96
2.30
3 - 73
4.09
3-04
3.40
3,21
3.22
3.07
4.10
K20
3.45
2.95
1.75
2.93
2.75
2.65
0.15
0.12
0.19
0.14
0.06
2.60
P2O5
0.12
0.11
0.86
0.82
0.-85
0.66
0.349
0.333
K 2 0/Na 2 0 CaO CaO + A Ik K20 CaO + Alk Na?0 CaO+Alk A. B. C. D.
2.70 0.15
0.96
0.86
0.51
0.98
0.331
0.358
0.427
0.356
0.385
0.411
0.328
0.297
0.19^
0.319
0.284
0.266
0.299
0.265
0.341
0.345
0.379
0.325
0.331
0.323
0.352
0.402
Yanshan, Archaean (Qianxi and Badaohe groups of North China) Wutai-Taihang, Archaean (Fuping group) Wutai, Lower Proterozoic (Wutai group) Yanshan, Lower Proterozoic (Shuangshanzi group)
E. F.G. H.
Baltic Shield, Archaean Ukrainian, Archaean Russian platform basement, Archaean Canadian Shield, Archaean
SUN DAZHONG & WU CHANGHUA KoO
130
Baltic Shield: Archaean Ukrainian: Archaean Russian platform basement: Archaean Canadian Shield : Archaean Wutai-Taihang : Archaean in North China Yanshan : Archaean in North China Wutai : L.Proterozoic in North China Yanshan: L. Proterozoic in North China
CaO
Na 2 0
Fig. 8. K 0-Na 0-Ca0 by weight for Archaean regional average composition of Wutaishan-Taihangshan and Yanshan regions. The line separates all Canadian Archaean averages from all U.S.S.R. shield averages, plotted according to Lambert et al. (1976). 2
2
DISCUSSION In summary one can see that the salient features of Archaean rocks in North China are remarkably distinct from those found in many other parts of the world, especially in Gondwana and the western hemisphere. (1) So far, no undoubted komatiite has been found. Although some komatiites have been reported from southern North China (Li Jiliang et al., 1979) and northern North China (Zhang Yixia et al., 1980), the high grade of metamorphism and strong migmatization of the rocks
obscures their origins. Their chemical composition is unusual for komatiite, with K 0 generally being much higher than that of typical komatiite. At the very least we can say that komatiite is underdeveloped in the areas studied in detail, and that peridotitic komatiite is especially rare. Some rocks, however, approximate to basaltic komatiite in composition. (2) So far no Archaean anorthosite has been found in North China. Although the so-called Damiao anorthosite body in the Yanshan area yields an inferred age of at least 1.6b.y. (Zhong Fudao & Xie Guanghong, 1978), actual dating 2
TABLE V I
The distribution of Archaean and Lower Proterozoic rocks in North China Archaean
Lower Proterozoic
Reg ion
Yanshan
Wutai-Taihang
Yanshan
Group
Qianxi, Badaohe
Fupiing
Shuangshanzi
Th ickness
(m)
(*)
(m)
Wutai
(m)
Wutai (m)
<*)
]'
Hornblend i te
30
0.3
60
0.7
Amphi boli te
903
9.0
496
5.4
Hb-plag-gneiss
_
658
1667
_ _
16.7
825
9.0
-
-
Bi-plag-gnei ss
580
7.4
108
1.1
1566
17.1
-
580.
7.4
30
0.3
2070
26.3
Melano-granuli te
287
2.9
Leuco-granuli te
1455
14.6
-
Bi-granuli t i te
.
1550
_ _
_ _
15.5
753
8.2
16.12
47.8
50
0.5
1073
11.3
Leuco-granuli t i te Meta-basic volcanics Meta-intermediate-acid volcanics Migmat i te-grani te.
-
_ 3933
39-4
Clastic rock Khondalite-schist Marble Total
_
_
2.8
220.
8.4
_ _
2.4
186
553
16.4
750
450
13.3
1267
9.5
16.1
-
_
~
-
-
-
540
6.9
105
1.2
760
22.5
365
4.6
3870
42.3
-
-
397
4.4
-
-
9983
100.0
9139
99-9
3375
100.0
8.3
650
•
7866
-
100.1
GREENSTONE-GNEISS SEQUENCES IN NORTH CHINA
• 21
25~
20-
1-5-
z
\ O ¥
CM
•31
to0-5
l 8
•14
3-5
O 0 © ^
gives an age of l.Ob.y., indicating that the rocks are not Archaean. (3) Tonalite and trondhjemite are found only in small amounts. Both in the Yanshan and Wutai-Taihang areas the massive migmatite or older granite bodies are mostly calc-alkaline, with a higher ratio of K 0 / N a 0 than the Archaean in other parts of the world. (4) Except for its central part, the Archaean strata- in North China are intercalated with marble. The best development of Archaean marble is found in the Wutai-Taihang region, where the accumulated thickness ranges from 397-509 m, with the thickest individual bed reaching 200 m and extending over 30 km. This resembles the marble within the gneiss of Madras, in India (Naidu, 1963). The CaO/MgO ratio in some individual samples from pure marble reaches 55.86, an unusually high ratio for Archaean carbonates. (•5) The K 0 content of the metamorphic rocks of Archaean age in North China is relatively high. The amphibolite approximates to continental tholeiite in average K 0 content. In addition, there is potassium-rich leuco-granulitite of sedimentary origin in the Archaean strata. These differences, particularly the average composition of the crust in Archaean times, demonstrate that the Archaean of North China can be assigned to a potassium-rich portion of the Earth's crust, similar to that of the neighbouring districts of the Soviet Union, but obviously different from the sodium-rich crust within the Canadian Shield, Gondwana, and the North Atlantic Craton (Fig. 9). At the same time, the whole of North China had a relatively stable tectonic setting, and from an early age was a mature portion of the "continental crust". Although in the late Archaean, large-scale regional thermodynamic events in North China consolidated the Earth's crust, this strip of potash-rich 'continental' ' crust deeply influenced the subsequent evolution of the crust in the region. 2
|5 0^ 3 o e o10® c * q A •
20
30
131
West U.S. 7. South Africa 6. South Africa 5. West Australia 4. South Africa 3. South Africa 2. Minnesota Valley—U.S. 1. Greenland B. Fuping Group—Wutai—Taihang region, China A. Qianxi Group & Badaohe Group—Yanshan, China.
3-OH
2-5 1-5 Age in Aeons ( 1 0 9 years before present)
Igneous,meta-igneous and igneous-looking Sedimentary and metasedimentary Igneous .sedimentary, and metamorphic, undifferentiated Metamorphic igneous and sedimentary, undifferentiated
2
2
Fig. 9. Approximations of weighted average ratios of K 0 / N a 0 for major volumes of sedimentary, igneous and undifferentiated rocks of Archaean and Proterozoic age (up to 1500 Ma B.P.) of Australia, Africa, North America and North China (after Engel et al. 1974). 2
2
t
Lower Proterozoic
32. Martin formation—Can. 31. Lower Belt— Can. & U.S. .30. Lower Proterozoic—West North America 29. Svecofennide—Baltic Shield 28. Et—Then, Cameron Bay, Echo Bay Groups—Can. 27. Lower Proterozoic—Can. 26. West Australia 25. Belcher Group & Manitounuk Supergroup—Can. 24. Lower Proterozoic—Fennoscandia 23. South Africa 22. Whitewater Group—Can. 21. Huronian Supergroup—Can. 20. Otish Mountain Group—Can. 19. Great Slave Supergroup & Epworth Snare, Goulburn Groups—Can. 18. Hurwitz Group—Can. 17. Union Island Group—Can. D. Shangshanzi Group— Yanshan, China C. Wutai Group—Wutaishan, China.
Archaean
16. Red Lake—Lansdown House area—Can. 15. Yellow Knife area—Can. 14. Yellow Knife area—Can. 13. Red Lake—Lansdown House area—Can. 12. Superior Prov.—Can. 11. Superior Prov.—Can. 11. Superior Prov.— Can. 10. West Australia 9. South Africa 8.
2
4
ACKNOWLEDGMENTS The preparation of this paper was materially assisted by our colleagues Xu Wenzheng and Han Guang. J. D. Lewis, Geological Survey of Western Australia, constructively criticized the text.
SUN DAZHONG & WU CHANGHUA
132
REFERENCES
CHENG YU-CHI, CHUNG FU-DAO, & S u YUN-JUN, 1973:
The Pre-Sinian of Northern and Northeastern China. Acta Geologica Sinica, 1973 (1), 72-81 (English text printed in Beijing in 1972).
CHENG YU-CHI, BAI JIN, & SUN DAZHONG, 1979: T h e
Pre-Sinian of China. Papers Submitted to the Second All-China Stratigraphic Congress, 1-20.
ENGEL, A . E . J . , ITSON, S. P . , ENGEL, C . G . , STICKNEY, D . M . , & CRAY, E . J . , 1974: Crustal evolution and
global tectonics: a petrogenic view. Bull. geol. Soc.
Am.,
85, 8 4 3 - 8 5 8 .
LAMBERT, R . ST J . , CHAMBERLAIN, V . E . , & HOLLAND,
J. G., 1976: The geochemistry of Archaean rocks; in Windley B. F. (Ed.) The Early History of the Earth, 377-387. Wiley, London.
LI JILIANG, CONG BOLIN, & ZHANG WENHUA, 1979: A
preliminary study on the tectonic evolution of the Late Archaean Iron Formation in the SW of North China Fault Block. Scientia Sinica, 22 (12), 1430-1442.
M A XINGYUAN, W U ZHENGWEN, TAN YINGJIA, & HAO CHUNGRONG, 1979: Tectonics of the North China
Platform Basement. Acta Geologica Sinica, 1979 (4), 3 0 4 - 3 1 4 .
MIYASHIRO, A . , & SHIDO, F., 1975: Tholeiitic and calc-
alkalic series in relation to the behaviors of titanium, vanadium, chromium and nickel. Am. J
Sci., 275, 2 6 5 - 2 7 7 . NAIDU, P. R. J . , 1963: A layered complex in Sittam-
pundi, Madras State, India. Spec. Pap. min. Soc
Am., 1, 116-123. NOCKOLDS, S. R., 1954: Average chemical compositions
of some igneous rocks. Bull. geol. Soc. Am., 65,
1007-1032. PETTIJOHN, F. J., 1949: Sedimentary
New York.
Rocks. Harper,
RONOV, A. B., & MIGDISOV, A. A., 1970: Evolution of
the chemical composition of the rocks in the shield and sediment cover of the Russian and NorthAmerican Platforms. Geochem. Int., 1970 (4),
294-325. WANG HONGZHEN ( H . C. WANG), 1980: Megastages in
the tectonic development of Asia. Scientia Sinica,
23 (3), 3 3 1 - 3 4 5 . ZHANG YIXIA, YAN HONGQUAN, WANG KUANDE, & LI FUYUAN, 1980: An approach to komatiite in the
Qianxi group of Jidong (in Chinese), J. Changchun geol. Inst., 1980 (1), 1-8.
ZHONG FUDAO, & XIE GUANGHONG, 1978: T h e age of
anorthosite event and its geological implications. Geochimica, 1978 (3), 2 0 2 - 2 0 8 .
THE ATLANTIC GRANULITE BELT, BRAZIL Eberhard Wernick
Department of Mineralogy and Mineral Resources, Sao Paulo State University at Rio Claro, P.O. Box 178, 13500 Rio Claro, S.P., Brazil ABSTRACT The Atlantic Granulite Belt of Brazil extends in a broad zone along the Brazilian coast from the State of Bahia (about 10°S) to the State of Rio Grande do Sul (about 31 °S). The belt represents a fundamental Archaean zone of weakness and contains sedimentary and volcanic supracrustals and igneous infracrustals of Archaean age that have been affected by polymetamorphism and tectonic overprinting during the Proterozoic. It borders cratonic areas containing Archaean greenstone belts which were also affected by the Proterozoic events. In part, the Atlantic Granulite Belt clearly shows an ensialic and intraplate nature. The Atlantic Granulite Belt is of low- to medium-pressure type, with a geothermal gradient of 40-50 °C/km indicating a crustal thickness of 15-20 km. The main episode of granulite formation was at 2700 Ma, but both the granulite belt and adjacent cratonic areas have yielded ages of up to 3100 Ma. Geological, penological and geochronological data suggest a polycyclic/polyphase history for the granulites, which originally developed from crustal material. Large Archaean sialic crustal plates must thus have existed in Brazil at least 2700 Ma ago.
INTRODUCTION Archaean (older than 2500 Ma) granulite and charnockite occur in the oldest terrains of the South American Platform (Almeida, 1971), which includes much of Brazil. Available geological and geochronological data make it possible to discuss them in terms of several structural provinces. Archaean granulite and charnockite of this terrain occur as enclaves of up to several hundred k m in migmatite/gneiss terrains in Archaean cratons, as continuous belts surrounding cratons, and as lenses and irregular bodies in younger mobile belts where they clearly show structural overprinting and various degrees of reworking in post-Archaean times (Wernick, 1979; Wernick & Almeida, 1979; Wernick etal., 1979). In central and eastern Brazil, three granulite/ charnockite belts have been recognised: the 500km-long Goiano Belt trends N-S to NNE and is situated in the central Goiano Massif; the Atlantic Belt (Fyfe & Leonardos, 1973, 1974; Leonardos & Fyfe, 1974) runs roughly parallel to the coast of the country; and the Alfenas Belt is close to the boundary between the States of Sao Paulo and Minas Gerais, and crosses the State of Goias with a NW-trend. The Alfenas and Goiano Belts and the northern branch of the Atlantic Belt delimit a large Archaean craton, the so-called Paramirim Craton, which underwent partial fragmentation by Proterozoic mobile belts. The largest preserved part of this old craton is the Sao 2
Spec. Pubis geol. Soc. Aust., 7 (1981)
Francisco Craton (Fig. 1), which is surrounded by a Late Precambrian mobile belt. The Alfenas Belt and the southern part of the Atlantic Belt seem to delimit another large Archaean craton which is largely covered by the Parana Basin (Fyfe & Leonardos, 1974). The main purpose of this paper is to describe briefly the Atlantic Granulite Belt, to calculate Archaean geothermal conditions and crustal thicknesses, and to establish the history of such terrains in the Proterozoic. The subdivision of the Precambrian and the names of the main tectonothermal cycles in Brazil used here are given in Table I. THE ATLANTIC BELT IN THE SAO FRANCISCO CRATON The best-preserved portion of the Atlantic Granulite Belt is in the Sao Francisco Craton (State of Bahia). Here, the belt runs roughly from lat. 10°S to lat. 17°S where it is truncated by the Late Precambrian Araguai Belt, the SE limit of the Sao Francisco Craton (Fig. 1). The width of the granulite belt is highly variable, ranging up to 200 km. Although the general trend of the belt is N-S, there are deviations to the NE and NW, the most important being a large arch south of Salvador (13°30Y2-SJ. Also, near Salvador, the main belt bifurcates with a branch extending to the NE. The western contact between the granulite belt and the Archaean craton is poorly known, f
E. WERNICK
134
SL N
E F
AM
PHANEROZOIC SEDIMENTARY AND/OR VOLCANIC COVER _ L A T E PRECAMBRIAN MOBILE BELTS AND FOLDED REGIONS: P A - P A R A G U A Y - A R A G U A I A BR-BRASILIA AR~* ARApUAl' N E F R - N O R T H E A S T E N FOLDED REGION S E F R - S O U T H E A S T E R N FOLDED REGION. - L A T E PRECAMBRIAN COVER -
— L A T E PRECAMBRIAN HOLASSIC DEPOSITS. - O L D BASEMENT MORE OR LESS REMOBILIZED DURING THE PROTEROZOIC. _ U R U A p U A N O MOBILE BELT - E S P I N H A p O MOBILE BELT — CRATONIC BASEMENT. A M - A M A Z O N A S S L - S A O L U L Z S F - S A O FRANCISCO R P - R I O DE L A P L A T A . - . M A I N G R A N U L I T E AND C H A R N O C K I T E OCCURRENCES A T - A T L A N T I C BELT A L - A L F E N A S BELT G O - GOIANO BELT. PROVED ANO P 0 S 8 I B L E GREENSTONE BELTS.
Fig. 1.
The main Archaean granulite belts of central and eastern Brazil and Uruguay.
13.5 Iyer, 1976, 1979). High initial Sr&VSrM ratios for Primary Precambrian tectono-thermal cycles in Brazil such old rocks from Mutuipe (0.714 ± 0.009) suggests the existence of even older sialic crustal Brazilian Cycle Late Proterozoic 1000 - 570 Ma material (Cordani & Brito Neves, 1978). Rocks 1800 - 1000 Ma linked to the Transamazonian Cycle occur as M iddle Proterozoic Urua^uano Cycle Early Proterozoic Transamazonian Cycle 2500 - 1800 Ma plutons intruding the granulite belt, as belts of Jequie Cycle 3000 - 2500 Ma gneiss and schist of Transamazonian age, and as Late Archaean > 3000 Ma older rocks reworked during the Early ProteroGuriense Cycle Early. "Archaean zoic (Caraibas Complex). Migmatite present in but greenstone belts belonging to the craton pro- the granulite belt seems to be linked to these provide an indication of the westward extent of the cesses. A similar situation exists at the southern Atlantic Belt.:Toward its northern end within the border of the craton, in the region of the Late Sao Francisco Craton, the granulite belt becomes Precambrian Aracuai Belt, where Brazilian Cycle progressively smaller, wedging between the plutons coexist with metasediments and reworked greenstone belts of Serrinhas to the east and Archaean granulite. Jacobina to the west. This indicates the existence THE ATLANTIC BELT IN THE of at least two Archaean cratons (Fig. 1). The Atlantic Belt within the Sao Francisco NORTHEASTERN FOLDED REGION Craton comprises three complexes: Caraibas North of the Sao Francisco Craton in the (Barbosa, 1970), Jequie (Cordani, 1973) and Northeastern Folded Region (Fig. 1), the continuSalvador. These complexes are lithologically ous granulite belt becomes disrupted with the heterogeneous with anorthosite, ultramafic formation of discrete granulite blocks. The rocks, basic gneiss of calc-alkaline affinity, basic Northeastern Folded Region comprises several to acid charnockite, and a supracrustal sequence Late Precambrian mobile belts with a radial disincluding biotite-hornblende gneiss, quartzite, position between ESE and NE that converge to pelite, marble, and banded iron-formation. the W and are separated by polygonal median Sighinolfi (1971) has identified possible mag- massifs of Pre-Brazilian age, which were remobilmatic components in some of the granulite. Rock ized during the Late Precambrian. In almost all assemblages belonging to the amphibolite as well of these median massifs Archaean granulite and as to the hornblende and pyroxene granulite charnockite have been recorded (Brito Neves, fades of metamorphism, have been recognized. 1975). The size of the granulite occurrences, and Metamorphic P-T conditions have been fixed their mineralogy and texture, vary with the from various critical mineral assemblages. degree of remobilization during the TransStormer (1973), using the A1 0 content of ortho- amazonian and Brazilian Cycles. Some occur as pyroxenes from sapphirine-bearing phlogopite continuous homogeneous Archaean rocks but bronzitite, inferred conditions of 1000°C and more frequently they form granulite palaeosomes 5-8 kb. Aluminous metasediments from Salvador in migmatites with Early to Late Precambrian (Sighinolfi & Fujimore, 1978) suggest conditions neosomes. In both cases, the rocks show clear of 780-850°C and 4-7 kb. Medium pressure is also signs of polymetamorphism, tectonic overprintindicated by plagioclase + hypersthene ± clino- ing and strong mineralogical and textural changes pyroxene ± quartz ± garnet from granulite due to retrograde metamorphism, cataclasis, etc. which occurs through the whole belt (cf. Green & In the more acid rocks, mesoperthite is almost Ring wood, 1967). Garnets from aluminous meta- completely converted to microcline. Geochronosediments show compositional variation which logical data confirm the polymetamorphic Sighinolfi & Fujimore (1972) relate to local dif- history of the whole region (Brito Neves et al., ferential oxygen and water pressure during meta- 1974). morphism. Fluctuations in the P -total fluid As the granulite bodies have continuity within pressure have also been suggested by these a given median massif, and as the different massifs have been moved for hundreds of kiloauthors. Components within the belt yield ages that fall metres along transcurrent faults at the end of the into four groups: 3100 Ma (Guriense Cycle), 2700 Brazilian Cycle (Wernick et al., 1978), it is diffiMa (Jequie Cycle), around 2000 Ma (Trans- cult to prove any link with granulite of the noramazonian Cycle), and 600 Ma (Brazilian Cycle). thern part of the Atlantic Granulite Belt in the Rocks of the Guriense Cycle form larger areas Sao Francisco Craton. However, it is possible to (e.g. in the Mutuipe region) or _ are dispersed show structural continuity of the median massifs through areas belonging to the Jequie Cycle, sug- and the adjacent Sao Francisco Craton (Brito gesting that at least some of the 2700 Ma-old Neves, 1975). The granulite enclaves in the rocks were derived from older rocks (Cordani & Northeastern Folded Region are interpreted as A T L A N T I C GRANULATE BELT, BRAZIL
TABLE I
2
3
c02
E. WERNICK
136
40 / C9
/
RIO
^
/
/
o
PARA
/
I
<MARANHAQ t I /
t\
J \
l
<
K '
E
''
GROSSO
/
/ /
(
^
Mutuipi. (
Salvador
Jequie
Brasilia / * /
Go i a n i a
J L \
J
MINAS GERAIS
.ESPERITO SANTO
PARANA
A
V
/ - - . S A N T A '
• ^ Rio ' S a o Paulo
de
Janeiro
CATARINA
G R A N D E ^ ^ DO _ 30° S U L / 7 p ® r t o Alegre
V 1000 km
Fig. 2.
^
T-
ALAGOAS LAGC >*r X\A> ^ SERGIPE
B A I A \ }
R
y
Y"**"
'
I GOIAS /
A
^ JPABAIBA. ;PEVa,V<
PIAU i
'
C
»
^
\
MATO
GRANDE DO NORTE
/
Central and eastern Brazil, showing places referred to in the text.
ATLANTIC GRANULITE BELT, BRAZIL
the northernmost part of the Atlantic Granulite Belt that were preserved as remnants in a crust highly reworked during the Proterozoic. This is deduced from the similarity between preserved granulitic rock assemblages in some of the median massifs and those from the Sao Francisco Craton, and the probable ensialic evolution in situ of the Northeastern Folded Region (Brito Neves, 1975; Almeida et al.; 1976; Hasui et al., 1978; Wernick et al., 1978).
THE ATLANTIC BELT SOUTH OF THE SAO FRANCISCO CRATON South of the Sao Francisco Craton, in the States of Espirito Santo, Rio de Janeiro and Minas Gerais, is the Southeastern Folded Region (Hasui et al., 1975; Wernick et al., 1978). Here, Atlantic Belt granulites occur as isolated areas or palaeosomes in migmatite of Early or Late Precambrian age, and as more or less continuous gneiss-granulite-charnockite belts (the most important being the Paraiba Group). West of this belt is the Barbacena Group, a 2800 Ma migmatite/gneiss complex surrounding the Rio das Velhas Greenstone Belt (Wernick, 1979). In the Paraiba Group, it is possible to distinguish several rock assemblages, including a charnockitic suite (charnockite, enderbite, opdalite, jutonite, birkremite and norite), a cataclastic granulite suite, augen gneiss, a metasedimentary association (pelitic gneiss, quartzite, amphibolite, banded iron-formation, calc-silicate rocks and marble), and various migmatites. These lithologies either form distinct units or are intimately intermixed. Based on critical mineral parageneses, conditions of recrystallization of about 680-725 °C and 3.5-6.0kb have been suggested for rocks from an area in the State of Rio de Janeiro by Costa & Marchetto (1978). Leonardos & Fyfe (1974) fixed the conditions at about 700 °C and less than 4.0 kb with a geothermal gradient of 60 °C/km in a nearby area. However, data from the adjacent Alfenas Belt at the boundary of the States of Sao Paulo and Minas Gerais gave P-T conditions of 740-770°C and 5.9-6.3 kb based on garnetcordierite-biotite-sillimanite-spinel granulite (Oliveira & Ruberti, 1979) and 680-760°C and 5.36.6 kb based on wollastonite-bearing calc-silicate rocks (Oliveira & Alves, 1976). Data based on the A1 2 0 3 content of coexisting clino- and orthopyroxenes from a charnockite suite of Madras composition in this area gave P-T conditions of 795-883 °C and < 8/0kb (Oliveira & Hypolito, 1978). Ages of 2700-3100 Ma have been recorded in this area (Hama et al., 1979).
137
THE SOUTHERN BRANCH OF THE ATLANTIC BELT The southern branch of the Atlantic Granulite Belt forms the larger granulite-charnockite bodies in the States of Sao Paulo, Parana, Santa Catarina and Rio Grande do Sul, in the domain of the Southeastern Folded Region. This geotectonic unit is defined by three Late Precambrian mobile belts approximately parallel to the coast and separated by median massifs of Pre-Brazilian age. Contacts between the mobile belt and the median massifs are marked by transcurrent faults (Hasui et at., 1975; Wernick et al., 1978; Wernick & Penalva, 1978). In the State df Sao Paulo, there are granulitic and charnockitic rocks in the Piacaguera Complex bordering the Joenville Median Massif (Ridge, 1974; Sadowski, 1974; Hasui & Sadowski, 1976). Farther south in the State of Parana there are extensive dioritic to granodioritic charnockites near Serra Negra (Basumallick et al., 1969). They are associated with metasedimentary rocks, including gneiss, calcsilicate rocks and acid to basic granulite. Nearby, at Antonina, is a possible greenstone belt (Kaul et al., 1979) which contains magnesian schist, amphibolite, quartzite and banded iron-formation (Cordani & Girardi, 1967; Girardi & Santini, 1973). In the State of Parana another charnockitic complex around Pien has been described by Girardi (1976). It is composed of metaperidotite, pyroxenite, norite, gabbro and diorite associated with serpentinite, magnesian schist and intermediate to basic granulite. Some of these rocks underwent reworking during the Early Proterozoic, and occur now as palaeosomes in Transamazonian migmatite. Metamorphic recrystallization conditions are 750-880 °C at a minimum pressure of 7.0kb, based on a two-pyroxene geothermometer applied to associated mafic and ultramafic rocks. The coexistence of sapphirine, bronzite and spinel indicates conditions of about 800 °C and more than 4.0 kb (Girardi & Ulbrich, 1978). Temperatures of around 1000 °C inferred from the A1 2 0 3 content (7%) of orthopyroxenes, like those in similar rocks from the State of Bahia (Stormer, 1973) are considered unrealistically high. A Pien-like charnockite-granulite suite, still in the Joinville Median Massif but in the State of Santa Catarina, occurs at Barra Velha. There, ages around 3400 Ma for the formation of the ultramafic-mafic magmatic suite and 2800 Ma for the regional granulitization are reported by Minioli (1972). The Barra Velha Complex is within an extensive granulite area, the so-called Santa Catarina Granulite Belt, which extends southward to the Blumenau lineament, where it
E. WERNICK
138
meets Late Precambrian rocks of the Tijucas granitic layer encircling the Earth and being successively reworked. Problems linked to the initial Mobile Belt. The Santa Catarina Belt is composed of meta- Sr /Sr ratio of reworked granulites can be sedimentary and meta-igneous rocks, the former overcome by a differential Rb/Sr ratio distribuincluding quartz-feldspar gneiss, calc-silicate tion through the crust. Therefore, the reworking rocks, quartzite and banded iron-formation. of deep granulite crust would not necessarily Marble is virtually absent. Ultramafic rocks like result in high initial ratios. Two hypotheses have been put forward for the those from Barra Velha and Pien are scattered through the whole belt. Geochronological data origin of Archaean granulites. The first considers show the main granulite-forming event occurred them to be the refractory residue of intensive at 2700Ma and two remobilization episodes anatexis, implying vertical metamorphic zoning occurred during the Early and Late Proterozoic (Fyfe & Leonardos, 1973, 1974; Leonardos & (Hartmann et al., 1979). The Santa Catarina Belt Fyfe, 1974; Glikson, 1976). White & Chappell is correlated with the granulite-charnockite (1977) added to this model, suggesting that durbodies in the western part of the State of Rio ing ultrametamorphism and melting, granitic Grande do Sul, geotectonically belonging to the melts ascend to shallower levels, taking with them Rio de La Plata Craton (Hartmann & Nardi, much of the refractory residue because of the 1980). high viscosity and inadequate separation of the The best-known granulitic body there is that of melt. This possibly led to reworked, migmatite/ Santa Maria Chico (Nardi & Hartmann, 1979). It granulite complexes such as the Ubendian Mobile comprises NW-striking, massive banded rocks in- Belt (Kroner, 1979). The second hypothesis postulates that granucluding basic gneiss, quartzo-feldspathic gneiss, anorthosite, ultramafics, metapelite and marble. lites are the result of the granulitization of preHypersthene is widespread in several lithologies. existing material (Robertson, 1968; James, 1975; The belt, remobilized during the Late Pre- Williams, 1977). The occurrence of granulite cambrian, strongly resembles the African belts adjacent to low-grade Archaean terrains Mozambique Belt, but the rocks are rather poor must then be explained. Bell (1971), Mathur in K 0 when compared with granulites from (1973), Glikson & Lambert (1973) and Pichaother continents. Metamorphic conditions in the muthu (1967) considered this to be the result of Santa Catarina Belt are around 700-800 °C and erosion of tilted Archaean blocks. The undation 5-10 kb. hypothesis postulates that the granulite belts were not produced during distinct orogenic periods, DISCUSSION but represent segments of the lower continental The key importance of granulite belts to under- crust brought to the surface through internal standing the evolution of our Earth has long been rotation of rigid blocks in large .continents recognized. On the basis of world-wide occur- and/or large-scale over- and underthrusting rences of 2500-3000 Ma-old granulitic rocks in along linear zones (Kroner, 1979). continental areas, Moorbath (1975a) postulated To explain the apparently ensialic nature of that continental crusts were already in existence. granulite belts, Sutton (1977) and Watson (1976) He proposed a large-scale production of dry and invoke intraeontinental movement between barren granulite around 2500-3000 Ma to a depth crustal blocks along linear zones of high strain of 15-25 km, implying that ancient continental from stresses which build up during movement of crustal thicknesses would have been comparable large supercontinent-carrying plates (Piper, with modern ones. This large-scale granulite- 1976). Williams (1977) finds the occurrence of forming process provided the incompatible granulite belts linked to strong shear zones with elements and water which migrated through the high heat flux which result from primary conveccrust, resulting in a petrologically and chemically tion cells, whereas adjacent greenstone belts are layered crust. Based on initial Sr /Sr86 ratios of associated with rift zones generated by secondary Precambrian granulites, Lewis & Spooner (1973) convection cells. In this case, both the granulite suggested a major depletion of lithophile belts and greenstones are younger than migmaelements in the early stages of crustal evolution terrains. with no later repeated extraction of Rb and K on tite/gneiss Available data on the Atlantic Granulite Belt a regional scale. Expanding this idea, and using indicate the following: Sr /Sr growth lines, the addition of mantle- 1. The Atlantic Belt is a geotectonic unit of conderived material by vertical accretion would sugextent which, in Brazil, can be followed gest that large-scale reworking of early sialic crust tinental more or less continuously from 10°S to 31 °S and did not take place. Alternatively, Hargraves may be (1976) developed a picture of a relatively thin Uruguay. continuous with the granulites in 87
2
87
87
86
86
ATLANTIC GRANULITE BELT, BRAZIL
139 cessively reworked in later periods. The initial Sr87/Sr86 ratios for the granulite of Bahia exhibit a range of values from 0.703 to 0.711. Cordani & Iyer (1979) employing Sr^VSr growth lines for the regions of Mutuipe and Itabuna found age values of precursors which are not different from the age values obtained for the main radiometric events (Guriense and Jequie Cycle). However, these interpretations rely on the assumption that the analysed rocks bear a direct geochemical relationship to the upper mantle, which may not be true for granulites of sedimentary origin. 6. The Atlantic Belt borders 2700-3100 Ma-old cratonic granite-greenstone-migmatite areas along its length. Relations in the northeastern and southeastern folded regions are not clear because of the extensive reworking of the belt in these geotectonic provinces. The• northern part of the belt in the Sao Francisco Craton has a clear intraplate setting, a similar situation to that of the Paraiba Group in the States of Rio de Janeiro and Minas Gerais. Relations between the granulite and greenstone from the cratonic areas can be best explained in the State of Bahia by the model of Katz (1976). The large Archaean cratons delimited by the granulite belt underwent successive reduction by reworking and fragmentation during the Proterozoic, so that they occur either as larger areas surrounded by Proterozoic belts or as fragments between them (Wernick eta I., 1979). 7. Archaean granulite belts of Brazil have NW, N-S or NE trends and impose a pronounced structural control on post-Archaean mobile belts. The granulite belts are also intruded by Proterozoic granitoids, syenites and alkaline rocks. They TABLE II Metamorphic conditions in the Atlantic Granulite Belt show successive tectonic reactivation which controls subsequent Proterozoic and Phanerozoic vulcanism and sedimentary basins. This agrees with the concept of Katz (1976) that Archaean granulite belts represent, in part, the expression of fundamental fractures and lineaments develin a protocrust. Such zones of weakness ko oped may reflect mantle influence by large and ductile MS zones, which control the position of the transform faults to which the greenstone belts are k linked through transverse rifts. These zones have remained unstable in post-Archaean times, acting 3 ho as zones of tectonic instability or as preferential heat flux channels. ho 8. The extent of the Atlantic Granulite Belt, in some places with a clear interplate position and an ensialic nature, points to the existence of very large Archaean crustal areas in Brazil prior to 3100 Ma ago. 9. The occurrence of a broad granulite zone ko along the Brazilian Atlantic coast is the combination of two distinct processes. One is the production of granulites along mobile zones, which
2. The belt is homogeneous overall, and contains a supracrustal volcano-sedimentary suite and an infracrustal charnockite suite. The two suites may form distinct areas or belts, or may be intimately intermixed. 3. The rocks of the granulite belt belong to the amphibolite, hornblende granulite and pyroxene granulite facies, and display no clear metamorphic zoning. In general there are marked metamorphic contrasts, even in small areas, with intermixing or interlayering of rocks belonging to the amphibolite and granulite facies. 4. Published PT conditions permit the classification of the belt as a low- to medium-pressure type. Different metamorphic recrystallization conditions along the length of the belt indicate a geothermal gradient of about 40 ± 5 °C/km (Table II). As the overall equilibrium temperature is about 750-800 °C, this indicates a crustal thickness of 15-20 km, agreeing with values proposed by Moorbath (1975&). The geothermal gradient of the Atlantic belt is significantly higher than that obtained by Clifford (1973) from a study of African granulites (25 ± 5 °C/km). 5. Geochronological data show that the main development of the Atlantic Belt occurred 2700 Ma ago (Jequie Cycle), but the existence of many 3500Ma ages both in the belt and in adjacent greenstone belts indicates the existence of older sialic crustal material. In several parts of the belt, reworked granulites occur in Early to Late Precambrian mobile belts. This agrees with the ideas of Hargraves of an old sialic primary crust, suc-
PRES-
a i itu no (o,\ AUTHORS)
Stormer
(19:73.)
Sighinolfi £ Fuj imore ( 1 9 7 2 )
.
INVESTIGATED ROCKS
TEMPERATURE (°C> .
Sapphirinebear1ng bronz i t i t e s
- about 1000
G a - C o - S i 1 1 - B j -Sp gne i s s e s .
about
(Kb)
GE0 " THERMAL .GRADIENT (°C/km)
5 - 8
750
about 6
- 725
3.5-6. 0
Moutinho da C o s t a £ Marchetto (1978)
Several su i t e s
Leonardos £ F y f e (197*0
G a - C o - S i l 1 - B i -Sp gneisses Ca 1 c - s i 1 i c a t e rocks
01iverra £ Ruberti • (1979)-
G a - C o - S i 1 1 - B i -Sp gnei sses
7^0
- 770;'
5.9-6.
01iveira & A Ives (1976)
Ca1c-s i1icate rocks
680
- 760
5.3-6. 6
rock
680
a b o u t 700
be low
01 i v e i r a £ Hypoli to ( 1 9 7 8 )
Charnockites
785
- 725
bel ow 8
Girardi £ Ulbrich .(1978)
'Sapph i r i n e bear ing bronz i t i t e s Charnocki tes
750
- 800
above 7
Nardi £ Hartmann (1979)
Several su i t e s
700
- 800
5.0-10
rock
50
60
40
86
140
E.
WERNICK
represent a "lateral" increase in metamorphism when compared with adjacent Archaean greenstone-bearing cratons. The other is the result of the westward tilt and erosion of the South American Platform (Wernick, 1979). In the second case, the granulites represent the lower part of both the granulite belts and cratonic areas, indicating a "vertical" increase in metamorphism. In the first case, the granulite occurs as distinct linear belts and in the second as irregular areas in cratonic terrains. Studies to characterize the differences between these types of granulites are recommended. 10. The origin of the granulite from the Atlantic Belt is still under discussion. Fyfe & Leonardos (1973), Leonardos & Fyfe (1974) and Oliveira (1973) considered it as residue of intensive anatexis. However, their conclusions are derived from highly migmatized parts of the belt. Several detailed studies (Choudhuri et al., 1978; Pedreira e r a / . , 1978) show clearly that part of the belt results from the granulitization of earlier crustal
material, in part having migmatitic structures. Costa & Marchetto (1978) developed, for the Paraiba Group, a complex polyphase/polycyclic evolution scheme involving first the migmatization of early, as yet unidentified, material which then underwent amphibolite-granulite facies metamorphism. Still later, the rocks suffered metamorphism in the Early Proterozoic. Cordani & Iyer (1976) suggest that rocks belonging to the Jequie Complex result from the granulitization of older crustal material. ACKNOWLEDGMENTS The author thanks Professor Dr Alberto Pio Fiori from the Department of Mineralogy and Mineral Resources—UNESP, for stimulating discussion, Eugenio Dezem and Amalia P. da Silva for the help during compilation work, and is indebted to Dr Hallberg, C.S.I.R.O. Division of Mineralogy, Floreat Park, Western Australia, for a critical review of the manuscript.
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1976: The Upper Precambrian of South America.
Bol. Inst. Geoc. Univ. Sao Paulo, 7, 45-80. BARBOSA, O., 1970: Geologia economica de parte da regiao de Medio Sao Francisco, Nordeste do Brasil. Dep. Nac. Min. Div. Fom. Prod. Min.,
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CORDANI, U . G . , & BRITO NEVES, B . B . , 1978: G e o -
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1978: Charnockitic gneisses and granulites of the Botelhos Region, Southern Minas Gerais. An. 30th Congr. Bras. Geol., 3, 1236-1249. CLIFFORD, T. N., 1973: African granulites and related rocks: a preliminary note; in Lister, L. A. (Ed.) Symposium on Granites, Gneisses and Related Rocks. Spec. Pubis geol. Soc. S. Afr., 3, 17-23.
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Leopoldensis, 3 (40), MAMA, M . , ALGARTE, J . P . , KAEFER, L . Q . , & ARTHUR, 45-75. A. €., 1979: Idades Rb/Sr e K/Ar na regiao Sul Sudoeste de Minas Gerais e Leste do Estado de Sao OLIVEIRA, M . A . F . , 1973: Petrologia das rochas metaPaulo. 11° Simp. Reg. Geol., Soc. Bras. Geol. SP, morficas da regiao de Sao Jose do Rio Pardo, SP. (Abstr.), 4. Rev. Bras. Geoc., 3 (4), 257-278. HARGRAVES, R. B., 1976: Precambrian geologic history. OLIVEIRA, M . A . F . , & ALVES, F . R., 1976: Wollastonita Science, 193 (4251), 363-371. em associates calcicas de facies granulitica Caconde, SP. Rev. Bras. Geoc., 8 (4), 249-261. HARTMANN, L. A . , DA SUVA, L. C . , & FILHO, V. O . , 1979: Complexo Granulitico Catarina-Descri^ao e OLIVEIRA, M . A . F., & HYPOLITO, R., 1978: Ortopiroxenios e clinopiroxenios coexistentes nos implicates geneticas. Acta Geol. Leopoldensis, 3 granulitos de Sao Jose do Rio Pardo, SP. Rev. (10), 93-112. Bras. Geoc., 8 (4), 249-261. HARTMANN, L. A., & NARDI, L. V. 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As regioes de dobramento Nordeste e Sudeste. An.
30th Congr. Bras. Geol., 6, 2493-2507. WERNICK, E., HERVE, F. A . , JARAMILLO, J. M., & CAMINOS^E., 1979: America do Sul: urn exemplo
de predominio de regenera^ao da crosta sialica 161-181. sobre acregao lateral. An. 2nd Congr. Geol. Chileno, B, 47-62. WATSON, J. V., 1976: Vertical movements in Proterozoic structural provinces. Phil. Trans. R. Soc. WERNICK, E . , & PENALVA, F,, 1978: Contribuigao ao Lond., A280, 629-640. conhecimento das rochas granitoides do sul Brasil. Rev. Bras. Geoci8 (2), 113-133. WERNICK, E., 1979: Arqueano no Brasil: tentativa de uma sintese. An. Acad, brasil. Cienc., 5 (2), WHITE, A . J. R., & CHAPPELL, B. W . , 1977: Ultrametamorphism and granitoid-genesis. Tectonophys., 287-310. 43, 7-22. WERNICK, E., & ALMEIDA, F. F. M . , 1979: T h e geo- WILLIAMS, H . R., 1977: African Archaean mobile belts tectonic environments of Early Precambrian granuand granite-greenstone terrain. Nature, Lond., lites in Brazil. Precamb. Res., 8, 1-17. 266, 163-164.
GEOCHRONOLOGY
THE OLDER PRECAMBRIAN GEOCHRONOLOGY OF WESTERN AUSTRALIA J. R. de Laeter, W. G. Libby, & A. F, Trendall 1
1
2
2
Department of Physics, Western Australian Institute of Technology, Hayman Road, Bentley, Western Australia 6102 Geological Survey of Western Australia, Mineral House, 66 Adelaide Terrace, Perth, Western Australia 6000 2
ABSTRACT This paper reviews the extensive geochronological data on the two major Archaean blocks of Western Australia, concentrating on results which have become available since the 1970 Archaean Symposium in Perth; these are summarised in tables and maps. The great bulk of data (both before and after 1970) come from the Rb-Sr method. Considering the amount of information (e.g. 837 whole-rock Rb-Sr analyses) disappointingly few firm conclusions can be drawn. On Rb-Sr evidence alone the new data confirm the suggestion of Arriens (1971) that the granitoids of the Pilbara Block are significantly (possibly about 0.5 b.y.) older than those of the Yilgarn Block, whose origin spans a comparatively narrow time span centred at 2.6b.y. The little evidence so far available, from the Pilbara Block, suggests that the greenstone belts are only slightly older than the sialic rocks. The available U-Pb evidence generally supports these conclusions, although recent zircon ages show that work of this kind has great potential for detailed elucidation of early local crustal development.
INTRODUCTION Isotope geochronology in Australia was initiated in Dr P. M. Jeffery's laboratory in the Physics Department of the University of Western Australia, in the mid-1950s. The accumulated data subsequently contributed by this and other laboratories are now substantial. A historical review of Precambrian geochronology in Western Australia has been given by de Laeter & Trendall (1979), without analysis of the geological significance of the data. Wilson et al. (1960) critically reviewed the data then available on the Australian Precambrian, and Compston & Arriens (1968) repeated the task, incorporating information added since 1960. No later all-embracing review has been made, although Arriens (1971) presented an assessment of the Archaean geochronology of Western Australia, together with an impressive bulk of new data, at the Perth Archaean Symposium in 1970. Trendall & de Laeter (1978) later compiled a complete bibliography of Western Australian Precambrian geochronology, but made no attempt to evaluate the data. The primary purpose of this paper is to compile and display the geochronological results which have become available from the older Precambrian rocks of Western Australia since the review of Arriens (1971). Major earlier contribu-
Spec. Pubis geol. Soc. Aust.,7 (1981)
tions, such as Arriens's own, are also included. Our review is coupled with a commentary on the overall significance of the newly available data for such broad issues as the development of the early sialic crust, the applicability of uniformitarian concepts to Archaean geology, and the definition of the "Archaean". GEOLOGICAL FRAMEWORK, AND FORMAT OF DATA PRESENTATION Western Australia has two major Archaean blocks (Fig. 1). The geology of the smaller, northern, Pilbara Block is described in this volume by Hickman, and in greater detail elsewhere (Hickman, in press). The geology of the larger, southern, Yilgarn Block is summarized by Gee et al. (1981). All of the Pilbara, and much of the Yilgarn Block consists of classical "granitegreenstone" terrain. However, Gee et al. (1981) have modified an earlier tectonic subdivision of the Yilgarn Block (Geological Survey of Western Australia, 1975) and have proposed a four-fold subdivision into three granite-greenstone provinces and a western gneiss terrain. As this is the first geological synthesis of the Yilgarn Block to appear since the systematic 1:250 000 scale mapping of the Yilgarn Block was completed by the Geological Survey in 1979, it is appropriate that we deal with the geochronology within this
J. R. DE LAETER, W . G. LIBBY & A. F. T R E N D A L L
Fig. 1.
Map showing the Archaean areas in Western Australia.
G E O C H R O N O L O G Y OF WESTERN AUSTRALIA
147
TABLE I
Geological data from the Yilgarn Block
Locali ty Kondini n Wongan H i l l s Wi1 una-Leonora-Kookynie Cue Mt Magnet Paynes Find P e r e n j o r i - P a y n e s Find Koolanooka H i l l s ' Kal goojfl ie-Norseman Warda-Warda Dalgaranger Mount Edoii Mount Mulgine I Mount Mulgine I I Wee1hamby Lake Koolanooka Moorine Rock Northam Toodyay Dale Bridge Armadale-Canning Dam Bowgada-Koo 1 anooka S i d i n g Koolanooka H i l l s Morawa-Mullewa East of P i n j a r r a Dumbleyung Corrigin-Brookton Dang in (Low Age) Dang in (Hi gh Age) Bruce Rock York Dood1aki ne-Mer red i n Doodlaki.ne-Kel l e r b e r r i n Doodlakine Boreho1e Dale B r i d g e - V o t i n g Bencubb i n-Wya1katchem Cadoux-Bonnie Rock Wi 11iams-Arthur R i v e r Poona-Da 1garanga (Inner) Poona-Dal'garanga (Outer) Mungari Mungari Kurrawang Kurrawang Kurrawang Menangina Menang i ha . MertondaleKamba1 da Kamba1 da KambaIda . KambaIda Kamba1 da Kamba1 da Kamba1 da Kamba1 da Kamba1 da Kamba1 da Kamba1 da ~ KambaIda Mount Windarra Mungar i . Karamindie Soak Wheeler Rocks Karon ie S i d i n g Stennet ROcks Menangina Rocks Kamba1 da Buldan ia Rocks Buldania Rocks Mungari Mengang i na Rocks Wheeler Rocks KambaIda Karon ie S i di ng Stennet Rocks Stennet.Rocks , Buldania Rocks Isolated Hill Marda Complex Mount Keith . Mount Kei th Mount Keith Mount K e i t h Mount Keith. Mount Kei th Mount Keith Mount K e i t h Mount Keith
Rock Type Granite Grani t e Granite Granite Granite G r a n i t e and porphyry Grani t e Lepidolite Lepidoli te Lepidoli te Muscovi t e Muscovi t e Muscovi t e Quartz porphyry G r a n i t e and p e g m a t i t e Gne i s s Gnei s s Gnei s s Gneiss Gneiss Gnei s s Gneiss Gneiss Gneiss Gneiss Gneiss Gnei s s Gneiss Gneiss Granite Granite Grani t e Grani t e Granite Grani t e Grani t e Grani t e - a d a m e l 1 i t e Granodiorite Granite Biotite Na g r a n i t e Biotite Biotite Gneiss, s l i c e s Grani t e Granite " f r e s h " Footwa11 b a s a l t Hanging wal l m e t a b a s a l t M e t a b a s a l t Pb-Pb P l a g i o c l a s e - t o t a l rock Blotite Biotite Microc1ine Plagioclase F e l s i c porphyry Pb i s o c h r o n F e l s ic porphyry B i o t i t e footwa11 Mi c r o c l i n e t o s o d i c g r a n i t e Granodiorite Granite Granite Adamel1ite to granite Grani t o id Granitoid Grani t o i d Granodiori t e Grani t o i d Granitoid Granite Granitoid Adame11i t e Granodiori t e Granitoid Grani t o i d Gran i t o i d Gran i t o i d Adamel1ite V o l c a n i c rocks Granod i o r i t e Pegmatite Apli te cobbles A p l i t e vein Brot i t e Bioti te Biotite Bioti te Biotite
Age and e r r o r
4 4 4 4 4 4 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 v 2 4 1 It 4 5 1 1 2 1 7 4 4 4 4 2 7 1 4 4 1 8 8 8 8 8 8 8 8 1 10 10 10 10 10 10 10 10 1 1 1 1 3 3 4 4 4 4 4
2811 ± 1560 2745 ± 565 2641 ± 41 2608 ± 68 2706 ± 264 25^9 ± 117 2562 ± 20 2628 2599 2623 2632 2614 2562 28 2505 ± 2534 ± 57 2918 + 207 3154 ± 57 3488 ± 654 3094 ± 792 2869 ± 127 2520 ± 632 2768 ± 71 2553 ± -106 2556 ± 162 2583 ± 2061 2622 ± 297 3236 ± 373 2638 ± 366 2631 ± 21.1 2674 ± 115 2676 ± 269 2707 ± 92 2632 ± 232 . 2639 ± 187 2573 ±. 553 2909 ± 768 23 2535 ± 2550 ± 51 20" 2565 ± 2665 ± 35 2824 ± 160 2521 ± 50 2496 + 50 41 2723 ± 2497 ± 256 36 2550 ± 30 2633 ± 50 2579 ± 2663 ± 270 80 2692 ± 2555 2572 2541 + 35 60 2592 ± 27IO + 200 2521 ± 40 2550 ±. 25 35 2559 ± 60 2565 ± 2640 ± 35 35 2705 ± 2632 ± 28 74 2691 ± 2671 ± 79 2640 ± 53 2760 ± 7 0 35 2655 ± 20 2581 ± 2450 2430 2410 •2515 2490 2400 2550 2550 25 2537 ± 80 2579 ± 2632 ± 17 18 2481 ± 20 -2565 ± 50 2510 ± 2599 2598 2568 2568 2552
Long i tude
•x. s: z o. .7010 ± .0220 .6860 ± .0620 .7030 ± .0010 .7020 ± . 0060 .7040 ± .0100 .7040 ± .,0050 .7030 ± .0007
2 3 3 4 3 3 4
4 4 11 6 5 6 27
.7090 ± .0030 .7160 ± .0100 .7040 ± .0030 .7000 ± .0010 .6920 ± .0070 .7020 ± . 0210 .7070 ± .0060 . 7 1 7 0 ± .0160 .7040 ± .0020 .7040 ± .0010 .7040 ± .0030 .7080 ± .0550 .7140 ± .0020 .7060 ± .0040 .7030 ± .0040 .7120 ± .0090 .7040 ± .0050 .6290 ± .0220 .7000 ± .0060 .7040 ± . 0060 .6990 ± .0120 .7040 ± .0730 .6990 ± .0260 .7028 ± .0010 .7031 ± .0009
2 2 2
• 7019
1
6 9 10 5 3 5 10 4 6 4 4 3 4 5 5 8 8 6 16 4 8 .3 6 7 5 1 1 10 1
.7100 .7010 ± .0020 .7023 ± .0015 .7086 ± .0123 . 7 0 1 3 ± .0001 . 7 0 1 3 ± .0002
3 4 2 4 2 3 3 2 3 2 4 4 3 3 3 2 3 2
1 1
.7014 ± .0003 .7014 ± .0003 .7016 ± . 0008 . 7 0 1 7 ± .0011
3
.7014 ± .0003 .7019 ± .0003 3 1 4 3 2
13 4 7 24 9 11 2 1 1 1 13 10 44 1 16 13 5 12 12 13 2 12
32° 30' 30° 53' 27° 57' 28° 30' 29° 22' 29° 13' 3 1 ° 30' 27° 35' 27° 47' 29° 20' 29° 12' 29° 12' 29° 12' 29° 18' 31° 18' 31° 39' 31° 33, 32° 05' 32° 09' 290 18'
118° 16' 116° 45' 120° 50' 1 1 7 ° 49' 1 1 7 ° 00' 116° 13' 121° 30' 1 1 7 ° 08' 1 1 7 ° 00' 117° 39. 116° 58' 116° 58' 116° 26' 116° 04' 119° 07' 116° 40' 116° 28' 116° 48' 116° 05' 116° 08' 116° 13' 29° 13' 29° 54' 115° 48' 32° 37' 115° 56' 330 1 1 7 ° 45' 19' 32° 21 1 1 1 7 ° 28' 32° 03' 1 1 7 ° 16' 32° 03' 1 1 7 ° 16' 118° 09' 3 1 ° 53' 116° 46' 3 1 ° 53' 33. 118° 05' 31° 31° 38' 1 1 7 ° 49' 3 1 ° 35* 1 1 7 ° 53' 320 03' 1 1 7 ° 17' 31° 00' 117° 39' 30° 38' 1 1 7 ° 49' 330 12' 116° 58' 27° 20' 1 1 7 ° 30' 1 1 7 ° 15' 27° 30' 121° 20' 30° 50' 30° 50' 121° 20' 121° 20' 30° 50' 121° 20' 30° 50' 121° 20' 30° 50' 122° 1 1 ' 29° 53' 290 53* 122° 1 1 ' 28° 30' 121° 05' 121° 40' 3 1 ° 12' 121° 40' 31° 12' 121° 40' 31° 12' l 121° 40' 31° 121° 40' 31° 12' 121° 40' 3 1 ° 12' 121° 40' 3 1 ° 12' 121° 40' 31° 12' 121° 40' 31° 12' 121° 40' 3 1 ° 12' 121 6 40' 31° 12' 121° 40' 3 1 ° •12' 122° 15' 28° 30' 121° 15' 30° 53V 121° 22' 3 1 ° 02' 120° 53' 32° 1 1 ' 122° 34' 30° 56' 121° 33' 32° 36' 121° 56' 29° 50'. 121° 39' 31° 13' 32° 05' • 122° 02' 32 p 05' 122° 02' 121° 15' 30? 29° 50' 121° 56' 120° 53' 32° i i ' 121° 39' 31° 13' 122° 24' 30° 56' 121° 33 32° 36 V 121° 33' 32° 36' 122° 02' 32° 05' 28° 45' 123° 45' 30° 10' 119° 20' 120° 33' 26' 27° 120° 33' 27° 26' 120° 33' 27° 26' 120° 33' 27° 26" 120° 3 3 . 27° 26' 120° 33' 27° 26' 120° 33' 2 7 ° 26' 120° 33' 27° 26' 270 26' 1 20° 33'
5 ' 5 330 205 140 5 160 0 0 0 0 0 0 2 0 10 10 . 1 10 12 5 90 5 5 80 2 2 2 10 40 15 0 75 60 120 35 35 80 0 0
1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 ] 1 1 1 1 1 1 1 1 1 I 1 1 1 1 1 2 2 3 3 3 3
. 1 2 1
4 4 5 6 6 6 6 6 6 6 6 6 6 6 6 7 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 9 10 11 11 11 11 11 11 11 .11
]2
3 0 1 .3 . 1 0 0
5V
.7015 ± .0005 .7029 ± • 0015 .7015 ± .00015 . 7020 .7008
1 3
6 4 21 4 2 3
1 10 25 .2 ]
2
11
J. R. D E L A E T E R , W . G. L I B B Y & A . F. T R E N D A L L
148
TABLE I—continued
Rock Type
Locali ty
Biotite Biotite Tonali te Leucogranite and gneiss Leucotonali te Perseverance gneiss Perseverance pegmatite Kathleen valley gabbro Feldspar from pegmatite Muscovite from pegmatite Quartz syeni te Greywacke Grani te Granite Granodiorite, adamellite Adamel1i te Granitoid Adamel1i te Homogeneous to gneissic grani toid Biotite Bioti te Biotite Biotite Bioti te Biotite
Mount Kei th Mount Keith Lawlers Lawlers Lawlers Wei 1 Lawlers Lawlers Lawlers Greenbushes Greenbushes Fitzgerald Peaks Coolgardie Rocky Bore Roadside Bore Dog Rocks Mount Boreas Meckering . Mundaring York Area Meeker ing Meckering Meckering York York Northam Pemberton-Grevi1 lea
a. <o
Age and error
4 4 1 1 1 1 1 1 4 4 1 11 1 1 1 1 1 1 1
2467 2565 2652 ± 20 14 2576 ± 23 2474 ± 34 2625 ± 18 2588 ± 2718 ± 50 2659 2647 2360 ± 96 80 3060 ± 2603 ± 149 2461 ± 92 52 2573 ± 2428 ± 30 2513 ± 246 2760 ± 614 38 2847 ±
4 4 4 4 4 4 1
2512 2454 2392 2473 2506 2376 2327 ± 1624
87 Sr/86Sr Initial Ratio • and Error
.7015 ± .0001 .7022 ± .0002 .7020 ± .0002 .7022 ± .0005 .7624 ± .0068 • 7007 ± .0004
Latitude
Longitude
2 2 2
1 27° 26' 1 . 27° 26' 28° 00' 20 28° 00' 10 28° 00' 15 10 28° 00' 28° 00" 5 21 28° 00' 330 53' 1 330 1 53' 6 32° 55' 31° 55' 25° 47' 5 2 5 ° 7 59' 8 25° 56' 28° 03' 5 4 31° 38' 4 31° 58' 31° 53' 5
120° 33' 120° 33' 120° 30' 120° 30' 120° 30' 120° 30' 120° 30' 120° 30' 116° 04' 116° 04' 121° 10' 121° 10' 117° 08' 118° 12' 118° 16' 121° 57' 116° 58' 116° 12' 116° 46'
11 11' 12 12 12 ' 12 12 12 13 13 14 15 16 .2 16 16 17 .1 18 .118 10 18
3
31° 38' 31° 38' 31° 38' 31° 53' 31° 53' 31° 39' 34° 04'
116° 58' 116° 58' 116° 58' 116° 46' 116° 46' 116° 41' 116° 19'
18 18 18 18 18 18 19
1 1 1 4 1 4
.7044 ± . 0020
4
• 7095 ± .0067 • 7058 ± .0021 • 7023 ± . 0005 .7081 ± .0020 .7062 ± .0073 .7023 ± .o4io .7019 ± .0010
2 1 1
.7030 ± ,.0092
6
0 23 0 0 0 10
* For reference see caption, Figure 2. t 1, Rb:Sr isochron, whole rock; 2, Rb:Sr isochron, multiple whole rock plus minerals; 3, Rb:Sr whole rock model age; 4, Rb:Sr mineral model age; 5, Rb:Sr whole rock on successive sawn slabs; 6, U:Pb zircon; 7, Pb:Pb whole rock (with or without mineral points); 8, Pb:Pb mineral isochron; 9, Pb:Pb galena; 10, Pb:Pb potassium feldspar model age; 11, Sm:Nd.
and mineral Pb-Pb ages, zircon U-Pb ages, and Pb model ages, are so far comparatively few, and only one Sm-Nd age had been published by January 1980, although more appear elsewhere in this volume. Our commentary is therefore correspondingly weighted toward the significance of the Rb-Sr results. The procedures of Rb-Sr isochron interpretation have become well established since the whole-rock method was established nearly two decades ago (Nicolaysen, 1961; Compston & Jeffery, 1959). The rectilinearity of the isochron measures the acceptability of a hypothesis that at some past time related to the isochron slope, all parts of the rock body concerned had a uniform initial Sr/86Sr ratio (Rj) corresponding to the ordinate intercept. Further, if the Rb/Sr ratio of the whole rock body is known to have been constant since formation, the R provides a measure of the crustal residence time of its component material, prior to the age indicated by the isochron. However, the elegance and simplicity of this standard" interpretational paradigm gloss over a number of associated problems which remain controversial. Firstly, chemical closure since the time recorded by the isochron must be assumed, COMMENTARY at least in respect to Rb and Sr; any large-scale General movement of either or both of these elements into The geochronological data available for the or out of the rock body in question will destroy older Precambrian of Western Australia are the credibility of the results. Secondly, the nature heavily weighted toward the Rb-Sr method. Rock of the geological event, if any, which is marked framework, and the boundaries defined by Gee et al. are therefore included on Figure 2. All isotopic analyses yielding ages older than 2400m.y. from rocks within the boundaries of these two blocks are listed in Tables I and II and displayed in Figures 2 and 3. The numbered references in Table I are identical to those given in the caption of Figure 2, and the references for Table II are similarly included in the caption of Figure 3. All data published or made available to us as in press" at January 1980 are included in these Figures and Tables, except for data published before 1968, and included in the review of Compston & Arriens (1968). Tables I and II list the locality (including the latitude and longitude), the rock type, the number of samples involved and their approximate physical separation. For Rb-Sr isochrons the initial Sr/ Sr ratios and error limits (mostly at the 95% confidence level) have been included, as well as the model number of Mclntyre et al. (1966) where this regression method has been used. All Rb-Sr ages have also been modified in accordance with a decay constant of 1.42 x IP - Vr - for Rt) (Steiger & Jager, 1977). 4
87
1
1
87
86
87
{
GEOCHRONOLOGY
OF WESTERN
AUSTRALIA
149
TABLE II
Reference*
Approximate Scatter (km
Mc1ntyre Model Number of Points
Geochron ological information from the Pilbara Block 87 Sr/ 86Sr Initial Ratio Locali ty Rock Type Age and Error and Error La t tude i Long itude Copper Hills Porphyry Porphyritic felsite 2819 ± 66 • 7303 ±.0119 6 21° 40' 119° 57' .3 20 Mount Edgar Batholith Mi gmat i te,gnei ss 55. 3059 ± 366 .7016 ± .0047 3 10 21° 07' 119° 18 21 Older Granite Grani te Moolyella Granite Granite 26 14 ± 95 • 7397 ±.0419 3 6 21° 05' 119° 57' 11 21 Shaw Batholi th Gnei ss 2889 ± . 7 020 ± . 0 010 83 1 0 21° 30' 3 20' 119° 29 22 Cooglegong Adamel1i te Adamel1 ite 2551 ± 128 .7303 ±.0280 2 7 21° 30' 1-19° 23 22 Bamboo Greek Porphyry Porphyritic felsite 2760 ± 516 .7010 ± .0201 3 10 20° 56' 120° 20' 14' .1 23 Mt Newman 70-Mile Quarry Dark granito id 2882 ± 60 21° 118° 3 52' 0 24 1 3 17' Woodstock 1527 Gneiss ic granite 2969 ± 45 11 21° 37' 118° 54' 0 24 Mt Newman 12J-Mile Quarry Foliated granite 2961 ± 21° 7 01' 0 24 5 59' 119° Coog1egong Gneissic granite ' 2936 ± 9 21° 21 ' 7 0 24 29' 119° Tambourah 1511-1514 Migmatite and granite 3070 ± 12 5 21° 44' 119° 21' 0 24 Mt Newman 13-Mi1e Quarry K-feldspar 2920 20° 26' 118° 7 0 24 39' 1 Woodstock 1526 Gneissic granite 2786 ± 38 21° 118° 54' 0 24 3 37 Woodstock 1527 Gneissic granite 2769 ± 13 4 21° 37' 118° 54' 0 24 Tambourah 1512-1513 Migmatite and granite 2938 ± 33 0 24 7 21° 44' 119° 21' Cookes Greek Granite Adamellite, biotite aplite 2514 4 37 -7307 ±.0097 1 4 21° 38- 120° 26' .1 25 Cookes Creek Granite Bioti te 2552 21° 38' 120° 26' 25 Lookout Rocks Granite, adamellite 2540 ± 60 .7086 ±.0129 3 12 21° 56' 121° 25' 10 25 Hardey Sandstone Shale 2707 ± 151 •6950 ± .0129 3 7 21° 27' 120° 32' .03 26 H i11s i de Station Galena in older gneiss 2420 1 21° 34' 119° 30' 27 Doolena Gap Galena in fault in u/basic 3360 1 20° 50' 119° 40' 0 27 and metasediment Galena in dolerite Braeside MC 450 2710 1 21° 16' 121° 10' - 0 27 Galena fault in basalt Braeside ML 295 2720 1 2]o 05' 121° 00' 0 27 Galena in chert, rhyoli te Big Stubby 3500 21° 13' 119° 46' 28 0 Duffer Format ion 3452 ± 16 21° 22' 119° 37' 29 Shaw Batholith Mi gmat i te 2 21° 40' 119° 30' 3417 ± 40 30 Shaw Batho 1ith Grani to id 3280 ± 20 2 21° 40' 119° 30' 30 Shaw Batholith Muscovi te i n pegmat i te 2830 ± 30 30 3 21° 40' 119° 30' " For references see caption, Figure 3+ 1, Rb:Sr isochron, whole rock; 2, Rb:Sr isochron, multiple whole rock pi us minerals; 3, Rb:Sr whole rock model age; 4, Rb:Sr mineral model age; 5, Rb:Sr whole rock on success ive sawn slabs; 6, U:Pb zircon, 7, Pb:Pb whole rock (with or wi thout mineral points); 8, Pb:Pb mineral isochron; % Pb:Pb galena; 10, Pb:Pb potassium feldspar model age; 11, Sm:Nd.
by Sr isotopic homogeneity, is not always clear. Thirdly, and as a corollary, it is often controversial what textural or other features of the rock are associated with this event. It is variously argued, under different circumstances, that isotopic homogenization has been effected by a late thermal episode which has left little or no other impression on the rocks, and that elsewhere, by contrast, the later imposition of a strong tectonic fabric has not affected Sr isotopic distribution, at least on a hand-specimen scale. Fourthly, the preferential movement of radiogenic Sr into and out of minerals, on a small scale, or rock bodies, on a regional scale, remains mysterious and contentious. Experimental evidence is simply not available to settle these questions, and the acceptability of the "standard" isochron interpretation, outlined above, must therefore be assessed to some extent by whether or not it "works": that is, whether or not it appears consistent, over a wide region, with hypotheses based on other geological evidence. Such an assessment, which is the main objective of our commentary, is most conveniently made by the use of the Sr evolution diagram (Fig. 5), which displays in an inter-related and easily assimilable way the main characteristics of a group of samples yielding an isochron.
Gneiss and Granitoid Ages in the Two Main Archaean Blocks Gneiss and granitoid ages form by far the greatest proportion of those included in Tables I and II, and a crude assessment of their general distribution can be obtained from Figure 4, which may be compared with the conceptual Figure 4 of Aniens (1971). For a closer assessment of the significance of the Rb-Sr isochrons included on Figure 4 in terms of the strontium evolution diagram, it is necessary that, for each isochron, Rj, age, and the Rb/Sr ratio of the rock body, all be reasonably well established. The data of Tables I and II were therefore required to meet a number of arbitrary conditions before acceptance for Figure 5. Isochrons were only accepted if: (a) the error in Rj did not exceed ± 0,002; (b) the error in age did not exceed ± 120 m.y.; (c) the Rj was not less than 0.70; (d) they were derived from multiple whole-rock analyses (i.e. not single rock plus separated minerals); (e) the Rb/Sr ratios of the component samples gave a credible measure of the total Rb/Sr ratio of the rock body (i.e. isochrons heavily reliant on a small number of highly enriched samples were excluded);
J. R. DE LAETER, W. G. LIBBY & A. F. TRENDALL
150
F l g . 2.
Map of the Yilgarn Block showing the locations of geochronological data available since 1970. Numbers ^ ™ t e , S T C e Publications: 1 - A r r i e n s (1971); 2-Muhling & de Laeter (1971); 3-Compston & Twek (1973); 4—de Laeter et al. (1973); 5 - W o r d e n & Compston (1973); 6 - R o d d i c k (1974); 7-Roddick (writt. (1975); 9 B u n t i n 8 e< (1976); 10—Hallberg et al. (1976); 11-Roddick et al, w « ? P e r 1' a i ( 1 9 7 8 ) ; 1 3 - d e L a e t e r & Buckley (1978); 1 4 - d e Laeter & Lewis (1978); 15.n o ™ Wasserburg(1978); 16-WilIiamsef <7/. (1978); 17-Bunting & Chin (1979); 18-Libby&de Laeter (1979); 19—Rosmane/al, (1980).
i/^V
GEOCHRONOLOGY OF WESTERN AUSTRALIA
151
T Y P E OF A N A L Y S I S Rb : Sr (w.r.)
A
GE
24 °
2600 - 2700 m.y.
• Ea
2500 - 2600 m.y.
DO
> 2800 m.y. 2700 - 2800 m.y.
Rb : Sr (min)
•
< 2500 m.y
Sm : Nd
Pb
•
A
€
i>
1
0
0
100
0
O
km
O
Outline of Archaean Pilbara Block, Sylvania Dome and Gregory Range —
Outline Of Pilbara - Hamersley craton
I Fig. 3.
(f)
—
Map of the Pilbara Block showing the location of gepchronological data available since 1970. Numbers indicate source publications: 20—de Laeter & Trendall (1970); 21—de Laeter & Blockley (1972); 22—de Laeter et al. (1975); 23—Trendall (1975); 24—Oversby (1976); 25—de Laeter et al. (1977); 26— Hickman & de U e t e r (1977); 27—Sangsfer & g f o o k (1977); 28—Richards (1978); 29— Pidgeon (1978a); 30^= Pidgepn (19786).
the, t m t e s t
ditfanee M w m
was
less than 10 j$m. Isochron§ g ^ j f i p n s are termed "superior isochrons" for the purpose of the present discussion. Only nine of the 69 Rb-Sr isochrons listed in Tables I and II are superior isochrons; three others from Table I which may qualify cannot be included because the analytical data have not been published. The nine superior isochrons, which are included in Figure 5A and identified in the figure caption, come from the
Eastern Goldfields Province, Southern Cross Province, and Western Gneiss Terrain. All fall within a broad zone, marked in the figure. One interpretation of this pattern, assuming singlestage derivation from the mantle, is a sialic crust with R b / S r about 0.28 emplaced over a timerange between about 258G and 2870 m.y. Multistage processes would require a greater age. Two of the points (1 and 3) individually support an older mantle derivation. If condition (f) above is not applied, eight
J. R. DE LAETER, W. G. LIBBY & A. F. TRENDALL 152 more Yilgarn Block isoehrons and one Pilbara Gneiss Terrain (Gee et al. 1981), it must be Block isochron become available for use, and pointed out that point 10 is derived from a 5these appear, together with the first nine, in sample isochron in which the farthest samples are Figure 5B, numbered 10-18. We will call these separated by some 10km. Using the criteria we ''superior regional isoehrons". Six of these have selected for isochron acceptance, the Rb-Sr superior regional isoehrons, from the Eastern evidence for a marked age difference between the Goldfields and Murchison Provinces, fall within Western Gneiss Terrain and the remaining parts the broad zone defined by the points of Figure of the Yilgarn Block cannot be seen as definitive. 5A. The remaining three (10, 14 and 17) fall out- The broad zone of crustal evolution defined by side this zone, and come from the Western Gneiss the bulk of these superior isoehrons is consistent with the data published by Chapman etal. (1981), Terrain and the Pilbara Block. Although these two superior regional isoehrons in that all the isoehrons of that paper could have from the Yilgarn Block lend support to the con- been derived secondarily from material within the cept of a significantly older age for the Western zone. f
PILBARA
BLOCK •10
-45 40 -35 YILGARN
BLOCK
-30 -25 -20
-15
Granitoid and Gneiss
Mafic and Sedimentary Rock -10
Hi
3600
3400
3200
3000 A g e ( m i l l i o n s of y e a r s )
2800
2600
2400
Fig. ,4. Wstogramshowing the distribution in time of ages greater than 2400 m.y. determined since 1970 on rocks of the two Archaean blocks of Western Australia.
GEOCHRONOLOGY OF WESTERN AUSTRALIA
Glikson (1977, 1978, 1979) has incorporated selected Western Australian data on strontium evolution diagrams in three publications in a way which clearly illustrates the need for care in their use. His plotting of much of the data of our Tables I and II without indication of Rj or age error, in his Figure 1 (Glikson, 1977) and Figure 6 (Glikson, 1978), imply a support for the common existence of Archaean gneisses with high Rjs, which our review does not confirm. Similarly, his Figure 6 (Glikson, 1978) selectively presents data to support a hypothesis involving very early mantle-derived precursors of the crust of the Eastern Goldfields Province. No superior isochron is available from granitic rocks in the Pilbara Block, and the only superior regional isochron in our sense, from the Shaw Batholith, is included on Figure 5B (17). While it supports the older age for the Pilbara Block by comparison with the Yilgarn Block which is suggested in Figure 4, the exact amount of this age difference, and its consistency, have clearly not yet been established on Rb-Sr evidence. Evidence from methods other than Rb-Sr is not yet sufficient from either block to provide widespread comparative age estimates. The Pb-Pb mineral ages of Oversby (1975) for eight granitoids of the southern Eastern Goldfields and Southern Cross Provinces varied between 2630 and 2760 m.y., and were slightly older than Rb-Sr mineral isochrons from the same rocks. Oversby suggested, largely on the basis of model fi values, that sialic rocks must have been present in the Norseman area 3300m.y. ago, whereas in the Kalgoorlie area the formation of continental crust must have occurred later . The early presence of continental crust in the Coolgardie area is also supported by the only Sm-Nd result so far available from the Yilgarn Block (McCulloch & Wasserburg, 1978); however, this isolated model age of 3060 ± 80 m.y. from a single greywacke sample, cannot be regarded as definitive. Oversby (1976) has also provided the only PbPb ages from the Pilbara Block (Table II), which give qualified support for an older age of the granitoids there by comparison with the Yilgarn Block. Nine selected dates from Pb-Pb analyses of total rocks plus, in some cases, minerals, ranged from 2769 to 2070m.y. with a mean of 2674m.y. Support for an age difference between the two blocks comes from the Pilbara zircon UPb age of 3417m.y. by Pidgeon (19786) from the Shaw Batholith. Greenstone-belt Ages in the Two Main Archaean Blocks Until recently there were negligible definitive data on the ages of the greenstone belts. The zir-
153
con U-Pb age of 3452 ± 16m.y. from the Duffer Formation dacite of the eastern Pilbara Block (Pidgeon, 1978a) was the first reliable evidence. This has been followed closely by corroborative Sm-Nd ages reported by Hamilton et al. (1981), too recently for inclusion in Table II. The Pb model ages of Richards (1978) and Sangster & Brook (1977) are also consistent with those results. In the granite-greenstone provinces of the Yilgarn Block there are virtually no convincing primary, greenstone-belt ages available. The Kathleen Valley Gabbro Rb-Sr age of Cooper et al. (1978), at 2718 ± 50m.y., is a superior isochron and is plotted on Figure 5B (18) with granite ages for comparison. It gives general support to hypotheses involving near contemporaneity of greenstone and granite genesis. The extensive work of Roddick (1974) in the Kambalda area found nothing older. The single SmNd age from the Coolgardie area (McCulloch & Wasserburg, 1978) of 3060 ± 80m.y., although from a greywacke, indicates the age of the parent material, rather than the depositional age of the greywacke. We are aware of the recent work of Nieuwland & Compston (1981), which reports substantially older depositional ages for sedimentary rocks within the Western Gneiss Terrain, but have not included it in our review. These results are clearly in good accord with the geological concepts of Gee etal. (1981). Other Archaean Areas and Rocks In Western Australia the term Archaean has generally been applied to rocks older than 2400m.y., by contrast with the 2500m.y. boundary currently recommended by the Subcommission on Precambrian Stratigraphy (James, 1978). Compston and others are reviewing the geochronological data for the rocks of the Hamersley Basin, which unconformably overlie the Pilbara Block along its southern edge. They conclude that the basal rocks of the basin, the Fortescue Group, may be substantially older than 2500 m.y., and legalistically "Archaean". The great regional unconformity at their base, which marks the southern margin of the Pilbara Block, has hitherto been accepted as the equivalent of unconformities on other continents that historically define the conceptual base of the Proterozoic. We do not comment further on this problem here, and have not included the Hamersley Basin in our review, except for incorporating the Hardey Sandstone of Hickman & de Laeter (1977) in Table II. South of the Hamersley Basin the Sylvania Dome (Fig. 1) is generally accepted, on tectonic
J. R. DE LAETER, W. G. LIBBY & A. F. TRENDALL
154
-0.707
0.706
0.705
-0.704
I
-0.703
0.702 r
0.701
0.70Q
0.699
Yea rs
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r• 0 . 7 0 7
/
X Granite, etc h
I
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A Gabbro'
I
>
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I
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~r 1 f I I I _L
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—~~9To i I _L i i
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—
'^ffl8
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if
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/
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/
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-0.702
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/
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•—f—-
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r ,*13 / IJ
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i I I
2.7
~T 2.6
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Strontium evolution diagram for Archaean superior and superior regional isochrons since 1970 fox gneisses and granitoids of the Yilgarn and Pilbara Blocks, A. Superior isochrons of the Yilgarn Block. B. Includes information on A plus superior regional isochrons of Yilgarn Block and Pilbara Block, and one superior isochron for a Pilbara Block gabbro. Vectors indicate the direction from which 87Sr/86Sr has evolved based on present «7Rb/86Sr ratios. Key to numbers 1-Sliced Menangina gneiss, 2-Menangina granite (de Laeter et al., 1973); 3-Lawlers tonalite, 4-Perseverance gneiss, 5-Lawlers Well leucdtonahte (Cooper et al 1978); 6-Isolated Hill (Bunting et al., 1976); 7 - D o g Rock (Williams \etal, 1978); 8-East of Pinjarra (Arnens, 1971); 9-Quartz syenite, Fitzgerald Peaks (de Laeter & Lewis, 1978); 10Toodyay gneiss (Arnens, 1971); 11-Mount Keith granodiorite (Roddick et aL, 1976); 12-Lawlers leuco? Q 7 i f A I " ' 1 5 8 ( S 0 ,° P e r aL > 1 9 7 8 ) ; *Wiluna-Leonora-Kookynie; 14-Morawa-Mullewa (Arriens, ; g a r a n g a ( i n n e r ) ; 16 ^ n -Poona-Dalgaranga (outer) (Muhling & de Laeter, 1971); 17Shaw Batholith (de Laeter et al., 1975); 18—Kathleen Valley Gabbro (Cooper et al
1978)
GEOCHRONOLOGY OF WESTERN AUSTRALIA and lithological grounds, to be likely to consist of Archaean rocks. However, the only geochronological evidence so far available (Blockley et al., 1980) provides a granite age of 2235 ± 54m.y., which cannot easily be accepted as a primary age. The equivocal isotopic evidence for an Archaean age for any rocks of the East Kimberley region was discussed by Compston & A n i ens (1968), and the situation has not changed since then.
General Comments and Conclusions If previously published results are included with the newly available data listed in Tables I and II the following analyses acceptable by current standards are available from the older Precambrian rocks of Western Australia: 837 wholerock and 142 mineral Rb-Sr analyses, which have been variously grouped to form 145 isochrons, 61 whole-rock and 131 mineral Pb isotope analyses, 15 K-Ar analyses, 7 U-Pb mineral analyses and one Sm-Nd analysis, reported in 40 papers to which a total of 55 individual authors have contributed. A substantial amount of additional data is in press. When the Rb-Sr evidence is examined critically, little of it is definitive, but the better- defined isochrons seem to support the following general conclusions: (i) for the Yilgarn Block . (a) the time of major generation of gneisses and granitoids was remarkably consistent over almost the entire sampled area of the block, and peaked at about 2700 m.y.; (b) there is little evidence for any wide spread about this peak, and evidence for much older ages is still not compelling; (c) over a large area of the block the granitoids had a relatively short prior crustal history, and by implication were generated over a relatively brief period; (d) the time interval between the formation of the greenstone belts and the peak of granitoid emplacement is of unknown length, but the available evidence suggests that it was small;
155
(ii) for the Pilbara Block (a) the main period of gneiss and granitoid generation was also uniform over the area of the block, and was earlier than that in the Yilgarn Block; (b) the error limits attached to granitoid ages make it uncertain whether emplacement was restricted within a narrow time range; (c) the highly fractionated small-volume granitoids ("tin granites") are significantly younger; (d) greenstone belt ages are about 300-500 m.y. older than the peak granitoid age. In suggesting these as the best present conclusions we have deliberately chosen to accept and apply the " s t a n d a r d " Rb-Sr interpretational method outlined earlier. Hypotheses involving large-scale regional migration of Rb, Sr, or radiogenic Sr have not been examined, since their acceptance involves speculative possibilities better pursued in a more general context. Our conclusions concerning the Yilgarn Block are based on single-stage mantle derivation which seems consistent with Figure 5. The Rb-Sr evidence does not preclude multi-stage derivation, as suggested for example by Davies & Allsopp (1976), provided that processes of large-scale isotopic and chemical homogenization at each stage are envisaged. The available Pb-Pb, zircon, and Sm-Nd results are consistent with the above conclusions, although the recent zircon results of Nieuwland & Compston (1981) suggest that wider application of this method could lead to modification of (i)b above. In interpreting Rb-Sr data from either block, care is needed to distinguish sampling accidents and slow cooling effects from real events. Probably the single most striking geochronological characteristic of either of the Western Australian Archaean blocks is the vastness of the area of the Yilgarn Block over which the isotope evidence points to penecontemporaneous generation of sial over a relatively short interval. ACKNOWLEDGMENTS This paper is published with the approval of the Director of the Geological Survey of Western Australia.
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LIBBY, W. G., 1980: TWO 4 'anomalous" isochrons from the vicinity of Newman. Ann. Rep. geol.
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J. R. DE LAETER, W. G. LIBBY & A. F. TRENDALL
CHAPMAN, H . J . , BICKLE, M . J . , DE LAETER, J . R . , B E T TENAY, L . F . , GROVES, D .
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R. A., & G O R T O N , M. S., 1981: Rb-Sr geochronology of granitic rocks from the Diemals area, Central Yilgarn Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 173-186. BINNS,
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GLIKSON,
W., & JEFFERY, P. M., 1959: Anomalous "common strontium" in granite. Nature, Lond.s 184, 1792-1793.
, 1978: Archaean granite series and the early crust, Kalgoorlie System, Western Australia; in Windley, B. F., & Naqvi, S. M. (Eds), Archaean Geochemistry, 151-173. Elsevier, Amsterdam. , 1979: Early Precambrian tonalite-trondhjemite sialic nuclei. Earth-Sci. Rev., 15, 1-73.
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HALLBERG, J . A . , JOHNSTON, C . , & B Y E , S . M . , 1976:
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cambrian geochronology of Australia. Can. J. Earth Sci., 5, 5 6 1 - 5 8 3 . COMPSTON,
COOPER, J . A . , NESBITT, R . W . , PLATT, J . P . , & M O R T I -
MER, G . E . , 1978: Crustal development in the Agnew Region, Western Australia, as shown by Rb/Sr isotopic and geochemical studies. Precamb. Res., 7, 31-59. DAVIES, R . D . , & ALLSOPP, H . L . , 1 9 7 6 : Strontium iso-
topic evidence relating to the evolution of the Lower Precambrian granitic crust in Swaziland. Geology, 4, 5 5 3 - 5 5 6 . DE LAETER, J . R . , CHAPPELL, B . W . , & COMPSTON, W . ,
1973: An internal whole-rock age from an Archaean migmatite from Lake Rebecca near the margin of the Menangina Batholith, Western Australia. Abstr. Aust. N.Z. Ass. Adv. Sci. Sect. 1, Geology, 45th Congr., Perth, 146. G., 1 9 7 2 : Granite ages within the Archaean Pilbara Block, Western Australia. J. geol. Soc. Aust., 19, 3 6 3 - 3 7 0 . —, 1978: Age of a tin-bearing pegmatite at Greenbushes. Ann. Rep. geol. Surv. West. Aust. for 1977, 48-49. DE LAETER, J . R . , & BLOCKLEY, J .
DE LAETER, J . R . , H I C K M A N , A . H . , TRENDALL, A . F . ,
J. D., 1977: Geochronological data concerning the eastern extent of the Pilbara Block. Ann. Rep. geol. Surv. West. Aust. for 1976, 56-62. & LEWIS,
DE LAETER, J . R . , & LEWIS, J . D . , 1 9 7 8 : T h e a g e o f t h e
syenitic rocks of the Fitzgerald Peaks, near Norseman. Ann. Rep. geol. Surv. West. Aust. for 1977 56-60. DE LAETER, J . R . , LEWIS, J . D . , & BLOCKLEY, J .
G.,
1975: Granite ages within the Shaw Batholith of the Pilbara Block. Ann. Rep. geol. Surv. West Aust for 1974, 73-79. DE LAETER, J . R . , & TRENDALL, A . F . , 1 9 7 0 : T h e a g e o f
the Copper Hills Porphyry. Ann. Rep. geol. Surv. West. Aust. for 1969, 5 4 - 5 9 . _ _ — , 1979: The contribution of geochronology to Precambrian studies in Western Australia. J. R Soc. West. Aust., 62, 21-31. G E E , R . D . , BAXTER, J . L . , W I L D E , S . A . , & WILLIAMS,
I. R., 1981: Crustal development in the Archaean Yilgarn Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 43-56.
HAMILTON, P . J . , E V E N S E N , N . M . , O ' N I O N S , R . K . , GLIKSON, A. Y., & H I C K M A N , A. H . , 1981: Sm-Nd dating of the North Star Basalt, Warrawoona Group, Pilbara Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 187-192. H I C K M A N , A. H . , in press: Geology of the Pilbara Block
and its environs. Bull. geol. Surv. West. Aust., 127. HICKMAN,
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LAETER,
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The
depositional environment and age of a shale within the Hardey Sandstone of the Fortescue Group. Ann. Rep. geol. Surv. West. Aust. for 1976, 62-68. L., 1978: Subdivision of the Precambrian—A brief review and a report on recent decisions by the Subcommission on Precambrian Stratigraphy. Precamb. Res., 7, 193-204.
JAMES, H .
LIBBY, W. G., & DE L A E T E R , J .
R., 1979: Biotite dates and cooling history of the western margin of the Yilgarn Block. Ann. Rep. geol. Surv. West. Aust. for 1978, 79-87.
M C C U L L O C H , M . T . , & WASSERBURG, G . J . , 1 9 7 8 : S m -
Nd and Rb-Sr chronology of continental crust formation. Science, 200, 1 0 0 3 - 1 0 1 1 . MCINTYRE,
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T U R E K , A.,
1966: The statistical assessment of RbSr isochrons. J. geophys. Res., 71, 5459-5468.
P. C . , & DE L A E T E R , J . R., 1 9 7 1 : Ages of granitic rocks in the Poona-Dalgaranga area of the Yilgarn Block, Western Australia. Ann. Rep. geol Surv. West. Aust. for 1970, 6 8 - 7 3 .
MUHLING,
L. O., 1961: Graphic interpretation of discordant age measurements of metamorphic rocks. AnnalsN. Y. Acad. Sci., 91, 198-206.
NICOLAYSEN,
D. A., & COMPSTON, W., 1 9 8 1 : Crustal evolution in the Yilgarn Block near Perth, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 159-171.
NIEUWLAND,
V. M., 1975: Lead isotopic systematics and ages of Archaean acid intrusives in the KalgoorlieNorseman area, Western Australia. Geochim., cosmochim. Acta, 39, 1107-1125.
OVERSBY,
—, 1976: Isotopic ages and geochemistry of Archaean acid igneous rocks from the Pilbara, Western Australia. Geochim. cosrriochim. Acta, 40, 817-829.
GEOCHRONOLOGY OF WESTERN AUSTRALIA R. T., 1978a: 3450m.y.-old volcanics in the Archaean layered greenstone succession of the Pilbara Block, Western Australia. Earth planet. Sci. Lett., 37, 421-428.
PIDGEON,
19786: Geochronological investigation of granite batholiths of the Archaean granite-greenstone terrain of the Pilbara Block, Western Australia; in Smith, I. E. M., & Williams, J. G. (Eds) Proceedings of the 1978 Archaean Geochemistry Conference, 360-362. Univ. Toronto, Ontario. J. R . , 1978: Lead isotopes and ages of galenas from the Pilbara region, Western Australia. J. geol. Soc. Aust., 24, 465-473.
RICHARDS,
J. C. ML, 1974: Responses of Strontium Isotopes to Some Crustal Processes. Thesis, Aust. Nat. Univ. [unpublished].
RODDICK,
RODDICK, J . C . , COMPSTON, W . , & D U R N E Y , D .
W.,
1976: The radiometric age of the Mount Keith Granodiorite, a maximum age estimate for an Archaean greenstone sequence in the Yilgarn Block, Western Australia. Precamb. Res., 3, 55-78. ROSMAN, K . J . R . , W I L D E , S . A . , LIBBY, W . G . , & DE LAETER, J. R., 1980: Rb-Sr dating of granitic rocks
in the Pemberton area. Ann. Rep. geol. Surv. West. A ust. for 1979, 97-100.
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in an Australian Zn-Pb-Ba deposit. Nature, Lond., 270, 4 2 3 . STEIGER, R . H . , & JAGER, E . , 1 9 7 7 : Subcommission on
geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth planet. Sci. Lett., 36, 3 5 9 - 3 6 2 . TRENDALL, A. F . , 1 9 7 5 : Preliminary geochronological results from two Pilbara porphyry bodies. Ann. Rep. geol. Surv. West. Aust. for 1974, 1 0 3 - 1 0 6 . TRENDALL, A. F . , & DE LAETER, J . R . , 1 9 7 8 : Chronological bibliography of Western Australian Precambrian geochronology. Rec. geol. Surv. West. Aust., 1978/3. WILLIAMS, S . J . , ELIAS, M . , & DE LAETER, J . R . ,
1978:
Geochronology and evolution of the eastern Gascoyne Province and the adjacent Yilgarn Block. Ann. Rep. geol. Surv. West. Aust. for 1977, 5 0 - 5 6 . WILSON, A .
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WORDEN, J .
CRUSTAL EVOLUTION IN THE YILGARN BLOCK NEAR PERTH, WESTERN AUSTRALIA D. A. Nieuwland & W. Compston Research School of Earth Sciences, Australian National University, Canberra, A.C.T. 2600, Australia ABSTRACT Detrital zircons from quartzite register a concordia U-Pb age of 3340 Ma, which implies the past and possibly present-day existence nearby of crustal rocks of at least the same age. Interlayered gneisses are identified as orthogneiss by field observations and zircon morphology, the latter showing euhedral igneous zircons pitted, necked and rounded to varying degrees by metamorphic corrosion. The U-Pb age of the igneous zircons is 3250 Ma, so that the time of deposition of the quartzite and other sediments is closely constrained. Overgrowths and separate crystals of metamorphic zircon are present in the gneisses and in a cordierite-anthophyllite xenolith. Their U-Pb age is 3180 Ma, which we interpret as a time-point within the first high-grade metamorphism in the region. Post-deformation granitoids have zircon U-Pb ages of 2670 Ma but migmatites contain distinctive zircons that are younger still at ca 2530 Ma. RbSr results for whole-rock samples of the migmatite neosome and granitoids, for recrystallized muscovites and quartz from schist and quartzite, and for primary muscovites from a granitoid and pegmatite all indicate internal ages between 2500 and 2560 Ma. We conclude that migmatisation in at least one locality occurred during a younger, amphibolitic-facies, metamorphic event that encompassed or followed the intrusion of the granitoids and terminated around 2500 Ma. Rb-Sr model ages for the orthogneisses range from 3200 Ma to 2500 Ma and show that the post-emplacement metamorphisms produced a variable loss of 87 Sr from whole-rock samples.
INTRODUCTION The purpose of this paper is to present new isotopic-age determinations on Archaean rock units within the Yilgarn Block near Perth. These include the first zircon U-Pb results from the region using standard laboratory methods. Our isotopic interpretations are constrained by field and petrological evidence for the structural order of events and by zircon morphology. This involved new mapping of sedimentary and tectonic features which have been described elsewhere (Nieuwland, 1980). Our knowledge of the general geology of the region is based on a succession of studies by the Geology Department of the University of Western Australia and on recent 1:250000 mapping by the Geological Survey of Western Australia The longer-term objective of our isotopic work is the definition of Archaean crustal growth. We wish to recognize the oldest material component of the region, to trace its later transformations by recycling, and to distinguish it from new material that may have been added later.
Spec. Pubis geol. Soc. Aust., 7 (1981)
All ages cited in this paper employ the following decay-constants: \87Rb, 0.0142 x 1 0 " 9 y r - 1 \238u, 0.122125 x 10~ 9 yr~ 1 X 235 U, 0.98425 x l O ^ y r " 1 REGIONAL GEOLOGY AND GEOCHRONOLOGY The hybrid gneisses and their amphibolitic inclusions of Prider (1940) have long been recognized as older components of the Yilgarn Block along its western margin at the Darling Scarp where they are intruded by massive granite. Similar rocks termed migmatite by Wilde & Low (1978) are described to the northeast in the Perth 1:250000 Sheet (Fig. 1). More extensive older components are the metasedimentary and metaigneous rocks of the Jimperding Metamorphic Belt (Prider, 1944; Stephenson, 1970; Wilde & Low, 1978), which are also intruded by granitic rocks.
D. A. NIEUWLAND & W. COMPSTON
160
DARLING FAULT
15K
ORTHOQUARZITE BANDED GNEISS (Anb) FELDSPATHIC GNElSS(Anl) MIGMATITE INTRUSIVE GRANITOIDS
REGIONAL GEOLOGY
JIMPERDING METAMORPHIC BELT (based on Geol. Surv.W.A. Perth 1 250.000 sheet
Fig. 1.
Outline of regional geology of the Jimperding Metamorphic Belt, showing localities and rock units discussed in the text.
The ages of the intrusive granitic rocks, where determined by the whole-rock Rb-Sr method, are predominantly ca 2600Ma (Arriens, 1971). The ages of the older gneisses by the same method are much less clear, presumably owing to the effects of later metamorphism. The whole-rock isochrons are poorly fitted even for single localities and the uncertainties in the age are correspondingly high (Arriens, 1971). Nevertheless, the regionally averaged gneiss age at ca 3000 Ma is systematically greater than that of the intrusive granites. This point was reinforced by Libby & de Laeter (1979) using Rb-Sr whole-rock analyses of additional samples.
ZIRCON U-Pb AGES Methods Zircons were separated from the crushed samples by magnetic and density methods, and purified by handpicking. They were dissolved chemically by the method of Krogh (1973), and a mixed 235 U,208pb spike added to an aliquot of the solution for determination of U / P b by isotope dilution. Details of all techniques are given by Williams (1977). The processing blank was ca 0.3 ng for both Pb and U and the spikes calibrated against SRM 613 trace-element reference glass. Replicate analyses of a zircon solution gave 0.3% for the precision of U / P b .
CRUSTAL EVOLUTION IN Y I L G A R N BLOCK
Fig. 2.
SEM photographs of zircons from various rock units in the Jimperding Metamorphic Belt. A: Detrital zircon from the Jimperding orthoquartzite. B: Igneous zircon from the Ringa orthogneiss showing pitting due to metamorphic corrosion. C: Twinned euhedral zircon from the Noondeening gneiss. / D: Very small zircons growing on the face of older zircon in the migmatite. E: Section of zircon from the migmatite showing euhedral core and (white) euhedral overgrowth.
161
162
D. A. N I E U W L A N D & W .
Windmill Hill and Gabidine Springs Quartzites Quartzites are a minor component of the Jimperding Metamorphic Belt, although prominent because of their resistance to erosion. In places they preserve graded- and cross-bedding including herringbone structures, which indicate tidal currents (Blatt et al., 1972). They also contain lenses of fuchsite (chrome-muscovite) interpreted as former clay-rich bands and probably representing flaser structures in the original tidal flats. Zircons in the quartzites would be either detrital or formed during later metamorphism. Samples were collected at arbitrary horizons from quartzite exposures at the above localities (Fig. 1). All zircons extracted (ca 2g per 40 kg of quartzite) appear to be detrital (Fig. 2A). The analytical results are given in Table I. The alignment of data in Figure 3 fits the two separate chords shown almost to within our estimates of experimental error. The upper intercept age is 3341 + 133 - 73 Ma for the Windmill Hill zircons and 3267 + 48 - 25 Ma for Gabidine Springs. We interpret these ages as reflecting different mixtures of detrital zircons which are probably slightly different to their primary igneous ages. We conclude from the good fit to the chords that the range in ages of rocks in the provenance was rather small, in contrast, for example, to the provenance of the Bens Quartzite in the Dalradian of Ireland, in which the zircon fractions are comparatively scattered (Pidgeon, 1969). Jimperding, Ringa and Noondeening
Gneisses
The nature of the protolith of the gneisses. Well-banded quartz-feldspar-biotite gneisses are a major component of the Jimperding Metamorphic Belt. Prider (1944) and McWhae (1946) view the gneisses near Toodyay (Fig. 1) as orthogneisses, originally emplaced as granitic
COMPSTON
sheets. Later workers have interpreted similar gneisses elsewhere as metamorphosed greywackes or arkoses (Stephenson, 1970; Johnstone, 1952) and undoubtedly both ortho- and paragneisses are present in the region as a whole (Wilde & Low, 1978). The nature of the gneiss protoliths is critical to our interpretation of the zircon-age determinations. Consequently field and other evidence for their , origin was reexamined, with the following observations: (i) There is an unequivocal intrusive contact between the Jimperding gneiss and quartzite 200 m upstream from the type section of Prider (1944) at Poison Creek. (ii) At this locality, the gneiss contains numerous xenoliths of quartzite. The xenoliths, bedding planes and the intrusive contact have been rotated almost parallel to the axial plane of the first and principal tectonic deformation Dj. (iii) The intense deformation within the Jimperding Metamorphic Belt has generated a pseudoconcordant contact between gneiss and quartzite for most outcrop planes. The true structural relationship can be seen only on a profile plane normal to the D j fold axis (Nieuwland, 1980). (iv) Zircons from the gneiss are short with recognizable crystal faces but a rounded appearance, which is due to metamorphic corrosion rather than mechanical erosion (Fig. 2B). Metamorphic corrosion of zircon has been described by Silver (1969) and Jocelyn & Pidgeon (1974). It has produced some necked crystals in the gneisses, and the well-rounded short grains observed could be the result of the corrosive subdivision of original crystals as the final stage of the necking process. (v) Most of the zircons are brown and cloudy but a small proportion have clear purple overgrowths and there are a few separate purple crystals which are clear and euhedral. Parallel-growth and twinning are not uncommon, and most of the twins
T a b l e
I
Analytical results for zircon from quartzites exposed at Windmill Hill, sample 76-256, and Gabidine Springs, 77-1174 * Size f r a c t i o n
Weight
M Windmill
(mg)
Total
P
b
(ppm)
u
777
"
Pb/20ltPb
.
207
^
Pb/2 06Pb
206
'
+
" 'I 5 " / 7°
2 06
(ppm)
H i l l , 76-256
261
36.
275
395
'"86
327
°'9' Gabidine Springs, 77-1174
3 6 5
5 5 0
Pb/238U
207
Pb/23sU
—
—
<728 3220 5231
.25709* , .25160 ± 7 .246,4 • ,0
,6046 ± 30 .5818 ± 29 .5608 i 28
21.43= .0 20. ,3 t * . ,3.03*10
5791
.23655 1 89
.5403 ± 27
17.62 i '
""3
.25821 ± 21
.6026 , 30 .
21.45 ± 10
93
'I"
-100/+85
!'2.90 °
135
ftc/+7n -05/+/O
4.10
171
99c
9
.6*36+32
23- 1A ± )2
.
"
2 2 5
- 2 6 1 2 9 ± 12
.6193 ± 31
22.31 ±
223
3 0 5 1020 - 2 5 2 3 7 ± 10 refer to the last d i g i t s o n l y .
.55^2 ± 28
19.28 ± 10
7Q
6
—
9ftn 2 8 0
3
* Uncertainties
22k
shown
in this a n d o t h e r T a b l e s
1 2 7 0
1 7 0 7
•
•26079 ±
U
—
5
CRUSTAL EVOLUTION IN YILGARN BLOCK
T
10
'"i i—T-
15
163
20
COMBINED 0. 6
QUARTZITES
- %m
JD Q_
error overlap
207Pb/235U Fig. 3. Concordia diagram for detrital zircons from two orthoquart^ite units. The upper intersects are interpreted as two different (composite) ages for mixtures of zircons from the provenance of the original sandstone.
are purple. The twinned zircons are too fragile to have survived a sedimentary process. We conclude that the brown zircons of the Jimperding, Ringa and Noondeening gneisses crystallized from an igneous protolith, and that the purple overgrowths are either a late phase of magftiatic growth or a growth of zircon during later metamorphism. We agree with Prider (1944) that these particular gneisses are orthogneisses, but in view of the probably large amount of tectonic flattening, they need not have been originally sheet-like in form as proposed by Prider (1944). Age of the granitic protolith. The purple zircons were separated as much as possible from the predominant brown and cloudy grains by handpicking. However, there is no detectable agedifference between the two types (Table II, Fig. 4), nor is there any detectable difference between the upper intersection ages of the Jimperding and Ringa Orthogneisses which are 3243 + 190 - 190 Ma and 3211 + 43 - 35 respectively. Consequently we have combined the data for both gneisses to obtain 3246 + 70 - 50 Ma, and interpret it as the age of emplacement of the granitic protolith of the gneisses. This is comparable to
the zircon y^Pb age of 3280 ± 20 Ma for foliated homogeneous granite from the Mt Edgar Batholith in the Piibira Block (Pidgeon, 1978). Age of metamojphic growth of zircon. Zircons from the Noondeening gneiss were analysed for several reasons: (i) the zircon population contained a much larger fraction of clear purple crystals which could fre easily separated, (ii) the Noondeening gneiss contained a 30 x 50 m xenolith of zircon-rich cordierite-anthophyllite rock (Prider, 1940) whicli might be appreciably older than the gneiss protqlith, and (iii) the structural position of the Noondeening gneiss relative to the Windmill Hill quartzite could not be established conclusively, so it was conceivable that the gneiss itself might be basement to the sedimentary succession. The data for all the zircons both from the gneiss and the xenolith lie on the §§me chord to within experimental error (Fig. 4). The upper intercept age is 3180 + 60 - 40 Ma. This cannot be proved statistically to be younger than the 3250Ma age for the gneiss protolith. However, its value excludes the detection of possible basement ages using these zircons, and is obviously consistent with the concept of new zircon growth during later metamorphism.
164
D. A. NIEUWLAND & W. COMPSTON 5
10
15
207p 235 b /
20
U
Fig. 4. Concordia diagram for zircons from three gneiss units, giving a combined upper intersect age of 3246Ma interpreted as the age of crystallization of the igneous protolith of the gneisses. TABLE II
Analytical results for zircons from banded gneisses and an included xenolith (\im)
Size fraction
Jimperding Orthogneiss +100 +85
Weight (mg)
2.22 (purple)
-100+85
Total Pb (ppm)
u
(ppm)
206pb/20^pb
207pb/206pb
206pb/238u
207
•5277 + 26
17.07
Pb/235U
(76-264) 186
301
9580
.23464 ±
8
0.98
146
207
8465
.24380 ±
5
1.11
174
281
9066
.23268 ±
8
-85+60
3.48
210
342
14378
.23313 ±
2
-60
3.22
278
462
8925
.22771
4
±
.5131
± 29
.5310 + 27 .5274 ± 26 .5218 ± 26
9
19.49 ±,io 17.04 ± 9 16.95 i
9
16.38 ±
8
Ringa Orthogneiss -100+85
0.6V
684
2339
.22005 ±
8
-85+70
.5067
+ 28
15.38 ±
8
1.53 .
309
522
2652
.22049 ±
8
-70
.5157
+ 26
15.68 ±
0.62
354
623
1092
.21578 ± 16
.4881 + 25
U.52 ±
1
(magnetic)
-70
1.41
397
308
8
535
2811
+100
(purple)
1.04
229
280
10301
-80
(purple)
.24442 ± 19
. 5996 + 30
. 20,21 ± 10
0.94
427
611
1109
.2341
+100
(brown)
± 16
2.60
• 5497 + 28
17.74 f
• 5137 ± 26
15.87 ± f
Noondeening Orthognei ss . - 4 5 (purple) -100+85 +150
(purple)
(purple)
.21958 ±
6
.5023
312
511
1029
.22411 ± 73
1.00
348
431
417
.23687 ±
9
1.73
227
349
7569
.24274 ±
2.80
3
152
225
7444
.24624 ± 10
.5956
,5798
25
15.21 ±
8
H
(77-1133) .4952
±
25
.5768
±
29
19.30 ± 10
+ 30
20.22 ± 10
16.16.±
8
Cordierite-anthophyl1 ite xenolith (77-1126) -100+85
(purple)
-85+60
(purple)
4.47
223
331
3871
3.48
.24539 ±
5
243
367
76700
+100
(purple)
.24362 ±
8
4.93
242
• 5780 + 29
19.62 ± 10 19.41 + 10
365
4576
-60
(purple)
.24555 ± 21
19 .46 * 10'
3.91
269
• 5746 + 29
448
3935
.23945 ± 52
• 5235
* 26
• 17.29 ± 9
+
29
C R U S T A L E V O L U T I O N IN Y I L G A R N B L O C K
165
TABLE III
Analytical results for zircons from the migmatite melanosome, sample 77-1111 :
ract ion and Treatment
Weight - (mg)
Pb (ppm)
U (ppm)
, 2 Pb/ °Vb
206
207
Pb/ 2 0 6 Pb
206
Pb/ 238 U
207
Pb/ 2 35 U
Undi fferenti a ted Fracti ons:
m- + 100
8. 31
849
1795
1838
.1687
999 888
1240
+85-100
5. 50 4.,46 -
2366
#67:
2087
2432
.1639 . 1658
#69:
-60
3-• 57
1251
2907
1160
.1605
m:
+ 100
± +
6
.4398
21
8
.3862
± 19
8.73 ±.
±
5
9.09 ±
*
5
• 3975 ± 20 .3696 ± 18
8.18 ±
10.23 ±
5 4 5 4
Cpres and Rims from 65: #70:
purple cores, hot HF
1.• 73
458
686
995
,.17489 ± 12
.5841
± 29
14.08 ±
7
#71:
Tims, cold HF
9..08
331
878
2592
.17458 ± 17
.35704 ± 17
8.59 ±
4
#72-:
clear brown cores, cold HF
1,.48
392
1086
11327
• 17953 ±
.34823 ± 17
8.62 ±
4
#73:
cloudy brown cores, cold HF
5..05
434
1872
12487 '
.17815 +
4
.20953
10
5.15 ±
3
#74:
purple cores, ultrasonic only
1.• 19
234
485.
2011
.17488 ±
4
.4511
+ 23
10.88 ± 20
A discussion of the multiple possibilities for the origin of the xenolith arid the growth of its zircons is given by Nieuwland (1980). Their indicated age of 3180 Ma as well as their observed occurrence in cordierite and in retrograded hypersthene suggest that the zircons formed during retrograde amphibolite-facies metamorphism that succeeded the original granulite-facies event. The Post-gneiss Migmatite Large areas in the Toodyay district (Fig. 1) are migmatites that appear to be derived locally from older gneisses by metamorphic differentiation and partial melting. Zircon data described below indicate that such areas probably reflect remobilization processes during a thermal and metasomatic event much later than the main deformations and first granulite-facies metamorphism dated here at ca 3200 Ma. Zircons were extracted from portion of a 10 ,x 50m amphibole-rich block within hornblendebearing migmatite (Pioneer Quarry, 5 km WNW of Northam, Fig. 1), which comprises a neosome of alternating quartzofeldspathic and biotitehornblende-rich layers, and a palaeosome of hornblende, biotite, plagioclase and quartz. Almost all grains have a large near-white euhedral overgrowth on euhedral purple and, rarely, brown cores. The overgrowths may be twinned, and very small (1 to 2pm) zircons grow on larger crystals. Four zircon fractions (Table III) were analyzed without attempting to separate core from overgrowths. The data-points are well aligned with an upper concordia intercept of 2595 + 100 - 60 Ma, strikingly younger than zircons from the banded gneisses.
4
Various acid leaching treatments were employed in attempts to produce separate concentrates of cores and rims. They are detailed by Nieuwland (1980). Results (Table III) showed that both uranium and lead were leached from the zircons but the data were consistent with zeroage U/Pb disturbance (Fig. 5). We consider that only the purple cores cleaned from white overgrowth by ultrasonic treatment in water, No. 74 in Table III, give an undisturbed result. This is displaced from the untreated fraction No. 65 by a small but significant amount relative to experimental error, toward an older 207pb/206pb age. These data are consistent with the hypothesis modelled in Figure 5 that the purple cores were formed at the same time as the clear purple zircons in the Noondeening gneiss and its xenolith, but that nearly all the radiogenic Pb accumulated between 3180 Ma and ca 2595 Ma was lost from the cores during the younger event. We infer also that the white overgrowth zircon was formed somewhat later than 2595 Ma, as the latter is based on composite samples. The P b / 206pb age of the leached rim fraction, No. 71 at ca 2530 Ma, is the only estimate for the age of the overgrowth available from the zircon data. Independent estimates for the age of the neosome from Rb-Sr analyses are described below. Undeformed Intrusive Granitoids The Mortigup granodiorite intrudes metasediments in the northwestern part of the map area (Fig. 1) and the Monday Hill adamellite appears from the outcrop pattern to intrude migmatites in the central east. Apart from minor chloritization of biotite and saussuritization of plagioclase, both bodies seem unaltered. Both are compo207
D. A. NIEUWLAND & W. COMPSTON
166
10
15
0.6
0.5
_Q Q_
8
9 Pb/235U
207
Fig. 5. Concordia diagram for the migmatite zircons. The broken line through 2600Ma was generated by (present-day) leaching experiments in the laboratory. The Pb-Pb age for the rim fraction No. 71 at 2530 Ma is a minimum estimate for the age of new zircon that formed "during the second amphibolite-facies metamorphism.
nents of a large granitic batholith which occupies which had a white overgrowth. The data (Table 50% of the area of Figure 1 and, where relation- IV) exhibit the usual zircon age discordance ships are observable, always post-dates the pattern, and the upper intercept ages at 2677 + Jimperding Metamorphic Belt. 50 - 43 Ma for Mortigup and 2642 + 315 - 166 All the zircons extracted were euhedral, most Ma for Monday Hill are not distinguishable. If were purple, anel none contained cores except for crystallization of the two bodies is assumed to be the few grains from the Monday Hill adamellite synchronous, regression of their combined data TABLE I V
Fraction Mortigup
A nalytical resyfafor zircons from the Mortigup Granodiorite and Monday Hill Adamellite Pb Pb/ V (pig)
G r a n o d i o r iite
+ 146 -140+110 + 104,
cloudy
+1QA,
clear
+ 140,
non-magnetic
+ 100 -100+85
M*
U l 2.82
-110 Monday W i l l
(ppm)
Adamell i t e
U (ppm)
2 06pb/2 0Upb
2 07pb/2 0 6pb
206
23(
207
Pb/ 0 ?35
(77-1130)
377
+
916
159
.1779^ ± 96
.2800
274
645
252
.18081 ± 10
.3197
+ 16
7.971 ± 40
622
1582
168
.17886 ± 74
.2730
+
6.732 ± 3 4
329
749
240
.18304 ± 23
. 3264
+ 16
14
14
6.868 ± %
8.238 ± 41
2.86
424
981
149
.17995 ± 17
.2784
± 14
1.93
1234
4032
67
.17681 ± 18
.1438
+
7
3.505 * 18
336
851
317.8
.17572 ± 14
.3067
+ 15
7.432 ± 37
7.130 ± 36
(77-1086) 1.18 0.88
225
663
379.1
.17554 ± 76
-85+70
. 2694
± 13
6.592 ± 33
0.45
235
935
390.4
-70
.17583 t
59
. 2584
± 14
6.264 i 31
0.90
413
1153
290.0
.17272 ± 70
. 2628.
+
6.258 ± 31
13
167 by the small dispersion in Rb/Sr. Primary muscovite (Table V) gives a model age of 2577 ± 20 Ma, and muscovite from a post-deformation pegmatite (normally associated with the granitoids) gives 2570 ± 20 Ma. All of these ages are younger than the zircon age of 2670 ± 45 Ma but agree with the new growth of zircon in the migmatite at ca2530 Ma. Despite its wider range in Rb/Sr, the age of the Monday Hill adamellite at 2474 + 88Ma (Model III, Mclntyre et al., 1966) is even less precise than the Mortigup WHOLE-ROCK AND MINERAL Rb-Sr AGES granodiorite. This is due to its greater geological scatter (MSWD 89.0), which we attribute to variMethods ation in the initial Sr/ Sr owing to local All Rb and Sr determinations were made by assimilation of migmatite (Nieuwland, 1980) isotope dilution using a mixed Rb, Sr spike. and/or prolonged hydrothermal effects during Standard dissolution and cation exchange pro- cooling. Estimates for the original igneous age of cedures were used (Compston et al., 1977). Repli- the granitoids can be found using the procedures cate analyses of SRM 607 K-feldspar using this described by Cameron et al. (in press). The initial spike gave 24.17 ± 3 (a mean) for R b / S r , Sr/ Sr is modelled on values expected for the 1.2006 ± 2 for ?Sr/ Sr and an age of 1415 ± "bulk Earth", and isochron scatter and age reset2Ma (assuming 0.71 for initial Sr/ Sr). ting are attributed to local Sr isotope exchange between samples. The results are 2676 ± 20Ma Intrusive Gran itoids for the Mortigup granodiorite and 2635 Whole-rock data for the Mortigup granodiorite + 1080 - 130 Ma for the Monday Hill adamel(Table V) are almost perfectly aligned at 2500 ± lite, both being equal to the igneous zircon age to 138 Ma (MSWD 2.4), the precision being limited within experimental error. All the isotope data are therefore consistent with granitoid crystallization and emplacement at TABLE V 2670 Ma, followed by prolonged cooling under Rb-Sr analytical data for intrusive granitoids. Model ages are calculated in this and other Tables using amphibolite-facies metamorphic conditions until 0.7049 and 0.085 as the present-day Rb-Sr parameters a? 2530 Ma. of the "Bulk Earth". Migmatite Model Total Rb Sr/ Sr age Compositionally-differeftt slabs between 2-3 Samp1e Sr Rb/ (ppm) (Ma) (ppm) cm thick as well as large whole-rock samples were analyzed from both the neosome and the palaeo.-2MT 1 whol e rock 128.5 349-9 1 .063 .74179 ± 12 2607 some of the migmatite analyzed for zircon. The MT 2 whole rock . 212.8 381.2 0.925 .73750 ± 3 2681 data for the neosome are aligned to within experiHT 3 whole rock 127.2 342.7 1.075 .74272 ± 11 2640 mental error at 2512 ± 115 Ma, in agreement MT 4 whole rock 116.4 358.3 0.941 .73806 ± 14 2677 with the internal Rb-Sr alignments of the intrue rock 116.8 356.8 O.947 .73822 ± 3 2671 sive granitoids and with the migmatite zircon age. WT 6 whol M whole rock 120.5 358.0 0.975 .73957 ± 6 2691 The close alignment and the particular age value MT 7 whole rock 128.6 346.6 1.075 .74308 ± 2 2665 signify that the neosome as such was first formed MT 8 Who 1 erock 126.3 355.8 1.028 .74150± 2 2682 during the ca 2530 event, and its high initial 2634 . 7 4149 ± 55 MT 9 who 1 erock 134.0 371.2 1,045 Sr/ Sr, 0.711 ± 4, indicates that it was deMT 10 -whole rock 132.6 325-8 1.179 .74679± 5 2646 77-1130 muscovi te 1192.4 33 - 9 161.16 6.7127 ± 3 2577 rived from an older source. The Rb-Sr data for the palaeosome present a Monday Hill adamel1ite (77-•1086) MH whole rock ' 255-Q • -8374 10 1:8 m t J79 2449 contrasting picture. Over the ~ 25 cm scale of MH j$ whole roek 25 K 7 81.7 9-17 1.02853± 38 2465 sampling, isotopic equilibrium at 2530 Ma was MH 2A whole rock 220.8 723.5 0.88 .73468 ± 7 2590 not established. One individual model age is ca MH 2B whole roek 226.7 738.6 0.89 .73437 ± 11 2532 4150Ma (sample No. la) and two others exceed MH 3 whole rock 122.9 320.6 1.11 .74759± 7 2874 orthogneiss age of 3250 Ma. Such excessive MH 4 who1 eroek 236.7 131.7 5.28 .89391 ± 31 2517 the ages must be expected if the palaeosome, MH 5 whole rock 231.3 70.6 9.78 1.05973± 23 2531 model in part constitutes a residue after MH 6 whole rock 213.7 172.3 3.63 .83774± 6, 2591 extractionorof whole, the neosome as a low-melting-point Int rusive peqmatite , Ringa Area (77-1160) fraction. No precise age information can be obmuscovite 4030.6 42.6 27133 1008.9 ±5-6 2569 tained from the melanosome Rb-Sr results. C R U S T A L E V O L U T I O N IN Y I L G A R N B L O C K
gives 2670 ± 45 Ma for the joint age. This result is detectably older than the growth of zircon in the migmatite, indicating that the migmatization process occurred, or continued, after crystallization of zircon in the granitoids. Some effect of the process might be expected therefore in the latter. It will be shown now that the Rb-Sr whole-rock data in both the migmatite and granitoids register ca 2500 Ma rather than 2670 Ma.
87
85
84
87
8
86
87
86
87
86
86
87
87
8
86
7
86
-r
5
86
168
D. A. N I E U W L A N D & W. TABLE VI
Rb-Sr analytical data for migmatite Sample
Rb (ppm)
Sr (ppm)
87
86
Rb/ Sr
87Sr/86:
Sr
age (Ma)-
2828
Neosome (77 -112?) 2.14
.78910 ± 19
NE0 IB
99-9
134.9
2.15
.78939 ±
NE0 2
124.2
130.1
2.79
.81297 ± 56
2759
.77449 ± 42
2900
.80448 ±
2756
NE0 1A
135.2
99.6
NEO 3
81.3
135.5
1.74
NEO 4
113.4
129.7
2.58
6
5
2824
NEO 5
123.8
132.1
2.72
.81008 ± 12
2756
NEO 6
110.0
131.6
2.44
.79978 ± 11
2782
4219
Palaeosome (77-1111) 86.6
201.0
1.25
.77682 ±
MEL 2
109.7
173.9
1.83
.76935 ± 26
MEL 3
107.8'
155.0
2.02
.78624 ±
9
2900
MEL 4
99.6
191.6
1.51
.77667 ± 20
3460
MEL 5
117.8
220.8
1.55
.77631 ±
8
3352
MEL 1
MEL 6
77.8
189.0
9
2554
2938
1.19
.75198 ±
9
MEL 7
82.2
162.4
1.47
.75867 ±
8
2682
MEL 8
79.3
170.7
1.35
.75419 ± 12
2692
MEL 9
61.6
216.5
0.82
.73667 ± 31
2980
Muscovites and Fuchsites from the Metasediments From studies of the textures and mineral assemblages, Nieuwland (1980) concludes that all the minerals in the schists and quartzites underwent static recrystallization during a middleamphibolite-facies metamorphism (600-650 °C). This metamorphism was subsequent to the three deformational phases recognized in the microstructures,* but no constraint on the time interval could be given. It may be very short so that the recrystallization was part of a single combined deformational and metamorphic event, or it
COMPSTON
could be much later during a subsequent metamorphic event. Using modelled " n o r m a l " initial 87 Sr/86s rj the ages of all the micas are nearly the same (Table VII), as expected for a regional metamorphism. However, the age value itself at 2560 Ma is younger than the emplacement of the postdeformational granitoids at 2670 Ma, and equal within error to the growth of new zircon in the migmatite. Thus the amphibolite-facies metamorphism ended late in the Archaean history of the area. It was therefore a different event, most probably, rather than a retrograde phase of the early granulite-facies metamorphism dated at 3180 Ma by zircons from the cordierite-anthophyllite rock. Concentrates of (recrystallized) quartz which contain small inclusions of microcline were also analyzed from two fuchsite bands to clarify the initial 87Sr/ 86sr of the fuchsite (Table VII). If the fuchsite and " q u a r t z " from single bands achieved Sr isotopic equilibrium during recrystallization, the two quartz-fuchsite tie-lines which average 2568 Ma measure the age of that event. The indicated values for initial 87Sr/ 86Sr are 0.71 and 0.75, both of which we regard as low for the Rb-rich environments. It follows that none of the micas recrystallized in a closed local environment. Instead they appear to have exchanged their radiogenic Sr with a large external reservoir of common Sr during the regional metamorphism. The Ringa and Jimperding
Gneisses
Multiple samples for Rb-Sr analysis were collected at sites over a profile length of ca 100m and near the zircon sampling sites. In addition* light and dark bands were taken from the Jimperding gneiss. The Noondeening gneiss was not TABLE V I I analyzed owing to its more pronounced surface Rb-Sr analytical data for recrystallized micas from schist (sample 77-1161) and quartzites, and microcline- weathering. Final sample selection was based on bearing recrystallized quartz fractions thin-section examination, including such criteria as the degree of chloritization of biotite, saussuriTotal Model Rb Samp Ie Sr Rb/ Sr S r / 86 Sr age tization of plagioclase and the abundance of (ppm) (ppm) (Ma) epidote. However, no samples were free of these 77- 1154 fuchs i te A 610 19.4 136 .2 5.8103 + 37 2593 effects. 77- 1154 fuchsi te B 616 20.0 132..1 5.6632 ± 36 2597 77- 1154 quartz 9-8 24..3 1-3 1.6054 ± 12 2571 All data are plotted in the Sr evolution dia77- 1157 fuchs i te 572 121.• 9 5-1871 + 36 19.5 2545 gram, Figure 6, including published results for 77- 1157 quartz 8.2 26.7 9.• 7 1.0712 ± 9 2633 five samples of Arriens (1971, "Toodyay 77- 1161 muscovi te 280 20.0 2.4656 ± 5 47..2 2584 77- 1164 fuchsite A 30.8 gneisses", Table I) which are from a 500m 535 61..0 2.9284 + 14 2525 77- 1164 fuchs i te B 510 29.6 60..4 2.8878 ± 8 2504 traverse through the "Upper Gneiss" (Prider, 77- 1170 fuchs i te A 2K 1 363 60.9 2.9824 + 84 2589 1944). In the absence of metamorphic or weather77- 1170 fuchsite B 376 20.4 66. 3 3.1930 ± 8 2598 77- 1254 fuchs i te ing effects, all points should fit closely on a 3250 595 18.5 140. 7 5.9576 ± 39 2583 Ma isochron based on the zircon results. Instead, 87
h ng
t 2 f0l,atl0n
(
Nieuwl
86
87
rnetamorphism occurred during D 2 . This is evident by the concentration ofmusco^ ! T o n C n ? o 2 c r e n u l a t l o n s > w h e r e they recrystallized with the 001 plane perpendicular to the
* n d , 1980). However no isotopic record of this earlier event can be expected
169
C R U S T A L E V O L U T I O N IN Y I L G A R N B L O C K
TOODYAY ORTHOGNEISSES
•85
/
87
Sr/«*Sr
3225MQ •80
-
/
•70
j f
BULK EARTH
|
#
V
A
2500MO
v
• /
•75
/
JIMPERDING V RINGA • A
87
Rb/wSr
I
1
(This work) (Aniens 1971)
I
Fig. 6, Sr evolution diagram for whole-rock samples of orthogneiss. All the Jimperding and most of the Ringa samples have lost Sr subsequent to original crystallization at ca 3225 Ma. Isochrons for undisturbed rocks derived from the average unfractionated mantle pass through the present-day "Bulk Earth" point (Cameron et al. in press). 87
t
TABLE VIII Figure 6 shows that most of them including all the Jimperding gneiss samples, are displaced to Rb-Sr analytical data for orthogneisses from Ringa and Jimperding areas the RHS of this line (young model ages), and that none of them lie to the LHS of the line beyond experimental error. The importance of the zircon U-Pb age as a basis for understanding the whole-rock Rb-Sr behaviour cannot be over-emphasized. The above observations can be interpreted only by the mechanism of variable loss of either total Sr or radiogenic Sr, relative to Rb, from the wholerock samples. This is commonly found for wholerock samples of recrystallized acid volcanics (e.g. Page, 1978). It excludes the age-resetting mechanism observed elsewhere, of local Sr isotope equilibration between whole-rock samples (Cameron et al., in press), which must produce samples having both old and young model ages in about equal abundance. No regression method can be validly applied to data affected by variable Sr loss, and we attribute no significance for example, to the moderately good internal alignment of the Jimperding samples at 2100 ± 300 Ma. In the light of the zircon results, we may assess the Rb-Sr data validly by examining the distribution of Rb-Sr model ages. As before, the initial Sr/ Sr may be taken as 0.7009, corresponding 86
.82271 ± 22 .82129 ± 10
3211 3174
.80937 ± 28 .80959 ± 21
2480 2476
RN 1A RN I B
162-0 163-6
181.0 183.3
2.61 2.61
RN 2A RN 2B
152.4 155-1
148.3 150.5
3-00 3.01
RN 3 A RN 3B
113.2 111.0
261,1 256.3
1.26 1.26
.75750 ± -75675 ±
9 6
3084 3041
RN 4
160.2
163-3
2.86
.81884 ± 14
2834
RN 5A RN 5B
122.4 122.7
150.0 150.7
2.38 2.37
.80887 ± 47 .81002 $ 2
3120 3167
RN 6
146.4
150.2
2.85
.82767 ±
2
205?
87
Rb/86Sr
87
gneiss
-
RN 7
154.9
133.0
3.41
,83180 ±
2
2638
RN 8
188.6
164,4
3.35
.83836 ±
8
2821
RN 10
162.1
251,7
1.88
•79059 ± 15
3284
Jimperding g n e i s s
87
Model age (Ma)
Tota 1 Sr (ppm)
Ringa
87
S r / 8 6 S >r
Rb (ppm)
Samp1e
(76-258 t o 76-264)
5
2904
127-4
97-4
3.83
.84386 ±
3
2566
JM 260
133-3
102.2
3-82
.84409 ±
3
2577
JM 261
137.7
157.5
2.55
.80571 ± 11
2833
JM 262
140.0
141.7
2.89
.82134 ±
2
2864
JM 263A
159.6
144.1
3.24
.82804 ±
5
JM 263B
134.3
104.5
3.76
.84472 ±
4
2630
JM 264
133.4
144.9
2.69
.81611 ±
2
2944
JM 258
150-4
JM 259
207,4
2.10
.78972 ±
2696
170
D. A . N I E U W L A N D & W . C O M P S T O N TABLE IX
. Model Rb-Sr ages for the Ringa orthogneiss from (1) Arriens (1971), and (2) Libby & de Laeter (1979), using 0.7009for the initial Sr/^Sr Sample No.
36950
87
Rb/86Sr
87
Sr/86Sr
Model age (Ma)
.92905
2847
(2)
5.53
GA4136 (1)
2.016
.7957
3236
GAAl37
"
2.22V
.8021
3138
GA4138
"
4.087
• 8791
3005
GA4139
"
2.022
.79^5
3187
GA4140
"
.587
. 7256
2903
to the contemporary bulk Earth Sr/ Sr at 3250 Ma. The expected precision in model age for these samples is about 1.5%, which is borne put also by the replicate data for the Ringa gneisses (Table VIII). Thus two model ages that differ by more than 4.5% are probably different and one or both of the samples were affepted by Sr loss. The six oldest model ages in Tables VIII and IX combined cover a range of 4.7% and have a mean value of 3197 ± 56 Ma (95% confidence limits for the mean). If the two youngest ages within this group are discarded because of Sr loss, the mean becomes 3225 ± 72Ma. The latter can be regarded as a reliable minimum estimate for the original igneous age, and is, incidentally, not distinguishable from the zircon result at 3250 Ma. We presume that the loss of Sr occurred during the middle-amphibolite-facies metamorphism that recrystallized the muscovites and generated the migmatite at ca2530 to 2560 Ma. The young87
86
87
87
est gneiss model ages at ca 2500 Ma are consistent with this hypothesis. The fact that a nett loss occurred on a regional scale suggests that exchange of Sr and/or loss of total Sr took place via a pervasive fluid medium, as inferred also for the recrystallized muscovites. This possibility has vital implications for the utility of Sr/ 86$r as a fingerprint for older source material. Had the exchange gone to completion, the 3250 Ma Toodyay gneisses would exhibit a well-fitted isochron at ca 2500 Ma which had very little if any indication of an elevated initial Sr/ Sr. Arguments using the initial Sr/ Sr and regionally closed-system anatexis have been employed in the Greenland Archaean to distinguish the genesis of the Nuk and Amftsoq gneisses (e.g. Moorbath, 1975). These arguments would not apply to the ageresetting of the Toodyay gneisses, on the above evidence. We do not exclude the possibility that still later metamorphic processes including deep weathering may have also affected the gneiss and other whole-rock Rb-Sr samples. For example, Libby & de Laeter (1979) observe an alignment at car 1500 tyla between whole-rock samples in the granitoid at Paid Hill (Fig. 1). The youngest model age given by the Bald Hill data is ca 2420 Ma, which is hardly any later than ca 2500 Ma measured here. Nevertheless, the hydrothermal alteration associated with the emplacement of postgranitoid mafic dykes which increase in abundance toward the Darling Fault may be an important factor. Discussion of published gneiss Rb-Sr ages in the light of the zircon U-Pb results will be given elsewhere by Pidgeon and Compston. 87
87
87
87
86
86
TABLE X
Archaean geological history: summary of conclusions Event 1.
Age
Crystallization of the unseen source rock of the quartzites
2.
Deposition of sand and clay in a shelf environment subject to tidal currents
3.
Intrusion of microcline granites, the protoliths of the orthogneisses
k.
Deformation and metamorphism of the sediments and intruded granites to granulite facies, metamorphic growth of zircon
5.
Retrograde amphibolite facies metamorphism, loss from metasediments
6.
Intrusion of granitoids with igneous crystallization of zircon, during
7-
inferred
87
Sr
A prolonged younger amphibolite facies metamorphism accompanied by recrystal1ization of micas, migmatization, growth of new zircon and localized 8 7 S r loss from granitoids and gneisses
(Ma)
> 33^0 ca 3300
32.50
Method zircon U-Pb (inferred)
zircon U-Pb, whole rock Rb-Sr
ca 3180
zircon U-Pb
ca 3000
whole rock Rb-Sr
2670
87 6
(Arriens,
1971)
,/-7n
zircon U-Pb, w h o l e rock Rb-Sr (modelled initial
2560
zircon U-Pb, w h o l e - r o c k and muscovite Rb-Sr
to 2500
Sr/8
Sr)
CRUSTAL EVOLUTION IN YILGARN BLOCK 171 CONCLUSIONS ACKNOWLEDGMENTS The zircon U-Pb results have provided the first We thank Dr M. J. Rickard for help with the reliable time frame for the Toodyay-Northam structural geology, Dr R. T. Pidgeon with the area. Table X summarizes our conclusions for its zircon morphology, and Professor R. T. Prider Archaean geological history. The extent to which and Dr S. Wilde for guidance in the field. Drs the proposed history is applicable to the re- Pidgeon, Wilde and de Laeter kindly reviewed the mainder of the Jimperding Metamorphic Belt text. should be tested by further isotopic studies. REFERENCES
ARRIENS, P. A., 1971: Archaean geochronology of Australia. Spec. Pubis geol. Soc. Aust., 3, 11-23. BLATT, H., MIDDLETON, G., & MURRAY, R., 1972: The
Origin of Sedimentary Rocks. Prentice-Hall, New Jersey.
CAMERON, M., COLLERSON, K. D., COMPSTON, W . , & MORTON, R., in press: The statistical analysis and
interpretation of imperfectly-fitted Rb-Sr isochrons from poly metamorphic terrains. Geochim. cosmochim. Acta.
COMPSTON, W . , FOSTER, J. J., & GRAY, C. M., 1977:
Rb-Sr systematics in clasts and aphanites from con-
sortium breccia 73215. Proc. 8th lunar Sci. Conf, 2, 2525-2549. JOCELYN, J., & PIDGEON, R. T., 1974: Examples of twinning and parallel-growth in zircons from Precambrian granites and gneisses. Min. Mag., 39,
587-594. JOHNSTONE, M. H., 1952: The geology of the Hamersley Siding area. J. R. Soc. West. Aust., 36, 45-75. KROGH, T. E., 1973: A low-contamination method for
hydrothermal decomposition of zircon and extraction of U and Pb for isotopic age determinations. Geochim. cosmochim. Acta, 37, 485-494. LIBBY, W. G., & DE LAETER, J. R., 1979: Biotite dates and cooling history at the western margin of the Yilgarn Block. Ann. Rep. geol. Surv. West. Aust.
for 1978, 79-87.. MCWHAE, J. R. H., 1946: The geology and physio-
graphy of the Lawns wood area. J. R. Soc. West. Aust., 32, 49-74.
MCINTYRE, G. A., BROOKS, C., COMPSTON, W . , & TUREK, A., 1966: The statistical assessment of RbSr isochrons. J. geophys. Res., 71, 5459-5468.
Geological interpretation of whole rock isochron dates from high grade gneiss terrains. Nature, Lond., 255, 391.
MOORBATH, S., 1975:
NIEUWLAND, D. A., 1980: Structural geology and geo-
chronology of the Toodyay district, Western Australia. Ph.D. Thesis, - Aust. Nat. Univ. [unpublished]. PAGE, R. W., 1978: Response of U-Pb zircon and Rb-Sr total-rock and mineral systems to low-grade regional metamorphism, in Proterozoic igneous rocks, Mount Isa, Australia. J. geol. Soc. Aust., 25, 141-164. PIDGEON, R. T., 1969: Zircon U-Pb ages from the Galway granite and the Dalradian, Connemara, Ireland. Scott. J. Geol, 5, 375-392. , 1978: Geochronological investigation of granite batholiths of the Archaean granite- greenstone terrain of the Pilbara Block, Western Australia (Abstract); in Smith, I. E. M., & Williams, J. G. (Eds) Proceedings of the 1978 Archaean Geochemical Conference, 360-362. Univ. Toronto, Ontario. PRIDER, R. T., 1940: Cordierite-anthophyllite rocks associated with spinel-hypersthenites from Toodyay, Western Australia. Geol. Mag., 77, 364-382. , 1944: The geology and petrology of part of the Toodyay district, Western Australia. J. R. Soc. West. Aust., 28, 83-137. SILVER, L. T., 1969: Geochronologic investigation of the Anorthosite Complex, Adirondack Mts, New York; in Isachsen, Y. M. (Ed.) The Origin of Anorthosite and Related Rocks Mem. N. Y. State Mus. Sci. Service, 18, 233-251. STEPHENSON, N. C. N., 1970: The high-grade metamorphic and associated igneous rocks at Mt Bakewell, near York, Western Australia. J. R. Soc. West. Aust., 53, 81-94. WILDE, S.. A., & Low, G. H., 1978: Perth, Western Australia 1: 250000 Geological Series. Explan. Notes geol. Surv. West. Aust., SH50-14. WILLIAMS, I. S., 1977: The Berridale Batholith: a Lead and Strontium Isotope Study of its Age and Origin. Ph.D. Thesis, Aust. Nat. Univ. [unpublished].
Rb-Sr GEOCHRONOLOGY OF GRANITIC ROCKS FROM THE DIEMALS AREA, CENTRAL YILGARN BLOCK, WESTERN AUSTRALIA H. J. Chapman, 1 M. J. Bickle,1 J. R. de Laeter, 2 L. F. Bettenay, 1 D. I. Groves, 1 L. S. Andersen, 3 R. A. Binns, 4 and M. Gorton 5 1
Department of Geology, University of Western Australia, Nedlands, Western Australia 6009
2
Department of Physics, Western Australian Institute of Technology, South Bentley, Western Australia 6102 3 Institute of Petrology, University of Copenhagen, Denmark 4 CSIRO Division of Mineralogy, North Ryde, New South Wales 2113 5 Department of Geology, University of Toronto, Ontario, Canada ABSTRACT Rb-Sr geochronology of Archaean granitic rocks from the central Yilgarn Block yields ages which may be related to sequences of deformational and intrusive events. Banded gneisses, possible candidates for basement to the greenstone sequences, give ages within error of available estimates for the greenstones. Initial 8 7 Sr/ 8 6 Sr ratios for the gneiss isochrons preclude their formation more than 100-200m.y. before deposition of the greenstones. The ages of the major deformational events are bracketed by the syn- or pre-tectonic plutons at Rainy Rocks dated at 2612* ± 37 m.y., 8 7 S r / 8 6 S r initial ratio (IR) = 0.7014 ± .0011 and the post-tectonic pluton at Milky Soak dated at 2557 ± 49 m.y., IR = 0.7099 ± .0040. The oldest mineral ages (2537 ± 17 m.y.) record cooling from the associated metamorphism. A late, undeformed leuco-adamellite is dated at 2467 ± 36m.y., IR = 0.7186 ± 0.12. There is local evidence for rehomogenisation of Rb-Sr isotope systematics several hundred million years after intrusion of some plutons. Pegmatites and adjacent whole-rock samples at Pigeon Rocks give an age of 2302 ± 58 m.y., IR = 0.966 ± .064 whereas Pigeon Rocks gives a whole-rock age of 2550 ± 53 m.y., IR = 0.7151 ± .0072. Consideration of average 8 7 Sr/ 8 6 Sr evolution with time suggests that the pegmatites may have been intruded at ca 2550 m.y. This evidence for resetting of Rb-Sr isotope systematics raises the possibility that some of the ages with apparent high initial * 87 Sr/ 86 Sr ratios are reset and this has important implications for interpretations of Archaean Rb-Sr geochronology.
INTRODUCTION This paper reports results of a detailed study of Rb-Sr isotopic systematics in Archaean granitic rocks from the central Yilgarn Block in Western Australia. The study was formulated to investigate the genesis and evolution of the granitic rocks to test a model of the structural and metamorphic evolution of the Yilgarn Block first proposed by Binns et al. (1976). However, initial results (Andersen et al., 1976) indicated possible resetting of Rb-Sr ages over an interval of several hundred million years after intrusion of the plutonic rocks. The possibility of such resetting may severely limit the interpretation of Rb-Sr * XRb = 1.42 x 10
11
yr
\ errors quoted at lcr level.
Spec. Pubis geol. Soc. Aust., 7 (1981)
ages in Archaean terrains. Not only will ages be too young but apparent initial 8 7 Sr/ 8 6 Sr ratios will be anomalously high. These initial 8 7 Sr/ 8 6Sr ratios provide evidence about the processes of crustal evolution in that high initial ratios are taken to imply derivation by remelting of older crustal material. The proportion of such reworked older crust is a critical parameter in models of Archaean crustal evolution. Hence in this project an attempt has also been made to ascertain the extent and cause of such resetting, The model proposed for the evolution of the Yilgarn Block by Binns et al. (1976), Archibald & Bettenay (1977), Archibald et al. (1978) and Gee (1979a, b), among others* postulates an early
H. J. CHAPMAN, ET
174
GREENSTONE BELTS Younger Metasedimentary Sequence \?p°oi Low strain metamorphism Older Voleano/Metasedimentary Sequence Low strain metamorphism High strain metamorphism B I F marker beds 0
10
20 km Fig. 1.
AL.
GRANITOIDS -;-»-;-,H Fractionated leucoadamellites rZZ\ Heterogeneous postkinematic granitoids Postkinematic granitoid plutons 1 Synkinematic granitoids IH1II11111 Foliated hornblende tonalites STRUCTURES --J--- Major fold Fault
G
Unconformity Sample site
Map of the Diemals area, central Yilgarn Block.
sialic crust, now at least partly represented by banded gneisses, on which greenstone sequences were deposited. Subsequent early folding was followed by a widespread deformational and metamorphic event. Marginal zones to greenstone belts were metamorphosed to high grade and intruded by granitic rocks, and both greenstones and granitic rocks were subjected to high strain. Greenstones in areas of lower grade were subjected to less internal strain with pre-existing structures often well preserved. During this metamorphic event, the underlying sialic crust was heated above minimum melt conditions and the widespread granitic rocks are inferred to be de-
rived by partial melting of the banded gneisses. This model is reviewed in detail by Archibald et al. (1981). PREVIOUS WORK Previous geochronological studies in the Yilgarn Block are summarized by de Laeter et al. (1981). Most of the studies on the central and eastern Yilgarn areas have given ages in the region of 2700 to 2500 m.y. (XRb ^ 1.42 x 1 0 " n y r _ 1 ) and only the more recent studies attempt to determine a relative chronology of events. The results of Roddick et at. (1976) and Cooper et al. (1978), constrained by
Rb-Sr GEOCHRONOLOGY, YILGARN BLOCK
the field studies of Piatt et al. (1978) on the Lawlers-Agnew-Perseverance region some 250 km northeast of the Diemals area, are the most pertinent to this study and are further discussed below. REGIONAL GEOLOGICAL SETTING OF THE DIEMALS-EVANSTON AREA The Diemals-Evanston area was chosen for study because of relatively extensive exposure combined with availability of detailed field studies including work by L. S. Andersen (summarized in Andersen et al., 1976), BSc Honours students (Porter, 1971; King, 1974; Walker, 1974) and unpublished maps by undergraduate students of the University of Western Australia. The area (Fig. 1) lies within the Southern Cross Province, and is a typical granite-greenstone terrain comprising rather small, narrow, arcuate greenstone belts intruded by ellipsoidal to wedgeshaped granitoids that include discontinuous bodies of gneiss. The greenstone belts are dominated by komatiitic to tholeiitic metabasalts with inter layered metasedimentary sequences, including prominent BIF units, and acid to felsic metavolcanic rocks at the top of the exposed pile (e.g. Hallberg et al., 1976): relative to other Eastern Goldfields greenstone sequences, ultramafic flow sequences are rare. This composite sequence is overlain unconformably in the western part of the Diemals belt by meta-arkoses, metaconglomerates and micaceous quartzites (Fig. 1). The maximum thickness of the total sequence is ca 10km. The greenstones have undergone Regional lowpressure, low- to high-grade metamorphism (e.g. Winkler, 1974) with internal domains of lower grade, low-strain metamorphism contrasting with marginal commonly linear, higher grade, highstrain zones (Fig. 1). At least three deformation events are inferred, largely by reference to mesoscopic structures within the BIF units. Intrafolial, isoclinal folding (¥{) was followed by open to tight asymmetric folds in low-strain areas and tight to isoclinal folds with associated strong linear fabrics in high-strain areas: this second folding event (F2) was possibly related to granitoid diapirism. Subsequent open to tight folding (F3) was coaxial with F 2 and composite F 2 -F 3 folding appears to be the dominant control of the present arcuate form of the greenstone belts. Metamorphic recrystallization outlasted deformation so that most rocks in high-strain zones are now granofelses. Granitic rocks underlie ca 75% of the DiemalsEvanston area but exposure is sparse; synkinematic and post-kinematic intrusive granites (see Fig. 2} dominate outcrop. Banded gneisses,
175
j1 f 1 -Granodiorites x
Mondie
Rocks
Pigeon
Rocks
East
Pigeon
Rocks
North Diemals - Mt. Elvire
Fig. 2.
Milky
Soak
Rainy
Rocks
Kellys
Rocks
Johnson
Rocks
Compositions of the granitic plutons sampled by this study from Bettenay (1977).
similar to those described elsewhere in the Eastern Goldfields Province, occur largely as discontinuous blocks cut by intrusive granites and pegmatites (e.g. Cockatoo Rocks) or as widespread smaller xenoliths in the intrusive granites. Synkinematic granites occupy a prominent NEtrending zone that includes outcrops at and west of Pigeon Rocks and the Rainy Rocks belt between Broadbents and Evanston, and also typify the eastern margins of the Manning Range and Evanston greenstone belts. The Rainy Rocks belt comprises biotite adamellite-granodiorites with lesser leuco-adamellites typified by extreme subhorizontal mineral lineation subparallel in trend to the greenstone margin and to fabrics in marginal greenstones. These lineated granitoids are cut by deformed, steeply dipping pegmatite and aplite dykes. The Pigeon Rocks body is also a biotite adamellite cut by aplites and pegmatites, but the adamellite is generally only weakly foliated and lineated. Post-kinematic granites are typically medium to coarse grained and variably porphyritic, and range mainly from granodioritic to leuco-adamellitic in composition. On the contact with synkinematic granites, relationships are complex, but in some outcrops (e.g. parts of Milky Soak) weakly deformed, porphyritic post-kinematic adamellites demonstrably cross-cut, and contain xenoliths of, lineated adamellite-granodiorites of Rainy Rocks type. North and northeast of Diemals, the post-kinematic granites are essentially undeformed, but in the southeastern part of the area (e.g. adjacent to Manning Range), the granitoids have a pervasive cataclastic deformation expressed as patchy, streaky, mineral lineations, chloritic shear-zones and/or mylonites. In this area, large xenoliths of hornblende-bearing tonalite that have a pre-cataclasis fabric occur within the post-kinematic granites. The Johnson Rocks
176
H . J . C H A P M A N , ET
AL.
TABLE I
Rb, Sr and 87Sr/86Sr analyses of samples Sample 2 No.
87
Rock Type
Rb/
86
Sr
87
Sr/ Sr
2x Standard 1 error
86
Pigeon Rocks - Whole Rock HC1 HC2 HC3 HC4 HC5 HC6 HC7 HC8 HC9 HC10
Granodiorite/adamel1ite
3.65 4.02 3-50 3-97 2.44 2.47 2.74 3.75 3.59 2.47
11.03 12.18 10.53 12.00 7.25 7.33 8.17 11.30 10.79 7.33
1.12938 1.16056 1.11048 1,15978 0.98320 0.98382 1.01570 1.12383 1.10599 0.98672
.0002 .0002 .0002 .0001 .0002 .0002 .0001 .0001 .0001 .0001
19-0
19.7 20.9 23.2
68.54 74.40 79-99 71.84 77.20 87.93
3.23064 3.43388 3.62983 3.36762 3.53621 3.87682
.0004 .0003 .0010 .0005 .0005 .0003
0.255 0.268 0.290 0.249 0.240 0.280 0.21,7 0.174 0.267 0.217 0.209 0.168
0.740 0.778 0.841 0.722 0.696 0.812 0.629 0.504 0.774 0.629 0.606 0.487
0.73076 0.73242 0.73320 0.72958 0.72694 0.73130 0.72718 0.72237 0.73199 0.72545 0.72552 0.72047
.0001 .0001 .00019 .00018 .00010 .0001.2 .00017 .00020 . .00015 .00002 .00009 .00010
~
1.234 1.326 0.656 0.415
3.614 3.888 1.91 r j .205
0.82819 0.84603 0.777026 0.74455
.00010 .00014 .00010 .00020
266 310 290 .345 305 305 285 330 308 305
73 77 83 87 125 124 104 88 86 123
390 394 375 364 396 450
20.6 19-4 17.4 18.6 19-1 19.4
Pegmatites and Samples adjacent to pegmatites at Pigeon Rocks HC11C HC11D HCllGb HCl.lGc HCllGd
Granodiorite/adamel1ite
HC11H
Pegmatite
"
Banded Gneiss from Cockatoo Rocks HC22 Banded gneiss HC23 HC25 HC26 HC27 HC29 HC30 HC32 HC33 HC34 HC35 HC36
100 116 138 113 113 142 109 78 120 107 94 88
395 434 475 455 470 510 500 447 450 495 450 .525
20.4
21 .5
1 on i sochron HC23 HC24 HC28 HC 31
Tonali t i c vein G ran i t i c ve i n Gneiss adjacent to pegma-tjte Banded gneiss
Compositional Bands from Sample HC22, Cockatoo Rocks HC22 HC22A HC22B HC22C HC22D HC22E HC22F HC22G HC22H HC22I HC22J HC22K HC22L HC22H
375 225 206
305 170 316
"
Whole-rock Band wi thin
100 143 56 148 56 125 152 170 56 107 115 100 90 116
395 442 458 420 3&5 425 412 405 425 385 305 340 320 285
0.255 0.325 0.123 0.354 0.146' 0.297 0.370 0.422 0.133 0.278 0.376 0.293 0.280 0.409
. 0.740 0.943 0.356 1.028 0.423 0.862 1.075 1.226 0.385 0.801 1.092 0.850 0.812 1 . 1 89
0.73076 0.73564 0.71553 0.73874 0.71874 0.73330 0.74357 0.74733 0.71732 0.73307 0.74370 0.73508 0.73299 0.74862
.0001 .0001 . OOO'l .0001 . 0001 .0002 .0002 .0001 .0001 .0001 . .0002 .0002 .000*1 , .0001 <
Leuco-adamel1i te
380 375 373 460 378 430 362 375
49 51 56 52 58 55 61 60
7-77 7.32 6.69 8.87 6.50 7.81 5.95 6 ; 25
24.4 22.8 20.8 28.2 20.2 24.6 18.3 19.3
1.58441 1.52571 1.46134 1.72979 1.43863 1.59744 1.37322 1.40836
.00029 . 0602 .0002 .0003 .00023 .00027 .00022 .•00013
T
212 195 200 200 210 198 206 200 200 202
123 134 U6 129 126 121 125 128 128 128
.726 .456 • 717 .550 .666 .632 .646 .553 .563 .572 • 75,0
5.09 4.28 5.06 4.57 4.91 4.80 4.85 4.57 4.60 4.62 5.17
0.89338 0.86627 0.89600 0.875^4 0.88947 0.86833 0.88421 0.87027 0.87767 0.87765 0.89890
.00019 .00016 .0001,6 , .000.2 .00014 .00018 .0002 .0001:5 . 0001 .0001 .001
203 217 232 211 183 218
96 132 115 123 110 94
2.118 1.646 2.017 1.716 1.666 2.322
6.267 4.846 5.964 • 5.057 4.906 6.886
0.93951 0.88821 0.93246 0.89784 0.88945 0.96354
.00017 .00021' .00021) .00025 .00018 .00018
Kellys Rock HC12 HC 1 3 HC 1 4 HC1 5 HC1 6 HC 17 HC 1 8 HC 19 Johnson Rocks HC37 HC3& HC39 HC40 HC4l HC42 HC43 HC44 HC45 HC46 G74024A Milky Soak 1) HC53 HC54 HC55 HC56 HC57 HC58
Main Phase
Coarse-grained gran'odior i te/adamel lite
Rb-Sr GEOCHRONOLOGY, YILGARN BLOCK
111
I—continued
TABLE
2x
Samp 1 e ?
No.
Rock Coarse-grained
HC59
Standard
Type
granodiori
HC60 HC61
8 7
103
2.215
6.563
190
109 - 101
1.737
5 . H 9 5.061
0.95311 0.89846
.00013 .00020
0.90016
.00017
173 260
HC62 2)
Low
. More
HC48
Rb-Sr
mafi c
granodiori
te
Fine-grained
leucogranite
Fine-grained
leuco-adamel1ite
HC51A
More
HC49
mafic
granod iori
(I'ineated) (lineated)
te
Coarse-grained
Rainy
100
I.717 2.532
117
336
0.344
137 136
260 246
169 142
HG51B HG52
8 7
Sr/86Sr
error
7.527
0.98817
.00013
0.74098
.00012
0.76160 0.76412
.0.0023 .00021
Group
HC50
N.
Rb/86Sr
228
Rb/Sr
te/adamel1ite
granodiorite/adamel1i
te
0.528 •
0.999 1.536
0.551 0.686
1.603 1.998
0.499
1.451
0.77654 0.75582
.00020
285 230
0.795
2.319
0.78991
.00017
252
245
183
.00014
Rocks Granod i o r i t e / a d a m e l 1 i t e
HC63:
More
HC64 •
biotrte-rich
-
182 206 270
284
0.723 0.722
2.104
0.78210
.00003 .00021
122
2.206
6.528
0.94031
.00018
220
206
0.893 1.016
2.608
0.80287 0.81222
.00023
Granodiori te/adamel1i More
HC6.7
Granodiori
HC6|
More
biotite-rich
te
te/adamellite
biotite-rich
2.107
te
granodiorite/adamel1ite
granodiori te/adamel1i
HC6,5* HC66
200
granodiorite/adame11ite
170 168
HC69
196 407
-
2.970 1.211
0.417 0.430
390
0.78261
'
.00013 .00012
0.74742
1.246
.00008
200 264
258
0.778 2.120
2.268
0.74759 0.78520
125
6.270
0.93426
.00013 .00024
205 170
300
0.690
2.010
0.77864
.0001
330
0.511
1.486
0.75870
.0002
210
212
.0.988
2.887
0.80866
.00014
HC75 HG76
2.997 2.510
HC70 Granodiori te/adame11i
HC71* Little
Rainy
W B
te,
Rocks Granod iori t e / a d a m e l 1 i t e
HC73.
;
Rainy
Rocks
HG74
More
biotite-rich
granodiorite/adamel1ite
210
205
1.025
-195 203
227
0.860
HC7) HC7$
215
HG79
227 215 181
HC80
nim niikh mm
HC.81
«
S.E.
Yilgarn
-
Original
sample
LB282
81855'
LB 2 2 1 . -
82026 82030 8203^
K-feldspar Biotite
and
1.955 3.221
0.428
1.221
0.888
2.570
0.224
0.647
.0001 .0001
0.7155
.0003 .0002 .0004 .0002 .0001
•
0.7138 0.7183 0.7158
-0.330 0.406
0.140
ioooi
0.7287 0.7315 0.7780 0.8265
0.377 0.287 0.464
0.099 0.157 0.114
447 518 620
0.7178
0.722
254
_
_
24.35 533 1161
89.6
0.272
88.5 34.1
6.016 34.0
76.9 2.204 minerals '
Whole-rock
.0001
0.7171 0.7510
.0005 .0003 . 0002 .0001
0.7999 0.7280
7.48
2.52
0.9756 0.7380
.0015 .0014
1.360 6.088
.0002 .026
0.655
0.7265
0.033
0.7035 0.7643 0.7440
.0003 .0002
0.789 18.5 150.0
0.226
341 194
0.0114
24.7
352 111
0.222
526
42.1
57.5
58.3 60.4 -
0.584 -
12.50
1.699 0.645 41.1
0.986
Fines
<
75-8 -
#350
Plagioclase
-
K^feldspar
'
1.25 0.392
.0001 .0010
2.104
. 004
2.880
0.80645
3.66
0.83139 0.7489
.00035 .0001
184.4
9.68
19.05
.
Biotite
426
9.00
47.3
Error
calculated
better
than
as
precision
on
individual
analysis
-
does
not
take
into
HC f r o m
second
1.139 5.618
1.892
Muscovi t e
.
account
-
between
run
.0001 .0001
0.9063 2.646
65.6 195
.005
5.035
variation.
Real
.007
precision
± .05%.
Samples
prefixed
Bettenay
(1977).
Sample
.0005 .0001
0.78667
Schist
77246
]
,
203
206 he
Biotite
*
.0001
0.76977 0.77374
O.385 0.682 ,
0.671 1.09 0.118
415
ioclase
K-feldspar
2
.00017 .00019
0.77165
1.886
92
Whole-rock Flag
: Sphene
;*'
0.82434 0.80804
Sample
77292-
)
1.911 2.188
59 87 109 180
Whole-rock Plagjoclase „ _
1
0.636 0.656 0.750
.00013 .00009 .00017
Sample
77265
Evanston
0.649
0.133 0.235 0.249
167 552 660
65 70
LB359 LB360 LB167B
82035 82167
477 332 435 198
65
LB371 LB364 LB367A
Mt £ l v i r e
-
63 78 108
LB374 LB372
82025
Rocks
-
-
133 184
LB384
82023 82024
Rainy
-
3.321 2.834 1.890
.0001 .00021 .
0.82161
no.
LBX25 SP6
81962 81975
222 280
0.81570 0.79677 0.82854
3.381 3.198
1.155 1.093 1.135 0.970
Gneisses
81884 81878
175 197 200
excluded
772
from
from
f i r s t
isochron
fit
col l e c t i o n
-
period,
prefixed
collection
period
and
prefixed
8
collected
by
not
-
1
H. J. CHAPMAN, ET AL.
178
tonalite, ca 0.5 km , almost certainly belongs to this group of xenoliths. The post-kinematic Kellys Rocks body represents the most fractionated type in the area, showing similarities to the fractionated leuco-adamellites of Archibald et al. (1978). These are undeformed leuco-adamellites of uncertain extent intruded by dilational dykes of muscovite pegmatites and aplites. To the west are less fractionated biotite adamellites that contain disrupted bodies of banded gneiss at Cockatoo Rocks. Granite-greenstone boundaries, where exposed, are now either modified intrusive or highly tectonized contacts, although clearly most granitoids were originally magmatic intrusions into greenstone sequences. Some geological evidence does, however,- favour the preferred regional interpretation that the granitoids represent the products of reactivation of a sialic (gneissic) basement to greenstone belts. This evidence includes the local and more regional occurrence of banded gneisses that clearly predate intrusive granites, and the predominance of granites with minimum melt compositions but unfractionated trace-element geochemistry (cf. Bettenay, 1977) which contain widespread, partly assimilated gneiss xenoliths. The shallow-water (ensialic-type) environment suggested by sedimentary rocks, the limited thickness of the greenstone pile relative to total crustal thickness (see arguments in Archibald et al., 1981), and the Andean affinity of some felsic-acid volcanics (Hallberg et al., 1976) also require a sialic environment. SAMPLING AND ANALYTICAL METHODS Samples were originally collected from most of the significant granitic outcrops in the area (twelve localities). To overcome problems of weathering, specimens were obtained by detorfating explosives in percussion drill holes about 1 m deep. During initial collection, sites were selected to enable sampling of both pegmatitic and host granitic rocks from the same blast site. However, as discussed below, substantial homogenization of Sr isotopes and migration of Rb may have occurred adjacent to pegmatites up to 250m.y. after intrusion of the pegmatites. During subsequent sampling care was therefore taken to avoid pegmatites although this was difficult in the gneissic rocks. Samples of 2-10kg were jaw-crushed, and a 100gm split finely ground in an agate Tema mill. Rb/Sr ratios were determined in duplicate by Xray fluorescence spectrometry modified after the method of Norrish & Chappell (1967). Values of U.S.G.S. standard rock analyses by de Laeter & 2
Abercrombie (1970) were taken as standards. Rb/Sr ratios reproduce to better than ± 1 % and Rb and Sr concentrations calculated using mass absorption coefficients estimated from the Mo Compton scatter peak have a precision of ±5%. Rb/Sr ratios on mineral separates were determined by isotope dilution techniques (see de Laeter & Abercrombie, 1970, for details). Strontium was separated using standard column techniques (e.g. Lewis et al., 1975). Isotopic ratios were measured in a solid-source mass spectrometer of 30.5 cm radius of curvature and 90° magnetic sector. Samples from the initial collection period (distinguished in Table I) were measured using a triple rhenium filament source and a faraday cup collector. Samples from the second collection period were measured using a single tantalum, filament and signals from the collector amplified by a vibrating reed electrometer were digitized and processed by on-line computer. Sr/ Sr ratios were normalized to 8.3752. Repeated measurements of NBS 987 gave a value of 0.7102 ± .0002 (2a on mean) during analysis of the first batch of samples and 0.71010 ± .00034 during the analysis of the second batch of samples. Quoted precision on individual analyses in Table I is 2 x standard error on the mean of six to twelve batches of ten sets of measurements. These errors do not take into account between-run variation and the real precision will not be better than the ± .00034 (.05%) estimated for the NBS 987 standard. A Rb decay constant of 1.42 x l O ^ " has been used to calculate ages. Regression analysis was by the method of York (1969) using the programme published by Faure (1977), It should be noted that many of the ages previously published in adjacent areas are calculated using the programme published by Mclntyre et al. (1966) and that the latter programme results in significantly larger error estimates for isochrons with few samples. Errors are quoted at the la level. The term isochron is used to describe these best fits irrespective of whether the data lie within analytical error of the line or not. RESULTS Rb, Sr and 87s /86s analyses for all rocks incorporated on isochrons are given in Table I, and calculated ages and initial ratios in Table II. Resetting of ages: small-scale studies Results from the initial sampling gave ages for individual localities ranging from ca 2100m.y. to 2700 m.y. These isochrons were constructed from samples of pegmatite, aplite and host granite collected from one or two blast sites at each locality and the isochrons were based on three to five 88
86
- 1
r
r
1
Rb-Sr GEOCHRONOLOGY, YILGARN BLOCK
179
TABLE II
Results ofisochron regression using Model 2 (reciprocals of analytical errors as weights) based on York (1969) using program by Faure (1977). Errors of 10.5% and 2.0% assumedfor 87Sr/86Sr and Rb/Sr ratios respectively (errors at la level). \Rb = 1.42 x 10~11 yr~
No. of , samples
Rock Unit S. E. Yi lgarn Gnei sses Cockatoo who1e-rock
Age
Goodness of fit (Chi squared)
IR
13
2692 ± 61
0. 7021 ± . 0005
38
12
2490 ± 161
0. 7035 ± .0016
24
Compositional bands, from HC22, Cockatoo Rocks
14
2561 ± 6 2
0. 7027 ± . 0007
27 17
(exc1udi ng samples HC28 and HC31)
Rainy Racks (excluding HC65 and HC71)
20
2612 ± 37
0. 7014 ± .0011
Pigeon Rocks
10
2550 ± 53
0. 7151 ± .0072
3.7
Pegmatite and adjacent whole-rock samples, Pigeon Rocks
6
2302 ± 58
0.966
± . 064
0.24
Johnson Rocks (excluding HC42)
10
2726 ± 1 8 5
0.695
± .013
3.2
Milky Soak (Main Phase)
10
2557 ± 49
0. 7099 ± .0040
1.6
M-ilky Soak (Low Rb/Sr group) -
6
2550 ± 8 3
0.7040 ± .0018 '
4.2
Kellys Rock
8
2467 ± 36
0.719
Minerals "from pegmat i te, Ra i ny Rocks
4-
2454 ± 1 4
0. 7100 ± . 0007
0,43
Minerals from Mt Elvire l)r Plagioclase, K-feldspar, whole-rock - 2) Plagioclase, K-feldspar, whole-rock biotite (excluding sphene-rich heavy mineral concent rate)
3
2537 ± 17
0. 7023 ± .00008
0.16
4
2466 ± 62
0. 7025 ± . 0004
8.0
4 2 2
2354 ±'60 2036±40 1536 ± 30
0. 7099 ± .0018
4.0
Minerals from Evanston Schist 1) Whole-rock, fines, K-feldspar, plagioclase 2) Whole-rock - muscovite 3) Whole-rock - biotite
points. The scatter in ages was geologically unreasonable and could have been caused by variable initial 8 7 Sr/ 8 6 Sr ratios, later intrusion of pegmatites and aplites, by rehgnjpgenization of isotopes during subsequent events or by a combination of such factors. Resetting of Rb-Sr isotope systematica within a few hundred million years of emplacement characterizes some plutonic provinces {e.g. Page, 1978) and is reflected in higher initial 8 7 Sr/ 8 6 Sr ratios. However, it is precisely these ratios that are used to constrain the sources of granitic rocks in attempts to unravel the evolution of the sialic crust (e.g. Moorbath, 1976; Cooper et at., 1978). To investigate the possibility of resetting, alternative methods pf geochronology are needed and the results presented here are part of a preliminary investigation before the use pf whole-rock Pb-Pb and zircon methods. Useful information on the scale of homogenization can be provided by sampling pn different scales. To investigate this, two localities were sampled t^oth pn fjie whole-rock scale (sample sites spaced 5 rp fp 500 m) and on a smaller scale (site samples selected every 10-20 cm). At Pigeon Rocks, pegmatite and host granitic rocks were sampled over a distance of 1.5 m between two 15 cm thick pegmatites (Fig. 3) in addition to the collection of whole-rock samples from
3.80
*
3.60
± .012
0.42
Pegmatites and Adjacent Host Rock from Pigeon Rocks Pegmatite Host Adamellite • Age = 2302+58 m.y. I.R. = 0.966 + 0.064
CO CO
3.40 25 CM 3.20
Fig. 3.
R]b/Sr isochron plot for pegmatite and adjacent whple-ropk samples, Pigeon Rocks. Inset shows field sketch of sample sites. The slab sampled was detached by blasting.
pegmatite-free sjfes spaced a£ tp }Q§m ^paff. J h e isochrpn (pig. 3) cpnstfijctgc} frpm ppe pegmatite sample and five samples adjacent to the pggmatitg gives an age of 2302 ± 58m.y., initial 87Sr/ 86Sr ratio (henceforth IR) = 0.966 ± 0.064, which is younger than the whole-rock isochron age from the pegmatite-free sites (Fig. 4) of 2550 ± 53 m.y., IR = 0.7151 ± .0072. Several possibilities exist for the younger pegmatite age: (1) The pegmatite could have been intruded at
H. J. CHAPMAN, ET AL.
180
Pigeon
Rocks
Whole-Rock
Data
A g e = 2 5 5 0 i 5 3 m.y. I.R. = 0.7151 t O . 0 0 7 2
87
Rb/
/86
Sr
Fig. 4. Rb/Sr whole-rock isochron from Pigeon Rocks.
pegmatite. The whole-rock samples adjacent to the pegmatite have substantially lower mean Sr (19 ppm compared with lOOppm) and higher mean Rb (395 ppm compared with 300 ppm) than the whole-rock samples from pegmatite-free sites. The Cockatoo Rocks outcrop is the only example of the potentially older banded gneisses sampled within the Diemals area. The outcrop area is restricted (maximum 300 m x 200 m), rather weathered and cut by numerous pegmatites. The whole-rock isochron (Fig. 6) exhibits considerable scatter and excluding sample HC31, and sample HC28 which was collected adjacent to a pegmatite, gives an age of 2490 ± 161 m.y., IR = 0.7035 ± 0.0016. This increases to 2610 ± 99 m.y., IR = 0.701 ± .001 if HC28 is included. Compositional bands approximately 10 cm wide
— 2300m.y. and derived from a source with a high IR such as pre-existing high Rb/Sr material or much older crust. The mean Sr/ Sr ratio of the Pigeon Rocks whole-rock samples would have only evolved to 0.752 by 2300m.y. and could not have been the source for the pegmatitic material at this time. (2) The pegmatite could have been intruded substantially earlier than 2300m.y. and Sr isotopes rehomogenized in its vicinity at 2300 m.y. Figure 5 shows that samples with the mean Rb/Sr ratio of the pegmatite and adjacent host rocks would have evolved to an 87Sr/ 86Sr ratio of — 0.97 between 2500m.y. and 2300m.y.; the pegmatite may have been intruded close to the time given by the whole-rock age of Pigeon Rocks. (3) The 2300m.y. isochron could repre- Fig. 6. Rb/Sr data from Cockatoo Rocks, Jest-fit sent a mixing line between whole-rock and regression line calculated excluding HC31 and Cockatoo
Rocks
Whole - Rock
87
Samples
86
A g e = 2 4 9 0 1 1 6 1 m.y.
I.R. = 0 . 7 0 3 5 i 0.0016
87_. / 8 6 0 Rb / Sr
HC28.
were sawn from half the sample HC22 and these, together with the whole-rock value (the other half sample) give an age of 2561 ± 62m.y., IR = 0.7027 ± .0007 (Fig. 7). Without an independent method of assessing the age of the gneisses no firm conclusions can be reached about the significance of the whole-rock age. However, the following points should be considered. The age from the compositional bands of HC22 is similar to the whole-rock age but better defined. If this represents the formation age of the gneisses, and if Rb-Sr isotopes have not migrated substantially on a scale of Fig. 5. Sr/ Sr versus age of initial ratios and time 20 cm or more since (as required by the composidevelopment lines for Pigeon Rocks average whole-rock and pegmatite and adjacent tional band isochron), then the scatter in the whole-rock groups of samples. Time develop- whole-rock isochron is surprising. It should also ment lines for the highest and lowest Rb/Sr be noted that the range of Rb/Sr ratios of indisample of the pegmatite group are shown to vidual compositional bands (0.123-0.422) exceeds illustrate the effect of uncertainty in mean the spread in Rb/Sr ratios of most of the wholeRb/Sr ratios. rock samples (0.168-0.290 excluding HC28). Pegmatites and adjac Whole-rock sample
Pigeon Rocks Whole-rock 2600
2500
2400
Age m.y.
87
86
2300
Rb-Sr GEOCHRONOLOGY, YILGARN BLOCK
181
tered outcrops over the south-east Yilgarn Block by Bettenay (1977) was analysed and these define a more precise isochron of 2692 ± 61 m.y., IR = 0.7021 ± .0005 (Fig. 9) but with a scatter considerably in excess of experimental error. These gneisses represent a component of the pregreenstone basement postulated by Archibald & Bettenay (1977). Again, unless the gneisses have been substantially enriched in Rb, their mean Rb/Sr of about 1.0 constrains their derivation from a low Rb/Sr mantle or basic source to later than 2850m.y. (Fig. 8). .73
Tig. 7. Rb/Sr isochron, from compositional bands of sample HC22, Cockatoo Rocks.
Roddick & Compston (1977) show that resetting of whole-rock isochrons need only require smallsgale migration of Sr or Rb and the Cockatoo gneiss results could reflect such a process. If the outcrop is relatively homogeneous, the range of whole-rock Rb-Sr values might reflect incorporation of differing proportions of each band at each sample locality. The outcrop need only homogenize on the scale of —0.5 to 1 m to incorporate all band compositions, but whole-rock samples of 5 to 10 kg would not represent realistic average compositions at each site (if the scale of homogenization is —0.5m, a 10kg sample represents less than 1% of a sphere 0.5 m radius around the sample) and would resurrect an isochron reset on the scale of the bands. Thus there is evidence of isotopic resetting in synkinematic adamellite adjacent to pegmatites at Pigeon Rocks and strong indications of resetting in the whole-rock scale in large xenolithic blocks of granitic gneiss within biotite adamellite at Cockatoo Rocks. In both cases, diffusion distances need not have exceeded a few tens of centimetres. Whole-rock ages (a) Banded Gneisses. The poor fit to an isochron derived from the Cockatoo Rocks banded gneisses is discussed above. Even if the 2490 ± 161 m.y. isochron is reset, the initial ratio of 0.7035 and mean Rb/Sr ratio of the gneisses of 0.24 limits their existence since derivation from a mantle-type source of initial ratio —0.7005 to less than ca 250m.y. The gneisses are unlikely to have been formed much before 2850m.y., unless significant Rb was added during later events, and their relatively low Rb/Sr ratio makes this unlikely. A second suite of gneisses collected from scat-
.72
t/i oo 00
.70 Age
m.y.
Fig. 8. *'Sr/ Sr versus age of initial ratios and time development for whole-rock suites from this study. Errors in ages and initial ratios are correlated and are shown approximately as error boxes with la confidence interval. S6
(b) Intrusive Granitic Rocks. The intrusive granitic rocks are subdivided into: (1) foliated synkinematic plutons that predate the main deformations F , F and possibly Fj, and (2) the slightly foliated or unfoliated post-kinematic plutons. (1) Rainy Rocks is a strongly lineated, recrystallized, synkinematic granodiorite-adamellite. The lineation most probably results from the superposition of two planar fabrics (Fj and F or F 2
3
2
2
in. 80
.7 5
.70 Fig. 9. Rb/Sr isochron from banded gneisses collected from scattered localities in the southeast Yilgarn province (Bettenay, 1977).
H. J. CHAPMAN, ET AL.
182
<0° J>
2
4
6
87Rb/86Sr
Fig. 10. Rb/Sr whole-rock isochron from Rainy Rocks. The two high Rb/Sr samples are excluded from the calculated isochron fit. and F 3 ). A more leucocratic, rarely porphyritic, phase and an earlier more biotite-rich phase are present. At least two phases of pegmatites can be distinguished, the earlier being flattened to parallelism with the lineation and extensively boudinaged. The extreme stretching of this phase made it difficult to avoid during sampling and some of this material was incorporated in analysed samples. A second set of folded but crosscutting pegmatites was avoided. Samples were collected from three m a j o r outcrop areas. All samples except two with high R b / S r ratios define a relatively good isochron (Fig. 10) of 2612 ± 37 m.y., IR = 0.7014 ± .0011. It is not possible to distinguish either the two rock types or the different outcrops. Inclusion of the two high R b / S r samples HC65 and HC71 lowers the age to 2588 ± 3 3 m . y . , IR = 0.7021 ± .0010. Given the susceptibility of high R b / S r samples to resetting, the preferred age is 2612 ± 37 m.y. The Pigeon Rocks outcrop contains variably foliated synkinematic adamellite with numerous cross-cutting pegmatite veins, which were avoided during sampling. The isochron (Fig. 4) shows some scatter and "gives an age of 2550 ± 53 rq.y., IR = Q.7151 ± .Q072. ' Th[$ somewhat yppnger tharj, though just \ykhjn error of, the ^ainy Rocks isochron. The rather high initial ratio suggests that this isochron may be reset. For example, if the Pigeon Rocks body were intruded at 2650m.y. with an initial ratio of 0.701, its mean 87s r /86s r ratio would have evolved to 0.7155 by 2550m.y. Without independent control on the age, it is impossible to constrain the age of Pigeon Rocks to better than the period 2650m.y. to - 2 5 0 0 m . y . (2) Johnson Rocks comprises a foliated synkinematic hornblende tonalite preserved as a large (0.5 km 2 ) xenolith surrounded by a porphyritic
.
/
HC 42 4.5
87
/ 86 Rb/ Sr
5.0
Fig. 11. Rb/Sr whole-rock isochron from Johnson Rocks. HG42 is excluded from the calculated isochron fit. post-kinematic adamellite similar to that at Milky Soak. The relatively small spread in Rb/Sr ratio (Fig. 11) results in a relatively large error on the age and initial ratio of 2726 ± 185m.y., IR 0.695 ± .013 excluding sample HC42. The Milky Soak outcrop is mostly a weakly foliated post-kinematic porphyritic granodiorite adamellite with igneous textures well preserved. Two blast sites on the northern margin of the outcrop exposed fine-grained lineated leuco-adamellites and a more mafic granodiorite as well as a granodiorite apparently similar to the main phase. The two minor phases appear to pre-date the main phase and have been included in a separate isochron fit. The other samples collected over 1 km across ihp outcrop of the main granodiorite/adameilite give a relatively good fit to an isochron of 2557 ± 4 9 m . y . , IR = 0.7099 ± .0040 (Fig. 12). The low R b / S r group gives an isochron of 2550 ± 83 m.y., IR == ^.7Q4() + .0Q18. J h e ages of the two data sets c^e very sirnijar i p d although the initial ratio§ are nof distinct ait fhe 2a level it does not appear yalid to include all the data on one isochron. Again, the high initial ratios raise the possibility that these isochrons are reset. Both groups of samples could have formed at — 2665 m.y. with an initial « 7 Sr/ 86Sr ratio of 0.701, and provided each set behaved as a closed systerr), each would have evolved to tlieir observed initial ratios at 2550m.y. (Fig. 8). Again, without independent evidence, no firm conclusions can be reached other than that the Milky Soak granitic rocks must have been formed in the period 26502500m.y. Kellys Rock is an unfoliated post-kinematic leuco-adamellite with well-preserved igneous tex-
Rb-Sr GEOCHRONOLOGY, YILGARN BLOCK
183
results are not reported here. Three samples gave consistent results as illustrated in Figures 14, 15, and 16. Sample 77292, a granodiorite from Mt Elvire gives a plagioclase: K-feldspar: wholerock age of 2537 ± 17m.y. (Fig. 14) and a biotite: whole-rock age of 2359 ± 30m.y. Sample 77265, a concordant pegmatite from Rainy Rocks gives a whole-rock : plagioclase: Kfeldspar : biotite isochron of 2454 ± 14m.y. (Fig. 15). Sample 77246 is a schist sampled from drill
87
/ 86
Rb/
Sr
Fig. 12. Rb/Sr isochrons from Milky Soak.
tures, and is characteristic of a small but distinct group of late plutons. Eight samples define a good isochron (Fig. 13) of 2467 ± 36m.y., IR = 0.719 ± .012. The high initial ratio could indicate that this isochron is reset although the small scatter of the data is consistent with this being a primary age. Even if resetting has occurred, the high mean R b / S r ratio of the body precludes its intrusion with a low initial 8 7 Sr/ 86$r ratio prior to ~ 2550 m.y. (Fig. 8) in agreement with the geological evidence that the Kellys Rock pluton is younger than the other plutons dated.
Kellys
18
20
Rock
22
87
24 /86
Rb/
26
Fig. 14. Rb/Sr mineral data from sample 77292, Mt Elvire.
core in the greenstone sequence near Evanston (Fig. 2). This gives a whole-rock: fines < # 3 5 0 : plagioclase: K-feldspar isochron of 2354 ± 60 m.y. (Fig. 16), a whole-rock: muscovite age of 2036 ± 40 m.y. and a whole-rock : biotite age of 1536 ± 30m.y. The older mineral isochron from Mt Elvire suggests that the main metamorphic episode had terminated by 2537 ± 17 m.y. in this area, and is consistent with other mineral ages f r o m the eastern half of the Yilgarn Block (Libby & de Laeter, 1979) which mostly range from 2600 to
28
Sr
Fig. 13. Rb/Sr whole-rock isochron from Kellys Rock.
Mineral Ages Mineral separates were analysed f r o m a number of the rocks f r o m the first collection. Many of the samples showed considerable scatter on 87Sr/86Sr versus, ^Rb/ ^St plots and these
Fig. 15. Rb/Sr mineral data from sample 77265, Rainy Rocks.
H. J. CHAPMAN, ET AL.
184
m.y. The ages of the less precisely dated Johnson Rocks and Pigeon Rocks, the other pre- or synkinematic plutons dated, are within error of the Rainy Rocks age. Bettenay (1977), Archibald et al. (1978) and Archibald et al. (1981) conclude that the synkinematic and post-kinematic granodioriteadamellite plutons are derived by partial melting of banded gneisses. Within the errors on the ages, and particularly with the possibility of resetting causing the higher initial 87Sr/86Sr values of Pigeon Rocks and the two groups at Milky Soak, it is not possible to preclude this hypothesis. It should be noted that Rainy Rocks has a lower mean Sr/ Sr ratio than both the S.E. Yilgarn banded gneisses and the Cockatoo Rocks banded gneisses, although the errors overlap the rather large errors of the Cockatoo Rocks banded gneisses. Both Pigeon Rocks and the main high Rb/Sr group of samples at Milky Soak have initial ratios substantially higher than the banded gneisses at 2550m.y. These could have been derived by melting substantially older gneissic crust at that time, or they could have been derived from the observed banded gneisses at 2600-2650 m.y. and reset at 2550m.y. They could not have been derived from the S.E. Yilgarn or Cockatoo gneisses at 2550 m.y. (Fig. 8). The low Rb/Sr group of samples from Milky Soak have a similar mean 87s /86s to the S.E. Yilgarn banded gneisses at 2550 m.y. The older of the mineral ages of 2537 ± 17 m.y. suggests that the major thermal peak of the metamorphism had terminated by that time. Petrographic evidence suggests that medium- or high-grade metamorphic conditions accompanied and outlasted the main deformation events. The Milky Soak age might either reflect derivation by partial melting late in this event, or resetting of Rb-Sr isotopic systems within a previously intruded pluton. The youngest "event" in the area was the intrusion of the undeformed Kellys Rock leucoadamellite at 2467 ± 36m.y. with a high initial Sr/ 86Sr ratio of 0.719 ± 0.012. The initial ratio could indicate derivation from the previously emplaced plutons with higher Rb/Sr ratios. Most of the ages in the Diemals area correspond very closely to the ages recorded by Roddick et al. (1976) and Cooper et al. (1978) in the Lawlers-Agnew-Perseverance area (Table III). Both areas are characterized by major pre- or syn-kinematic plutonic activity within the interval 2670-2600m.y. In the Lawlers area, Cooper eial. (1978) bracket the D, deformation by the ages of the Lawlers tonalite 2652 ±20m.y. and the Lawlers leuco-granite 2576 ± 14m.y., and the D deformation by the Lawlers leuco-granite and a 87
Fig. 16. Rb/Sr mineral data from sample 77246, Evanston.
2500m.y., although ages as young as 2200m.y. are recorded. It is also consistent with derivation of pegmatites at Pigeon Rocks at -2550m.y., and resetting of the Pigeon Rocks and Cockatoo Rocks whole-rock isochrons to -2550 m.y. The younger mineral ages may reflect progressively later uplift of more deeply buried crust and lower blocking temperatures, particularly of biotite. Even so, the 1536m.y. biotite: whole-rock age from 77246 is substantially younger than any other biotite age reported within the central or eastern Yilgarn Block. DISCUSSION The ages of both banded gneisses and greenstones are still too poorly constrained to elucidate whether the former are candidates for basement to the latter. The minimum age of 2718 ± 50 m.y., IR = 0.7007 ± .0004 for the Kathleen Valley Granophyre and Gabbro 250 km northeast of Diemals (Cooper et al., 1978), and an age of 2635 ± 80m.y., IR = 0.7029 ± 0.0015 from the Marda felsic volcanic complex in the south of the Diemals area (Hallberg et al., 1976), provide the best constraints on the age of the greenstones: As discussed above, precursors to the banded gneisses cquld have been intruded at any time between about 2850 m.y. and 2600m.y. The ages of Rainy Rocks (pre- or synkinematic) and Milky Soak (post-kinematic) constrain the main deformation events to within the period 2612 ± 37m.y. to 2557 ± 49m.y. The younger age of Milky Soak is consistent with the observation that it contains xenoliths of the distinctive lineated granodiorite-adamellite characteristic of Rainy Rocks. The low initial 87$ / 86s ratio of Rainy Rocks precludes its derivation from a basic or mantle Rb/Sr type source more than a few tens of millions of years prior to 2612 1
r
r
86
r
r
87
2
Rb-Sr GEGCHRONOLOGY, YILGARN BLOCK TABLE
185
III
Tectonic events in the Diemals area compared with the Lawlers-Agnew-Perseverance area (all errors at 95 % confidence level) Diemals Area Event Mineral ages, Evans ton whole-rock - b iot i te whole-rock - muscovi te Pegmatites at Pigeon Rocks Mineral ages, Evanston Mineral ages, Rainy Rocks Mineral ages, Mt E1vire inc 1. biotite Ketlys Rock teuco-adarnel l ite Lafe cataclasis over S.E. of pie/rial's area Mt Elvire minerals uplift Milky Soak Compositional bands at Cockatoo Rocks Ma i n De fo fma t i on s , D1 (?) , Da and D3 P i geon. Rocks Rainy Rocks Johnson Rocks. Marda Felsic Volcanics S.E. Yilgarn Gnei.sses 1 Age may be reset Hal Iberg et al. , 1976
Age (m.y.) 1536 ± 60 2302 4 120 2354 ± 120 2454 ±28 2466±124 2467 ±72
12036 ± 80
2537±34 2550 ±98 2561 ± 124
2
x
IR
O.719± .024
.7099 ±.0080 .7027 ±.0014
0.715± .014 0.701 ±.002 0.695 ±.02 0.703± .002
1
0.7021 ±.0010
2692 ± 122
JJR
0.966 ±.128
2550 ± 106 2612 ± 74 • 2635±80
22726±370
Lawlers-Agnew-Perseverance Area Roddick et al. (1976), Cooper et al. (1978) Event Age (m.y.)
late leuco-tonalite 2474 ± 14m.y. In the Diemals area the main deformations are thought to have occurred before the intrusion of the Milky Soak granodiorite/adamellite (2557 ± 49 m.y.). In both areas cooling from the metamorphic event started during the interval 2600-2500 although possibly somewhat earlier in the north where Roddick et al. (1976) record biotite ages of 2565 ± 15m.y. The last plutonic event in both areas was at - 2 4 7 0 m . y . with the Lawlers leuco-tonalite at 2474 ± 14m.y. and Kellys Rock leucoadamellite at 2467 ± 36 m.y. CONCLUSIONS Rb-Sr geochronology in the Diemals-Evanston area reveals the following sequence of events (summarized in Table III). The age of the earliest events including the intrusion of precursors to banded gneisses and deposition of the greenstone sequence is not well established. Banded gneisses may have been formed at any time in the interval 2850m.y. to 2600m.y. Evidence from outside the present study area suggests that greenstones were deposited at some time in the interval 2800m.y. to 2650 m.y. The greenstones were intruded by widespread granodiorite and adamellite plutons between 2650m.y. and 2500 m.y. Earlier members of this suite (Rainy Rocks 2612 ± 37 m.y.) were deformed and then intruded by later "postkinematic" members of the suite (Milky Soak
Leucotonalite and aplite Pegmat i te Second Deformation - D2 Uplift (Mt Keith biotite ages) Lawlers 1eucotona1ite Pegmat i te First Deformation Di Jones Creek-Lawlers conglomerates Perseverance gne i ss Mt Kei th Granodiori te Lawlers Tona1i te Greenstones (Kathleen Valley Granophyre 6 Gabbro)
2474 ±14' 2481 ± 18
.7019 ±.0001 .7253 ±•.0002
2565 ± i 5 2576 ± 14 .7022 ±.. 0002 2588 ± 18 . 762 ±. 007 . 2625 ± 34 2652±20 2718± 50
. . 7022 ±.005 • 7015 ±.0002 . ,7015± .0001 • .7007 ±.0004 .
2557 ± 49m.y.). The crust underwent a major metamorphic episode over the period —2650 m.y. to 2550m.y., and the oldest mineral ages of 2537 ± 17 m.y. record uplift and cooling from this event. The metamorphic event may have reset Rb/Sr isotopic systematics in the Pigeon Rocks pluton and the Cockatoo Rocks at —2550 m.y. and caused pegmatite formation at Pigeon Rocks at this time. The last plutonic event was the intrusion of a post-kinematic leuco-adamellite at 2467 ± 36m.y. with a high initial 8 7 Sr/ 8 6 Sr ratio of 0.719 ± .012. Mineral ages around 2350 ,±60 m.y. and resetting of Rb/Sr isotopes adjacent to pegmatites at Pigeon Rocks at 2302 ± 58m.y. record events of uncertain significance. Younger muscovite and biotite ages of —2000 and — 1540 m.y. presumably reflect cooling of these minerals below relatively low blocking temperatures during progressive uplift of the terrain. The ages within this sequence correspond surprisingly closely to those of similar events in other areas within the Yilgarn Block. The granodiorite-adamellitic and more rarely tonalitic plutonic event most likely at 2660-2600m.y. appears to be particularly widespread as does the associated and subsequent metamorphic episode. There are strong suggestions of rehomogenization of Rb/Sr isotopes by metamorphic events, subsequent to intrusion of individual plutonic bodies. Evidence for this invariably includes the high initial 87 Sr/ 86Sr ratios of reset isochrons
186
H . J . C H A P M A N , ET
AL.
with isochrons preserved on a small scale also been derived by partial melting of a high 87$ / S r source, or may have been reset, and in the being resurrected on the whole-rock scale. There is also metasomatic alteration of host rock adja- absence of ages f r o m alternative isotopic systems cent to pegmatites. This evidence of resetting that are less easily reset, it is not possible to promakes interpretation of granitic bodies with high vide unambiguous interpretations of such data. initial ratios ambiguous. Such bodies may have REFERENCES KING, A. C . , 1 9 7 4 : Regional and structural geology of ANDERSEN, L . S . , BETTENAY, L . F . , BINNS, R . A . , DE the Diemals Find area, North Yilgarn Goldfield, LAETER, J . R . , GORTON, M . P . , & GROVES, D . I . , Western Australia. B.Sc. (Hons) Thesis, Univ. 1976: Archaean crustal history of the Central YilWest. Aust. [unpublished]. garn Block, Western Australia. Abstracts, /, 3-4. 25 th internal. geol. Congr., Sydney. LEWIS, J . D . , ROSMAN, K . J . R . , & DE LAETER, J. R., 1975: The age and metamorphic effects of. the ARCHIBALD, N. J., & BETTENAY, L. F., 1977: Indirect Black Range dolerite dyke. Ann. Rep. geol. Surv. evidence for tectonic reactivation of a pre-greenWest. Aust. for 1974, 86-88. stone sialic basement in Western Australia. Earth planet. Sci. Lett., 33, 370-378. LIBBY, W. G., & DE LAETER, J. R., 1979: Biotite ages and cooling history at the western margin of the ARCHIBALD, N . J . , BETTENAY, L . F . , BICKLE, M . J . , & Yilgarn Block. Ann. Rep. geol. Surv. West. Aust. GROVES, D. I., 1981: Evolution of Archaean crust for 1978, 79-87. in the Eastern Goldfields Province of the Yilgarn Block, Western Australia. Spec. Pubis geol. Soc. M C I N T Y R E , G . A . , BROOKS, C . , COMPSTON, W . , & Aust., 7, 491-504. T U R E K , A., 1966: The statistical assessment of Rb/Sr isochrons. J. GeophysRes., 71, 5459-5468. ARCHIBALD, N . J . , BETTENAY, L . F . , BINNS, R . A . , GROVES, D . I . , & GUNTHORPE, R . J . , 1 9 7 8 : T h e MOORBATH, S. M . , 1976: Age and isotope constraints evolution of Archaean greenstone terrains, Eastern for the evolution of Archaean crust; in Windley, Goldfields Province, Western Australia. Precamb. B. F. (Ed.) The Early History of the Earth, 351Res., 6, 103-131. 360. Wiley, London. BETTENAY, L. F., 1977: Regional geology and pedoNORRISH, K., & CHAPPELL, B. W., 1967: X-ray fluoresgenesis of Archaean granitoids in the south-eastern cence spectrography; in Zussman, J. (Ed.) Physical Yilgarn Block, Western Australia. Ph.D. Thesis, Methods in Determinative Mineralogy, 161-214. Univ. West. Aust. [unpublished]. Academic Press, London. BINNS, R . A . , GUNTHORPE, R . J . , & GROVES, D . I . , PAGE, R. W . , 1 9 7 8 : Response of U.Pb zircon and Rb/Sr 1976: Metamorphic patterns and development of total-rock and mineral systems, to low-grade greenstone belts in the Eastern Yilgarn Block, regional metamorphism in Proterozoic igneous Western Australia; in Windley, B. F. (Ed.) The rocks, Mount Isa, Australia. J. geol. Soc. Aust., Early History of the Earth, 3 0 3 - 3 1 3 . Wiley, Lon25, 1 4 1 - 1 6 4 . don. PLATT, J . P . , ALLCHURCH, P . D . , & RUTLAND, COOPER, J . A . , NESBITT, R . W . , PLATT, J . P . , & MORTIR. W. R., 1978: Archaean tectonics in the Agnew MER, G. E., 1978: Crustal development in the supracrustal belt, Western Australia. Precamb. Agnew region, Western Australia, as shown by Res., 7, 3-30. Rb/Sr isotopic and geochemical studies. Precamb. PORTER, D . , 1 9 7 1 : Metamorphic and structural geology Res., 7, 31-59. of an area near Lake Giles, Yilgarn and North DE LAETER, J. R., & ABERCROMBIE, I. E., 1970: Mass Coolgardie goldfields, Western Australia. B.Sc. spectrometric isotope dilution analyses of rubidium (Hons) Thesis, Univ. West. Aust. [unpublished]. and strontium in standard rocks. Earth planet. Sci. RODDICK, J . C . , COMPSTON, W . , & DURNEY, D. W., Lett., 9, 327-330. 1976: The radiometric age of the Mount Keith DE LAETER, J . R . , LIBBY, W . G . , & TRENDALL, A . F . , Granodiorite, a maximum age estimate for an 1981: The older Precambrian geochronology of Archaean greenstone sequence in the Yilgarn Western Australia. Spec. Pubis geol. Soc. Aust., Block, Western Australia. Precamb. Res., 3, 55-78. 7., 145-157. RODDICK, J . C . , & COMPSTON, W . , 1977: Strontium FAURE, G., 1977: Principles of Isotope Geology. Wiley, isotopic equilibration: a solution to a paradox. New York. Earth planet. Sci. Lett., 34, 238-246. GEE, R. D., 1979a: Structure and tectonic style of the W A L K E R , L, 1974: The geology of the Mt Elvire area, Western Australian Shield. Tectonophys 58 North Yilgarn and North Coolgardie goldfields. 327-369. B.Sc. (Hons) thesis, Univ. West. Aust. [unpub, 1979&: Explanatory notes on the Southern lished]. Cross 1:250,000 geological map sheet. Rec. geol. WINKLER, H . G. F., 1 9 7 4 : Pedogenesis of Metamorphic Surv. West. Aust., 1979/5. Rocks. Springer-Verlag, Berlin. HALLBERG, J . A . , JOHNSTON, C . , & BYE, S . M . , 1 9 7 6 : 1969: Least squares fitting of a straight line The Archaean Marda igneous complex, Western YORK,withD.,correlated errors. Earth planet. Sci. Lett., 5, Australia. Precamb. Res., 3, 1 1 1 - 1 3 6 . 320-324. r
86
Sm-Nd DATING OF THE NORTH STAR BASALT, WARRAWOONA GROUP; PILBARA BLOCK, WESTERN AUSTRALIA Hamilton, P. J., 1 Evensen, N. M., 2 O'Nions, R. K., 1 Glikson, A. Y. 3 & Hickman, A. H . 4 1 Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, England 2 Department of Geology, University of Toronto, Toronto, Ontario, Canada 3 Bureau of Mineral Resources, Geology and Geophysics, PO 378, Canberra City, ACT2601, Australia 4 Geological Survey of Western Australia, 66 Adelaide Terrace, Perth, Western Australia 6000
ABSTRACT Sm-Nd isotopic data are reported for metavolcanics f r o m the North Star Basalt Formation of the W a r r a w o o n a G r o u p in the Pilbara Block of Western Australia. The samples, which include komatiitic, basaltic, andesitic and dacitic rocks, are dated 3560 ± 32 m.y. which is interpreted as the age of volcanism. The initial 1 4 3 N d / 1 4 4 N d ratio is 0.508104 ± 34. These results are within error of those obtained f o r the Onverwacht G r o u p Volcanics f r o m the Kaapvaal craton of southern A f r i c a .
INTRODUCTION The Pilbara Block in Western Australia is similar to many other Archaean cratons in that it is a granite-greenstone terrain. The original stratigraphic relationships between the oldest granitoid components and the volcanic associations in the greenstone belts remain uncertain. Radiometric data have often failed to provide solutions to such problems because decay schemes, such as K-Ar, Rb-Sr and U-Th-Pb, otherwise suitable for dating Archaean rocks, frequently do not record primary crystallization ages. Rather, such data generally reflect varying degrees of isotopic re-equilibration and parentdaughter fractionation during subsequent mineralogical changes. This situation has been somewhat alleviated in recent years by the application of U-Pb dating of zircons and Sm-Nd whole-rock dating which yield primary crystallization ages in such terrains. For example, radiometric results obtained from Sm-Nd whole-rock analyses by Hamilton et al. (1978) and from UPb zircon studies by Michard-Vitrac et al. (1977) are in excellent agreement (3769+ m.y. and 3770 ± 42m.y. respectively) at dating the igneous activity of the Isua greenstone belt in West Greenland. Volcanism of the Onverwacht Group
Spec. Pubis geot. Soc. Aust., 7 (1981)
in the Kaapvaal craton of southern Africa was first successfully dated at 3540 ± 30m.y. (Hamilton et al., 19796) using the Sm-Nd technique. As in southern Africa, a similar divergence of opinion has existed about the temporal relations in the Pilbara Block of Western Australia between granitoid gneisses and greenstone belts (c/. Glikson, .1972; Hickman, 1975). The present study is specifically aimed at acquiring a precise Sm-Nd date for metavolcanics near the base of the greenstone belt succession in the eastern Pilbara Block. GENERAL GEOLOGY AND PREVIOUS GEOCHRONOLOGY The Pilbara Block is the smaller of the two Archaean cratons in the Western Australian shield (Fig. 1). The regional geology of the area is summarized by Hickman (1981). The greenstone succession, known as the Pilbara Supergroup, occurs as synclinal belts of metavolcanic and metasedimentary successions between granitoid batholiths which are generally dome-shaped. Preand syn-tectonic foliated migmatitic and granite gneiss batholiths are intruded by virtually unfoliated post-tectonic granites.
188
P J. HAMILTON, N. M. EVENSEN, R. K O'NIONS, A. Y. GLIKSON & A. H. HICKMAN
— 20 PILBARA BLOCK
.Marble Bar * area (Fig 1b)
25t MARBLE bAM •
= X » B i g Stubby
.
^ v ^ - . V ^ ^ . * -;/;.
• : • 9°5?' y V - • j-'. V • .V \ • n
!
Proterozoic rocks
-30°
Post-tectonic granitoid
Inclined strata dip
^
Top of bed
x
Sampling site
Foliated granitoid
(pillows)
Duffer Formation Mount Ada Basalt
^ ^
McPhee Formation
River Main road
North Star Basalt
— 35
115°
120
0 I
km
10
Fig. 1. (a) Locality map of study area in Western Australia, (b) Geological setting of the sampling sites.
This study has centred on the North Star Basalt—the oldest formation of the Taiga Taiga Subgroup. More recent pertinent geochronologic results of reasonable precision are summarized in Table I, and the constraints established on temporal relationships of the various lithologic units are briefly discussed below. , The Duffer Formation and Salgash Subgroup are conformable and must be > the 3452 ± 16m.y. age obtained for zircon from dacite in the Duffer Formation (Pidgeon, 1978a). The Rb-Sr and Pb-Pb studies of older granitoids (Oversby, 1976) indicate that the main period of deformation (D2) occurred at ca 2950m.y. and post-D2 granitoids intruded ca 350m.y. later (see Table
I). The crustal residence time of the precursory materials to some of the older granitoids is poorly constrained but, on the basis of inferred initial Sr and Pb isotope compositions (Oversby, 1976; de Laeter et al., 1975), is mainly ca 200 m.y. prior to D2. However, some of the granitoids are known, from Rb-Sr (de Laeter, pers. comm.; Cooper et a/., 1980), U-Pb (Pidgeon, 19786) and Sm-Nd (McCulloch, pers. comm.) to range in age from ca 3200 to 3490 m.y. SAMPLES AND RESULTS Six samples of the North Star Basalt metavolcanics from the Marble Bar area (see Fig. 1) were selected to give as wide a range in Sm/Nd ratio as possible. These include one komatiite,
Sm-Nd DATING OF NORTH STAR BASALT TABLE
189
I
Geochronologic data for the east Pilbara region Locality and Rock Type A.
Age (m.y.)*
Reference
Warrawoona -Group
Glen Herring area. Lat. 21°21'25" S. Long. 119°36'11" E. Dacite in Duffer Formation 2. Doolena Gap. Lat, 20°54' S. 1 Long. 119°43 E. Sal gash Subgroup
U-Pb z i rcon
3452 ± 1 6
Age of igneous crystal 1ization Extrus i ve/Intrusi ve?
Pb-Pb model ages for galena
3340
Only one sample measured but close to result for 1.
3. 'Big Stubby'. Lat. 21°43' S. Long. T l f ^ ' E. Duffer Formation
3500
Older Granitoids (affected by D 2 deformation; migmatites
Pidgeon (1978a)
Richards (1978)
Sangster & Brook (1977)
granitic gneisses)
k. Gneissic granites widely spaced ' within Pilbara Block
Rb-Sr whole rocks
2986 ± 176
Compston & Arriens (1968)
5. Mount Edgar Bathoii th near Moo1 ye 1 la, west' of Marble Bar
Rb-Sr whole- rock errorchron
3059 ±358
de Laeter & Block ley (1972)
6. a) Shaw Bathoiith
Rb-Sr whole rock
2889± 81
b) As above at Cogglegong, 50 km SW of Marble Bar 7. Tambourah, 20 km south of 6. b
Pb-Pb age for whole rock and minerals Pb-Pb age for two whole rocks and feldspars ditto for two different samples Rb-Sr internal isochrons
2936 ± 9
8. Woodstock, 50 km west of 6.b
Pb-Pb ages for whole rocks and minerals
Good agreement and interpreted as age of metamorph i sm
a) 3070 ± 12
Possibly a relict primary age
b) 2938 ± 33
Metamorphic age in good ageement with 6 Metamorphic perturbation of Rb-Sr system
c) 2698 ± 14 2610± 14 a) 2961 ± 45 b) 2786 ± 38 c) 2769 ± 13
Mt Newman Ra i1 way Line. Lat. 22°. , Long. 119°
Pb-Pb ages for whole rocks and minerals
10. Mt Newman. Ra i 1 w a y v a r i o u s local it res north of 9
Rb-Sr and Pb-Pb internal isochrons
11. War eery Gap. Tona1i te-g ra nodi ori te including Corunna Downs Batholith
Rb-Sr whole-rock
5280 ± 20
12. Shaw Batholith, migmat i te
U-Pb, zircon
3417 J
40
13. Mount Edgar Batholi th, foliated gran i te
U-Pb, zi rcon
3280 *
20
2961 ± 7
Oversby (1976)
a)=*siightly perturbed primary age Metamorphism at ca •2770 m.y. Metamorphic age. Good agreement with 6, 7-b, 8.a. Suggest 2950 m.y. time of D 2 deformat ion Possible metamorphic disturbances at ca 2750, 2200, 2100 and 2000 m.y. Intrusive into Lower Warrawoona Group
Cooper et al. (l Pidgeon (19782?)
Post-Tectonic Granites (unaffected by~D 2 deformation) Rb-Sr who 1 e-rock
2614± 93
Intrusive into 5.
de Laeter £ Block ley 0972)
Rb-Sr whole-rock
2547 ±129
Intrusive into 6.
de Laeter et al. (1975)
16. Cookes Creek Granite, 50 km SE of 13
Rb-Sr whole-rock
2514 ± 36
Intrusive into Warrawoona Group volcanics
de Laeter et al. (1977)
17. Gregory Range Granite
Rb-Sr whole-rock
2585 ± 59
18. Copper Hi M s
Rb-Sr whole-rock
2820 ± 64
1ntrus ive Forma t i on
de Laeter & Trendall (1970)
K4. Moolye1 la Adame 11 i te 15. Cooglegong Adame11i te
Porphyry
,
* Al I errors in this paper are quoted at the 2a level.
Decay constants used are as recommended by Steiger & Jager (1977).
one high-Mg basalt, one tholeiite, one andesite and two dacites. The rocks have been altered by greenschist-facies metamorphism during which certain major and trace elements were mobilized as documented by Glikson & Hickman (1981). Details of the chemical separation of Sm and Nd and subsequent mass-spectrometric, isotoperatio measurements have been described previously (O'Nions et al,, 1977; Hamilton et al., 1977). The Sm and Nd contents and i43Nd/ * 44 Nd ratios a r e presented in Table II. H3Nd/ * 44 Nd ratios range f r o m 0.514196 ± 30 to 0.510334 ± 36, and are plotted on a Sm-Nd evolution
diagram in Figure 2. The data form, a good linear array with each datum being within error of the regression line obtained by the method of York (1966). The slope of this line corresponds to an age of 3560 ± 32 m.y. (2crm) and yields an initial l43Nd/ i 4 4 Nd (I N d ) of 0.508104 ± 34 (2a J . DISCUSSION It is not known whether the dacite samples analysed here are directly related by simple igneous differentiation processes to the more mafic rocks. However, all six samples fall within error of a single line on the Sm-Nd evolution diagram (Fig.
190
P. J. HAMILTON, N. M. EVENSEN, R. K. O'NIONS, A. Y. GLIKSON & A. H. HICKMAN
0-514
AGE =3560 ± 32 m.y. iNd =0-508104*34
0-513
3Nd 4Nd
0-512
0-5110
0-01
0-15 7
Fig. 2.
0-25
0-20
Sm/ 1 4 4 Nd
Sm-Nd •evolution diagram for North Star Basalt samples. The decay constant for 147 Sm was taken as 6.54 ~> _ •
2) and must have crystallized within a short time interval and have experienced similar prehistories with respect to S m / N d evolution. The S m / N d age obtained is interpreted as the time of the volcanicity responsible for the North Star Basalt, and is the oldest age yet obtained from the Australian continent. The oldest published age for the foliated granitoids to the southeast of the study area is 3280 ± 20m.y. and for all granitoids of the Pilbara Block is 3417 ± 40m.y. (Pidgeon, 19786). Unpublished preliminary Sm-Nd data (M. McCulloch, pers. comm.) indicate ages of 3490 to 3380 m.y. for granitoids east and southwest of Marble Bar. Parts of the granitoid gneiss terrain investigated by the Rb-Sr and U-Pb methods are apparently-car 300 m.y. younger than the North Star Basalt, the oldest result being a Rb-Sr age of 3279 ± 162m.y. (J. R. de Laeter, pers. comm.) and 3280 ± 20 m.y. (Cooper et al„ 1980). Further isotopic studies using Sm-Nd whole-rock and U-Pb zircon dating might establish the existence of older granitoids, as postulated by Hickman & Lipple (1978). - The Sm-Nd age obtained for the North Star Basalt is ca 100m.y. greater than the U-Pb zircon age (Pidgeon, 197.8a) obtained from dacite in the
TABLE
II
Sm-Nd data for the North Star Basalt Sample and Rock Type
S.m (ppm)
75040006
1.206
75040016P
komatiite high-Mg basalt •
2.410
75040016H
tholeiite
75040024D
andesite
3-996
75040025A
dacite
.2.915
2.165
75040025B
dacite
7.802
Nd (ppm)
"7Sm/ ^Nd*
,v
Nd/ Nd
: 20
0.514196 ±30 8.101 0.1789 0.512306 + 26 7.408 0.1758 0.512256 + 28 0.510725 i 24 21.55 0.1114 0.510864 ±22 14.91 - 0.1175 50.09 0.09364 0.510335 136 2.808
0.2584
overlying D u f f e r Formation. It is uncertain from field relationships whether this dacite is extrusive or part of a high-level intrusion. Four other workers have dated the D u f f e r Formation at six localities using P b - P b , U-Pb and Sm-Nd methods, yielding an age range of 3550-3452m.y. A time-integrated S m / N d ratio may be calculated for the source of the North Star Basalt. This parameter describes single-stage evolution from 4550m.y. ago, the assumed time of solar system formation, to the time of crust-mantle differentiation. The i43Nd/i44Nd ratio at 4550m.y. ago is assumed to be the same as I N d ( = 0.50682) for the Angra dos Reis achondrite (Lugmair & Marti, 1977). The model S m / N d ratio is 0.308 ± 4(2<7m) for the North Star Basalt source region. With the
Sm-Nd D A T I N G O F N O R T H STAR BASALT Fig. 3.
\Sm/Nd=chondritic
191
Plot of age (t) and I N d values obtained f r o m whole-rock isochron studies on Archaean meta-igneous rocks (Hamilton et ah, 1978, 1979*7, b, and this work; Zindler et at., 1978). These data together with less precise results from other whole-rock suites (Hamilton et a/., 1977; Jacobsen & Wasserburg, 1978) and single whole-rock analyses (DePaolo & Wasserburg, 1976a, b) are consistent with a chondritic S m / N d ratio for the Earth.
& W a s s e r b u r g , 1979), A r c h a e a n m e t a - i g n e o u s rocks appear to have been derived f r o m sources with similar time-integrated S m / N d ratios a n d c l o s e t o t h e c h o n d r i t e v a l u e o f 0 . 3 1 ( E v e n s e n et at., 1978). T h i s p o i n t is i l l u s t r a t e d in F i g u r e 3 w h e r e t h e r e a s o n a b l y p r e c i s e a g e (t) a n d I N d values obtained f r o m whole-rock isochron studies a r e c o m p a r e d w i t h , a n d seen t o p l o t c l o s e t o , t h e c h o n d r i t i c S m / N d e v o l u t i o n line.
ONVERWACHTW>,LBARA
ACKNOWLEDGMENTS 3500 Time m.y.
exception o f t h e S t i l l w a t e r C o m p l e x , w h o s e parental m a g m a m a y h a v e interacted with older continental crust with low S m / N d ratio ( D e P a o l o
This work was supported by US National S c i e n c e F o u n d a t i o n G r a n t E A R 75 2 0 8 9 1 . It is p u b l i s h e d w i t h t h e p e r m i s s i o n o f t h e D i r e c t o r of t h e G e o l o g i c a l S u r v e y of W e s t e r n A u s t r a l i a , a n d t h e D i r e c t o r of t h e B u r e a u o f M i n e r a l R e s o u r c e s , Geology and Geophysics.
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&
ARRIENS,
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cambrian geochronology of Australia. Earth Sci., 5, 5 6 1 - 5 8 3 .
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COOPER, J . A . , J A M E S , P . R . , & R U T L A N D , R . W .
R.,
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1980: R b / S r dating of granitic intrusions in relation to the stratigraphic and deformational history of the Pilbara Block. Extended Abstr., 2nd Archaean Symp., Perth, 14-15. G., 1 9 7 2 : Granite ages within the Archaean Pilbara Block, Western Australia. J. geol. Soc. Aust., 19, 3 6 3 - 3 7 0 .
DE LAETER, J . R . , & B L O C K L E Y , J .
DE LAETER, J . R . , H I C K M A N , A . H . , T R E N D A L L , A .
F.,
& LEWIS, J. D., 1977: Geochronological data concerning the eastern extent of the Pilbara Block. Ann. Rep. geol. Surv. West. Aust. for 1976, 56-62. DE LAETER, J . R . ,
LEWIS, J .
D . , & BLOCKLEY, J .
G.,
1975: Granite ages within the Shaw Batholith of the Pilbara Block. Ann. Rep. geol. Surv. West. Aust. for 1974, 73-79. DE LAETER, J . R . , & T R E N D A L L , A . F . , 1 9 7 0 : T h e a g e o f
the Copper Hills Porphyry. Ann. West. Aust. for 1969, 5 4 - 5 9 .
Rep. geol.
: •—
EVENSEN, N . M . , HAMILTON, P . J . , & O ' N I O N S , R .
K.,
1978: Rare-earth abundances in chondritic meteorites. Geochim. cosmochim. Acta, 32, 1199-1212. G L I K S O N , A . Y . , 1 9 7 2 : Early Precambrian evidence of a primitive ocean crust and island nuclei of sodic granite. Bull. geol. Soc. Am., 83, 3 3 2 3 - 3 3 4 4 . GLIKSON, A . Y . , & HICKMAN, A . H . , 1 9 8 1 : G e o c h e m i c a l
stratigraphy and petrogenesis of Archaean basicultrabasic volcanic units, eastern Pilbara Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 287-300. HAMILTON, P . J . , O ' N I O N S , R . K . , & EVENSEN, N .
M.,
1977: Sm-Nd dating of Archaean basic and ultrabasic volcanics. Earth planet. Sci. Lett., 36, 263-268. HAMILTON, P .
J., O'NIONS,
R.
K . , EVENSEN, N .
M.,
B R I D G W A T E R , D . , & ALLAART, J . H . , 1 9 7 8 : Sm-Nd isotopic investigations of the Isua supracrustals, . and implications for mantle evolution. Nature, Lond., 272, 4 1 - 4 3 . HAMILTON, P . J . , EVENSEN, N . M . , O ' N I O N S , R . K . , &
DEPAOLO, D . J . , & W A S S E R B U R G , G . J . , 1 9 7 6 a : N d
topic variations and petrogenetic models. Res. Lett., 3, 249-252.
Surv.
, 1979: Sm-Nd age of the Stillwater Complex and the mantle evolution curve for neodymium. Geochim. cosmochim. Acta, 43, 999-1008.
iso-
Geophys.
, 1916b: Inferences about magma sources and mantle structure f r o m variations of 1 4 3 N d / 1 4 4 N d . Geophys. Res.. Lett., 3, 743-746.
TARNEY, J., 1979tf: Sm-Nd systematics of Lewisian gneisses: implications for the origin of granulites. Nature, Lond., 277, 25-28. HAMILTON,
P.
J.,
SMITH,
H.
S . , & ERLANK,
EVENSEN, N .
M., A.
O'NIONS, J.,
19796:
R.
K.,
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dating on Onverwacht G r o u p Volcanics, southern Africa. Nature, Lond., 279, 2 9 8 - 3 0 0 .
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HICKMAN, A. H. 1975: Precambrian structural-geology of part of the Pilbara region. Ann. Rep. geol. Surv. West. Aust. for 1974, 68-72. , 1981: Crustal evolution of the Pilbara Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 57-69. H I C K M A N , A . H . , & LIPPLE, S . L . , 1 9 7 8 : M a r b l e
Bar,
Western Australia— 1:250,000 Geological Series. Explan. Notes geol. Surv. West. Aust., SF50-8. JACOBSEN, S . B . , & WASSERBURG, G . J . , 1 9 7 8 : I n t e r p r e -
tation of Nd, Sr and Pb isotope data from Archaean migmatites in Lofoten-Vesteraalen, Norway. Earth planet. Sci. Lett., 41, 245-253. LUGMAIR, G. W . , & M A R T I , K . , 1 9 7 7 : Sm-Nd-Pu timepieces in the Angra dos Reis meteorite. Earth planet. Sci. Lett., 35, 2 7 3 - 2 8 4 . M I C H A R D - V I T R A C , A . , LANCELOT, J . , ALLEGRE, C . J . , &
U-Pb ages on single zircons from the early Precambrian rocks of West Greenland and the Minnesota River valley. Earth planet. Sci. Lett., 35, 4 4 9 - 4 5 3 . MOORBATH, S . , 1 9 7 7 :
O ' N I O N S , R . K . , HAMILTON, P . J . , & EVENSEN, N .
M.,
1977: Variations in 1 4 3 Nd/ , 4 4 Nd and 8 7 Sr/ 8 6 Sr ratios in oceanic basalts. Earth planet. Sci. Lett., 34, 13-22. OVERSBY, V . M . , 1 9 7 6 : Isotopic ages and geochemistry of Archaean acid igneous rocks from the Pilbara, Western Australia. Geochim. cosmochim. Acta, 40,
817-829.
PIDGEON, R. T., 1978tf: 3450m.y.-old volcanics in the Archaean layered greenstone succession of the Pilbara Block, Western Australia. Earth planet. Sci. Lett., 37, 421-428. , 1978&: Geochronological investigation of granite batholiths of the Archaean granite-greenstone terrain of the Pilbara Block, Western Australia (Abstract); in Smith, I. E. M., & Williams, J. G. (Eds) Proceedings of the 1978 Archaean Geochemical Conference, 360-362. Univ. Toronto, Ontario. J . R . , 1 9 7 8 : Lead isotopes and ages of galenas from the Pilbara region, Western Australia. J. geol. Soc. Aust., 24, 4 6 5 - 4 7 3 .
RICHARDS,
SANGSTER, D . F . , & BROOK, W . A . , 1977: Primitive lead
in an Australian Zn-Pb-Ba deposit. Nature, Lonci., 270, 423. STEIGER, R . H . , & J A G E R , E . , 1 9 7 7 : Subcommission on
geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth planet. Sci. Lett., 36, 3 5 9 - 3 6 2 . YORK, D., 1966: Least squares fitting of a straight line. Can. J. Phys., 44, 1079-1086. Z I N D L E R , A . , BROOKS, C . , A R N D T , N . T . , & HART, S.
R., 1978: Nd and Sr isotope data from komatiitic and tholeiitic rocks of Munro Township, Ontario. U.S.G.S. Open File Rep., 78-701, 469-471.
EVOLUTION OF THE GONDWANA ARCHAEAN SHIELD: ZIRCON DATING BY ION MICROPROBE, AND RELATIONSHIPS BETWEEN AUSTRALIA AND WILKES LAND (ANTARCTICA) J. F. Lovering, G. A. Travis, D. J. Comaford & P. R. Kelly 1
1
2
3
1
Department of Geology, University of Melbourne, Parkville, Victoria 3052, Australia Exploration Division, Western Mining Corporation Limited, P.O. Box 91, Belmont, Western Australia 6104, Australia Hasler Applications Laboratory, Applied Research Laboratories, P.O. Box 129, Sunland, California 91040, USA 2
3
ABSTRACT The ion-microprobe mass-analyser has been used to study P b / P b ratio variations between 15 fim sites on single zircon grains from coastal rocks on either side of the rift in the Gondwanaland Precambrian Shield between southwest Australia and Wilkes Land, East Antarctica. Zircons in gneisses from the Albany-Fraser Province in the Esperance region of Western Australia show apparent P b / P b ages ranging from a minimum of ~ 1600m.y. (in the highest uranium sites) to 3400-3500 m.y. (in the lowest uranium sites). These minimum apparent ages are identical with the results of conventional Rb/Sr dating of the time of highgrade metamorphism of these rocks, but the maximum apparent ages may reflect an ancient Archaean origin for some low-uranium zircons. Similar maximum apparent ages have been found by ion-microprobe studies in zircons from gneisses at Canning Dam on the western edge of the Archaean Yilgarn Block. Zircons from gneisses and associated rocks occurring in the Windmill, Donovan and Frazier Islands, and the East Antarctic mainland in the vicinity of Casey Station, Wilkes Land, show an identical range in apparent P b / P b ages (i.e. from -1600 to ~3500m.y.). These ages had not been recognized in previous studies. The close similarity in the total range of apparent 0 P b / P b ages determined in these zircons is strong evidence in support of those Gondwana reconstructions which place the Esperance region of Western Australia close to the Casey region of Wilkes Land in East Antarctica. 2 0 7
2 0 7
2 0 6
2 0 6
2 0 7
2
7
INTRODUCTION Lovering et al. (unpub. ms.) show how the unique spatial resolution of the ion-microprobe mass-analyser (IMMA) can provide information on apparent 207pb/ 206pb ages of 15 /xm diameter sites on individual zircon grains extracted from the Amitsoq gneisses from the Archaean of West Greenland. Their results show that individual 15/im analysis sites on single zircon grains retain a complex P b / 206pb isotopic-ratio history which can be interpreted as reflecting different apparent 207pb/206pb ages retained at each site, depending on its uranium content, and thus the degree of radiation damage suffered by the zircon lattice at that site. For example, the lowest 207pb/206pb apparent ages of 3000m.y. are found in the high-uranium sites and are taken to reflect the partial loss of radiogenic lead from these more metamict sites in response to later geological events, such as the intrusion of the Nuk 2 0 7
Spec. Pubis geol. Soc. Aust., 7 (1981)
2 0 6
206
gneiss parental rocks at 2800-2900m.y. On the other hand, it is suggested that the maximum apparent ages, of 3900-4000m.y. found in the low-uranium sites might reflect the existence in the Amitsoq gneisses of some zircons with crustal histories extending back to ~ 3900-4000m.y. The important implications of this study to the fundamental problem of the time of appearance of the Earth's earliest sialic crust led to an ionmicroprobe study of zircons from other ancient gneissic terrains. The present work describes observations on zircons from gneissic rocks exposed on the coastlines of southwestern Australia (in the Esperance region) and Wilkes Land, East Antarctica (in the region of Casey Station). The purpose of the study was to examine the detailed 207pb/ 206pb isotopic histories of zircons in both regions in an attempt to understand better the geochronological evolution of this part of the Gondwana Precambrian Shield, and also to
194
J. F. LOVERING, G. A. TRAVIS, D. J. COMAFORD & P. R. KELLY
examine the data in terms of detailed reconstructions of these parts of Australia and Antarctica into Gondwana before rifting and drifting. ION-MICROPROBE MEASUREMENTS OF 207pb/206pb RATIOS IN ZIRCONS A representative sample of the zircon population within each rock studied was extracted by conventional mineral separation techniques, a random selection of about 50 grains from each population was mounted in epoxy resin on glass, and a polished thin section mount was prepared. The distribution of uranium within individual zircon grains was then photographically documented using the fission track ' T e x a n " print method (Lovering et al,, unpub. ms.). Specific areas 15/mi across on individual zircon grains were selected for ion-microprobe analysis to cover the complete range of uranium concentrations found in each extracted zircon sample. The selected areas were then studied optically to ensure freedom from visible inclusions or fractures at, or close to, the analysis surface. The 207pb/ 206pb isotopic ratio at each analytical site was then measured using an Applied Research Laboratories ion-microprobe mass-analyzer (IMMA) under analytical conditions described by Hinthorne et al. (1979) and Lovering et al. (unpub. ms.). Briefly, the primary 16 0~ ion beam used had a 20 kV accelerating voltage, a beam size of 15/mi and a sample current of around 1.0 x 10~ 8 amps. Hinthorne et al. (1979) have described, in detail, methods developed to strip the important, complex, molecular-ion interferences from the lead-isotope peaks in the zircon secondary ion spectrum. Lovering et al. (unpub. ms.) give evidence in support of the accuracy of 2 0 7 Pb/ 206 Pb isotopic ratios in zircons measured by this technique but emphasize that the zircon site analysed must be relatively depleted in minor impurities, such as aluminium, magnesium and rare-earth elements (REE), otherwise complex molecular ion peaks interfere with the lead peaks and cannot be eliminated by normal stripping techniques. No further corrections have been applied to the calculated 207pb/206pb ratios for any possible common Pb present in the analysis sites, since no 204 Pb signal remained after correction for molecular-ion interference (see also Hinthorne et al., 1979 and Lovering et alunpub. ms.). If common Pb were indeed present at each analysis site, but its presence hidden by the correction procedure, then the calculated 207pb/206pb ages would be too old. We are, of course, aware that common Pb contamination of radiogeneic Pb is
observed in most conventional isotope-dilution studies of bulk zircon samples. However, we are also aware that the amount of such contamination (as measured by the 206pb/204pb ratio) normally decreases very markedly as the relative purity of the zircon concentrate increases. This would imply that most, if not all, common Pb contamination is introduced from contaminating phases occurring as solid and fluid inclusions and along fractures within, and surface coatings on, individual zircon grains. The obvious advantage of the ion-microprobe technique is that its unique spatial resolution allows the analyst to choose areas on individual zircon grains around 15/an across for analysis which are demonstrably free of such inclusions, fractures or coatings. On this basis, we contend that any possible common Pb contamination at each 15/mi analysis site is beyond our ability to measure within the constraints of the existing ion-microprobe techniques. Th/U atomic ratios (from 248Th +/254UO + ion ratios) and relative uranium abundances (from 254uo + /30Si + ion ratios) at each analysis site have also been determined. Uncertainties (la) for all corrected ^ T P b / ^ p b ratios have been calculated by a method given previously by Hinthorne et al. (1979) and are probably maximum assessments of the uncertainties. ESPERANCE REGION (SOUTHWESTERN AUSTRALIA) Geology, Tectonics and Geochronology The rocks of the southwestern coast of Australia in the Esperance region form part of the Albany-Fraser Province, a Precambrian mobile belt that sweeps around the southeastern and southern margin of the Archaean Yilgarn Block (Doepel, 1975; Gee, 1979). Structural trends of the Albany-Fraser Province truncate the northnorthwest trends of the Yilgarn Block at a high angle, and the boundary appears to be a gradational tectonothermal front termed by Gee (1979) the Fraser Front. The eastern boundary of the province is concealed by Tertiary sediments of the Eucla Basin. Migmatites, granitic gneisses and intrusive granitoids, the principal rocks in the province, display complex relationships most likely resulting from the superimposed metamorphism, tectonism, and igneous activity associated with several periods of basement-cover reactivation. Recognizable supracrustal sequences, dominantly shallow-water shelf-type sediments, occur intermittently in the province. They include the Stirling-Barren Beds in the western portion, which show a progressive increase in deformation
G O N D W A N A SHIELD: ZIRCON DATING BY ION MICROPROBE
Fig.. 1. . .
195
Tectonic outline map of the Albany-Fraser Province in the region around Esperance, southwest coast of Western Australia (after Gee, 1979). Sample localities for this study shown with abbreviated sample numbers.
from north to south, reflecting the transition to greater heat flow and basement mobility away from the Yilgarn Block. Turek & Stephenson (1966) obtained a Rb/Sr isochron age of 1150. ± 40m.y. on five shale samples from the Stirling Range. This was interpreted as a metamorphic age, and it is probably the youngest significant tectonothermal event in the province. Other supracrustal successions include the Mt Ragged Beds east-northeast of Esperance (Fig. 1), a sequence of quartzite, mica schist, quartzpebble conglomerate and felsic volcanics (Lowry & Doepel, 1974), and the Fraser Complex, a large synformal block of mafic granulite, with minor felsic granulite, gneiss, granite and metagabbro, forming the core of the Fraser Range (Tyrwhitt & Orridge, 1975; Gee, 1979). No radiometric dating Has been reported from the Mt Ragged Beds, but Compston & Arriens (1968) reported a Rb/Sr metamorphic age of 1330 ± 15m.y. for the Fraser Range granulites. Supracrustal remnants occasionally observed in the gneisses and migmatites attest to the presence of even older technically dismembered cover sequences. Gee (.1979) placed the rocks of the coastal stretch from Cape Pasley to Hopetoun into two general tectonic subdivisions—a Gneiss and Migmatite Zone and a Voluminous Granite Zone. The gneisses and migmatites, dominantly granitic in composition, appear to represent a tectonic interleaving of supracrustals and granitoids, this complex in turn being intruded at all scales by granitic rocks. Porphyritic adamellite, often
gneissic, and locally an augen gneiss, is the most common member of the Voluminous Granite Zone, and Gee (1979) shows these rocks as not only younger but also overlying the regional gneisses and migmatites. Arriens & Lambert (1969) report Rb/Sr ages ranging from 1590-1900 m.y. for a variety of rocks in the Gneiss and Migmatite Zone, and about 1100 m.y. for porphyritic adamellites at several localities in the province from Albany to Balladonia (Doepel, 1975). Richards et al. (1966) reported a U/Pb study of a zircon sample extracted from a metasedimentary gneiss near Ravensthorpe which indicated an upper intercept of concordia at about 2800m.y. Ion-microprobe Study of 207Pb/ 206Pb Zircon Ages In this preliminary study, samples from both tectonic subdivisions were analysed: 7710/109B is a supracrustal gneiss from the Gneiss and Migmatite Zone west of Esperance; 7710/101, 102, 112 are gneissic granites from the central Zone of Voluminous Granitoid; and 7710/81 is a supracrustal gneiss from the Gneiss and Migmatite Zone east of Esperance (Fig. 1) . All zircons extracted from the gneissic granites tend to be relatively large (150-300 /mi) and euhedral, typically with complex concentric growth zones and numerous inclusions. A large proportion of grains have relatively high uranium contents and those grains with uranium zoning usually show uranium-depleted cores and
196
J. F. L O V E R I N G ,
G. A. TRAVIS,
D. J. C O M A F O R D
& P.
R.
KELLY
TABLE I
Ion-microprobe analyses of zircons from supracrustal gneisses of the Gneiss and Migmatite Zone, Albany-Fraser Province, Esperance region Rock Number
Locali ty
7710/109B
Butty Head, Butty Harbour
7710/81
Point Malcolm
Zi rcon Gra i n Number 427 432 439 431 426
51>
U0+
Average Av. Av. Av. Av.
0.011 0.013 0.97 0.48 0.16
0.11 0.14 0.011 0.0088 0.0070
0.09701 ± 0.0060 0.1060 ± 0.0053 0.1546 ± 0.0173 0.2402 ± 0.0219 0.2491 ± 0.0334
Apparent Pb/20«Pb age (m.y. 1 1568 ± 121 1732 ± 92, 2398 i m 3122 ± 153 3180 ± 231
381 387 388
[ Core Av. Av. Rim
0.67 141 0.19 0.23 1.98
0.055 0.0055 0.084 0.074 0.043
380
Core Av.
1.46 158
0.0058 0.0084
0.1172 ± 0.0066 0.3022 ± 0.0281 0.1534 ± 0.0050 0.1752 ± O . O O 6 9 ( 0 . 1 8 6 9 ± 0.0177 { 0.1647 ± 0.0134 0.2736 ± 0.0276 0.2866 ± 0.0216
1915 ± 106 3482 ± 152 2385 ± 57 2608 ± 67 2716 ± 165 2505 t I4f 3327 ± 168 3^00 f 12|
Analys i s Site
2
*0ThO+/2
(Th/U)
385
uranium-enriched rims. Zircons from the two supracrustal gneisses form a distinctly different population in that they are rather smaller ( ~ 150/mi}, rounded or even irregular grains with a large proportion showing relatively low uranium contents. Growth zones are not prominent but many grains show uranium zoning with low-uranium cores and high-uranium rims. Ion-microprobe analyses of 15 /mi sites on individual zircons from each rock have been selected with regard to a range in relative uranium contents as shown by the range of 254yo + / 30Si + ion ratios at each site (Tables I & II). Apparent 207pb/206pb ages calculated at each site (Tables I and II) range from ~ 1600 to 3500m.y. and have then been plotted against relative uranium contents in Figure 2. These plots generally show an inverse relationship between apparent 207pb/ 206pb ages and relative uranium content at each site, the same correlation found in similar plots
25
W/30Si +
207pb/206pb
207
of ion-microprobe data on zircons from the AmTtsoq gneiss (Lovering et al., unpub. ms.). However, the geological interpretation of these apparent 207pb/ 206pb ages is extremely difficult without detailed knowledge of the 206pb/ 238y and 2 0 7 Pb/ 235 U ratios at each analysis site and their relationship to the concordia diagram. Classic studies by Wetherill (1956) and others have shown that fractionation of uranium and/or lead in zircons during their geological evolution can lead to major modifications in their 207pb/206pb ratios which can complicate the assignment of meaningful 4 'ages'' to these ratios. For example, it is well understood now that uranium gain and/or lead loss from a zircon during a subsequent geological event can lead to a reduction of the original 2 0 7 P b / 2 0 6 P b ratio and an apparent 207p b /206pb age which is less than the true age for that zircon sample. The amount of the apparent age decrease as recorded by the
TABLE
II
Ion-microprobe analyses of zircons from gneissic granites of the Zone of Voluminous granitoid, Albany-Fraser Province, Esperance region Rock Number
Locali ty
7710/101
High Island
7710/112
Termination Island
7710/102
Wylie Head
Zircon Grain Number 492 494 491
ThO+/2 S1,U0+ (Th/U)
2S4uo+/30Sj+
Average Av. Core
0.92 1.03 0.98
0,.015 0,.013 0..012
0.1602 ± 0.0118 0.1959 ± 0.0185 0.2149 ± 0.0162
age (m.y,) 2548 - 131 2793 ± 164 2943 ± 128
463 460 478
Av. Av. Av.
0.35 1.04 1.02
0..12 0.,013 0..0067
0.1716 ± 0.0099 0.2475 ± 0.0225 0.2780 ± 0.0348
2574 i-100 3169 ± 15:3 3352 ± 211
520 534
Av. 1 Rim I 1L Core Core 1k Rim Av.
0.61 0.32 0.87 0.55 1.39 0.38
0. 066 0. 15 0. 009 0. 047 0. 011 0. 086
0.1135 ± 0.0068 0.1170 ± 0.0051 0.2112 ± 0.0231 0.1242 ± 0.0090 0.2879 ± 0.0286 0.1250 + 0.0075-
530 531
Analys i s Si te
(
248
2o7
Apparent 207 pb/206pb
Pb/:?06pb
.
1856 t 113 1912 ± 81 2916 - 189 2017 ± 135 3407 t 16A 2029 ±110
GONDWANA SHIELD: ZIRCON DATING BY ION MICROPROBE
197 lost virtually all of their previously formed radiogenic lead during this same event. In this regard, it is highly significant that the lowest apparent P b / P b ages are generally associated with the sites containing the highest uranium abundances (Fig. 2). These uranium-rich zircon sites have suffered high lattice damage from a-bombardment and are metamict, so that radiogenic lead in these sites could diffuse or leach out during subsequent geological events. The significance to be attributed to the maximum apparent ages of 3400-3500m..y. is more difficult to assess. The only conventional zircon U/Pb study in the region is that by Richards et al, (1966) in which they proposed an upper intersect of concordia at about 2800 m.y. for a zircon sample from a metasedimentary gneiss collected 56 km east of Ravensthorpe in the Gneiss and Migmatite Zone. If the true ages of the zircons analysed in the present ion-microprobe study are also about 2800m.y., then the high apparent Pb/ Pb ages calculated here in certain sites could reflect radiogenic lead gain from a nearby site in the same grain relatively rich in radiogenic lead, and/or uranium loss from the site during some subsequent geological event. The best established subsequent event in the area is that delineated both by the previous total-rock Rb/Sr data and the present minimum apparent P b / Pb zircon ages by ion microprobe at around 1600m.y. However, the chord to concordia joining 1600m.y. to 2800m.y. is essentially parallel to the chord between 3500m.y. and the origin, so that the measured Pb/ Pb isotopic ratios for the apparent 3500m.y, old zircon sites cannot have been formed by any reasonable scenario of lead gain and/or uranium loss from a 2800m.y.
measured P b / P b ratio is a function of the proportion of uranium gained and/or lead lost during the event. Similarly it can be shown that uranium loss and/or lead gain in a zircon during a geological event subsequent to its formation or true age will lead to an increase in the original P b / P b ratio. A P b / P b age calculated on the basis of this modified P b / P b ratio will be higher than the true age of the zircon by an amount which is a function of the proportion of uranium lost and/or lead gained. The conventional way to examine the effect of uranium-lead fractionation in zircon evolution requires the determination of P b / U and 20 Pb/ U isotopic ratios for each sample analysed and the plotting of these data points on the concordia diagram. This is not possible for the present study since U/Pb ratios cannot be measured using existing ion-microprobe analytical techniques (Lovering et al., unpub. ms.). However, inspection of the concordia curve does allow certain conclusions. The youngest apparent P b / P b ages preserved in zircons from these first studies of gneisses and gneissic granites from the AlbanyFraser Province rocks in the Esperance region are in the range of 1600 to 1800m.y. These data are in good agreement with previous total rock Rb/Sr ages for rocks in the Gneiss and Migmatite Zone reported by Arriens & Lambert (1969). Consequently the simplest interpretation of the youngest apparent P b / P b ages reported in the present ion-microprobe study is that they represent either zircon grains which were formed during a geological event at this time or else particular sites in pre-existing zircon grains which 207
206
207
207
207
206
206
7
206
206
238
235
207
207
206
207
206
207
206
207
206
206
207
ESPERANCE
i
206
AREA. WESTERN AUSTRALIA Gneissic granite. High Island Gneissic graniiu. Termination Is. Gneissic granite. Wylie Head Supracrustal gneiss. Pt Malcolm Supracrustal gneiss. Buttv Head
CC '
001
\A
0001
16
1-8
20
22
26
2-4
28
32
30
3-4
36
38
Zircon 2 0 7 Pb/ 2 0 6 Pb Age(x 109years)
Fig. 2. Relative uranium concentration (254uo V S i ion ratio) vs 207 /206 b apparent age measured by ion microprobe on individual - 15 /^m.sites with zircons from gneisses and gneissic rocks from the Esperance region, Western Australia. 3 0
+
Pb
P
198
J. F. LOVERING, G. A. TRAVIS, D. J. COMAFORD & P. R. KELLY
old zircon during a 1600m.y. geological event. Such high 207pb/206pb ratios could only have been formed in a 2800m.y. old zircon if it suffered a massive radiogenic lead gain and/or uranium loss in the period from 2800m.y. to about 2000m.y., for which there is no independent evidence. Another possible explanation of the maximum apparent 207pb/206pb ages around 3500m.y. follows from the observation that these maximum apparent ages are normally associated with the lowest uranium sites in the zircons. These sites show little evidence of any lattice damage as a result of a-bombardment so that the radiogenic lead accumulating in them might well be retained without significant isotopic modification by either radiogenic lead gain and/or uranium loss during even severe metamorphic processes such as the gneiss-forming events in the region. On this basis, these low-uranium zircons in the AlbanyFraser Province gneissic terrain could represent — 3500m.y. old zircons derived from the adjacent Archaean rocks of the Yilgarn Block or else indicate that the Albany-Fraser Province gneissic rocks are simply reworked Archaean of the Yilgarn Block. Whichever of these two interpretations is favoured, the Archaean Yilgarn terrain is indicated as a possible ultimate source for these uranium-poor zircons. In this respect, it is of interest that the oldest apparent 2 0 7 Pb/ 2 0 6 Pb ages determined by ion microprobe on low-uranium zircons from the complex gneisses of the West Yilgarn Gneiss Domain in the Canning Dam Quarry about 32 km. southwest of Perth are also around 3500m.y. (Lovering, unpub. ms.). CASEY-WINDMILL ISLANDS REGION (WILKES LAND, EAST ANTARCTICA) Geology, Tectonics and Geochronology The region of the Antarctic mainland in the vicinity of Casey Station and the various groups of islands—Windmill, Donovan and Frazier Islands—adjacent to Casey Station consists of a layered sequence of schist, gneiss and migmatite (the Windmill Metamorphics) intruded by a charnockite (the Ardery Charnockite), a younger porphyritic granite (the Ford Granite), and various mafic dykes (Fig. 3). Blight & Oliver (1977) have mapped the Windmill Islands and mainland outcrops in the region of Casey Station. The Windmill Metamorphics on the Donovan and Frazier Islands have been mapped by two parties from the Department of Geology of the Universit v of Melbourne in austral summers 1977-78 (J. F. Lovering & G. A. Travis) and 1978-79 (A C Griffin & I. R. Duddy).
Blight & Oliver (1977) have shown that the metamorphic grade of the Windmill Metamorphics ranges from upper amphibolite facies in the north, through biotite-cordierite-almandine granulite to hornblende-orthopyroxene granulite in the south. They also concluded on geochemical grounds that the Windmill Metamorphics probably represents a supracrustal succession of acid and basic volcanics interbedded with sediments ranging from greywacke-type sandstone to shale. Partial melting of these rocks is thought to have produced some of the more acidic gneisses in the area. The geochronology of the region has been studied by a number of workers. Cameron et al. (1960) reported K/Ar ages of biotite from a gneiss from Haupt Nunataks (950 m.y.), a migmatite gneiss from a small unnamed island about 8km north of Wilkes Station (1120m.y.), and a charnockite from Ardery Island (1100 m.y.). The biotite from the migmatite gneiss also indicated a Rb/Sr age of 1110m.y. Webb et al. (1963) carried out a K/Ar age study of biotites from 5 gneiss samples from the Windmill Metamorphics and a sample of the Ford Granite and reported consistent ages of 1050-1140m.y. Later total-rock Rb/Sr studies in the area by Arriens (1975) yielded isochrons ranging between 1100 and 1400 m.y. According to Blight & Oliver (1977), Arriens believes the 1400m.y, date represents the pre-metamorphic age of sedimentation of the parental rocks to the Windmill Metamorphics, and that the 1100m.y. date corresponds to the subsequent high-grade metamorphic event. Ion-microprobe Study of ^Pb/2^Pb Zircon Ages For this study samples of various rock types within the Windmill Metamorphics were collected from both the southern high-grade metamorphic area and the lower-grade area in the north (Fig. 3 and Table III). A sample of the Ford Granite from Cloyd Island was also studied, but zircons extracted from a specimen of the Ardery Charnockite (7810/103), collected from the southwest side of Browning Peninsula, proved to be too high in REE and too low in radiogenic lead for satisfactory ion-microprobe analysis. Most of the gneisses from the Windmill Metamorphics (7810/28, 06, 21, 43) contain rather small 50-200/mi), often rounded zircons with relatively poorly developed growth zones and simple uranium zoning usually with low-uranium cores and high-uranium rims. Very similar zircons were found in the Ford Granite sample 7810/25. The zircons from 7810/41, the layered granite gneiss from Clark Peninsula, are very dif-
G O N D W A N A SHIELD: Z I R C O N DATING BY ION M I C R O P R O B E
Fig. 3.
199
Generalized geological map of the Windmill, Donovan and Frazier Islands and the regjgft arpund Casey Station, Wilkes Land, East Antarctica (based on Blight & Oliver, 1977). Sample loc^fffp for this study shown with abbreviated sample numbers.
200
J. F. L O V E R I N G , G . A. T R A V I S , D. J . C O M A F O R D & P . R. K E L L Y TABLE
III
Ion-microprobe analyses from the Ford Granite and Gneisses of the Windmill Metamorphics in the Casey Station region, Wilkes Land, East Antarctica Analysis Site
-uo+/ 30 S j +
-
195 193 197
Average A v. Av.
0.09 0.99 1.14
0.040 0.013 0.018
0 . 1 3 2 5 ± 0.0074 0 . 1 9 8 9 ± 0.0149 0.2396 ± 0.0161
2 1 3 2 ± 101 2817 ± 129 3 U 8 ± U2
7810/28
231 227 241 246
Core Av.. Hi-U end Av.
0.12 0.24 0.31 0.46
0.033 0.076 0.036 0.011
0.1457 ± 0.0117 0.1719 ± 0.0058 0.1734 ± 0.0074 0.2355 ± 0.0241
2296 ±. 145 • 2577 i 58 2591 ± 73 3090 ± 174
Migmati tic gneiss, Nelly Island, Frazier Islands
7810/06
266 263 260
Av. Av. Av.
0.04 0.04 1.18
0.033 0.024 0.01'2
0.1238 ± 0.0061 0.1412 ± 0.0077 0.2363 ± 0.0137
2013.±. 91 2242 ± 98 3096 ± 96
Layered granite gneiss, Clark Peninsula
7810/41
82 81 87 80
Rim Av. Rim Av.
0.12 0.29 0.21 .0.30
0.14 0.065 0.11 0.041
0.1824 ± 0.0052 0.2124 ± 0.0086 0.2404 ± 0 . 0 0 8 7 0.2847 ± 0.0131
2675 ± i»8 2924 ± 67 3 1 2 3 4 59 3390 * %
Granite gneiss, Haupt Nunataks
7810/21
304
f Rim 1 Core f Rim [ Core Av. Av. Core Core
0.16 0.81 0.10 0.07 1.09 1.13 0.77 f 0.26 I 0.17
0.090 0.007 0.029 0.034 0.037 0.016 0.018 0.014 0.0073
0.1059 0.1795 0.1569 0.1936 0.1829 0.1839 0.2278 0.2864 0.2976
± ± ± ± ± ± ± ± ±
0.0044 0.011.7 0.0157 0.0161 0.0094 0.0099 0.0137 0.0286 0.0182
1729 ± " 79 2649 ± 113 2423 ± 181 2773 ± 143 2680 ±" 88 2689 ± 32 , 3037 ± 100 3399 165 3458 ± 98
Av. ( Rim { Core Av. Core
0.03 0.02 0.60 0.36 0.56
0.020 0.084 0.013 0.10 0.015
0.0999 0.1097 0.2282 0.1129 0.2271
± ± ± ± ±
0.0040 0.0057 0.0115 0.0054 0.0160
1622 ± 76 1794 + 98 3040 ± 84 1847 ± 90 303? ± 1 1 8
Av. Av. Av. Av.
0.84 0.50 0.61 1.14
0.058 0..053 0.018 0.012
0.1150 0.1184 0.1945 0.2423
i ± ± ±
0.0049 0.0057 0.0123 0.0207
1881 ± 79 1933 ± -90 2781 ± 107 3136 ± 143
Cloyd
Ford Granit_e
Windmi1] Metamorphics
2 8 * Th0+/ 2S.U0+
Z i rcon Gra in Number
Rock Number
Rock Type and Locali ty
Rock Series
Island
7810/25
Gnei ss, Chappel Island, Donovan Islands
307 293 308 306 300 Mi gmat i t i c gnei ss, errat i c, Mi tchel1 Peni nsula
160 165
7810/43
• 161 164 Mafic dyke, Mi tchel1 Peni nsula
7810/104
325 326 321 337
ferent from those in the other gneisses f r o m the Windmill Metamorphics in that they are larger ( - 3 0 0 / x m ) and euhedral with very strongly developed concentric growth zoning. The rather rare zircons extracted f r o m the mafic dyke (7810/104) are highly angular and seem to represent fragments of large crystals fractured during the extraction process.
(Th/U)
,
25
Apparent Pb/*« 6 P b age (m.y.)
207pb/206pb
207
15 fini sites on individual zircon grains for ion^ microprobe analysis were chosen using the same criteria as discussed previously for the zircons from the Esperance rocks. Analytical data from the ion microprobe, and calculated 207pb/206pb apparent ages at each site, are shown in Table III. In Figure 4 these 207pb/ 206pb apparent ages have been plotted against relative uranium contents
WINDMILL ISLANDS - CASEY A R E A O
7810143
Migmatitic gneiss erratic 1 h Mitchell Pen. '
Mafic dyke
10
Gneiss. Chappel Is Migmatitic gneiss, Nelly Is Granite gneiss Haupt Nunataks Ford granite. Cloyd Is. Layered granite gneiss, Clark
Peninsula
0-1
0 01 300C(1)
16
1 * 2 ¥
2-2
24
207 ,
2-6
2^8~
30
32
3-4
36
/ 206r
n Zircon 207Pb / 206 Pb Age (x 109years)
Fig. 4.
+
Relative uranium concentration (254 U O / 30Si + ion ratio) vs 2 0 7 P b / 2 0 6 P b a p p a r e n t a g e m e asured by ion TrZ ^ n °n c V ~ 1 5 ^ m s i t e s w i t h i n z i r c o n s f r o ™ Ford Granite and Windmill Metamorphics from the Casey Station region, East Antarctica.
GONDWANA SHIELD: ZIRCON DATING BY ION MICROPROBE
(from 254uo + /30Si + ion ratios). As with Esperance zircons, there is a strong inverse relationship between apparent 207pb/ 206pb ages and relative uranium content at each site. The minimum 207pb/ 206pb apparent ages recorded in the zircons from the Windmill Metamorphics fall in the range 1600-1800 m.y. and occur in individual grains with the highest uranium contents. These minimum ages are significantly older than the 1400m.y. maximum Rb/Sr whole-rock ages of Arriens (1975) discussed above but completely agree with the minimum 2 0 7 Pb/ 2 0 6 Pb ion-microprobe apparent ages reported here for zircons from the gneisses around the Esperance region in Western Australia. -The maximum apparent 207pb/206pb ages in zircons found by this ionmicroprobe study occur in the lowest uranium sites and fall in the range 3400-3500m.y. and mirror exactly the maximum apparent ages found in the zircons from the Esperance gneisses. The zircons from the Windmill Metamorphics do not seem to record specifically the later —1100m.y. event indicated by the biotite K/Ar and Rb/Sr whole-rock studies previously discussed. Zircons from the Ford Granite, which intrudes the Windmill Metamorphics, show apparent 207pb/206pb ages which fall within the range of those observed in zircons from the Windmill Metamorphics (Table III, Fig. 4). These observations support the view that the Ford Granite magma was derived from the partial remelting of Windmill Metamorphics country rock (Blight & diver, 1977) and so might well contain relict zircons from those parental rocks. However those mechanisms responsible for the variations in apparent 2 0 7 Pb/ 2 0 6 Pb ages in the Esperance zircons described before are also presumably responsible for the exactly parallel variations in the Casey zircons. AUSTRALIA-ANTARCTICA RELATIONSHIPS WITHIN GONDWANA A ustralia-A ntarctica Reconstructions Since the earliest reconstructions of the ancient southern super continent of Gondwana, the southern rim of Australia has fitted rather neatly around the keystone of East Antarctica. However, recently published predrift reassemblies of Australia and Antarctica show significant disagreement in the detailed positioning of one continent with respect to the other. At one extreme there are the computer-generated, essentially identical fits on the 1000 fathom isobath (Sproll & Dietz, 1969) and the 500 fathom isobath (Smith & Hallam, 1970) which place the Casey region in East Antarctica relatively close to Albany on the southwest coast of Western Australia (Fig. 5A).
201
At the other extreme Griffiths (1974) has proposed a fit based on "reversing" sea-floor spreading using a mapped pattern of magnetic lineations and fracture zones. A similar reassembly has been suggested by Laird et al. (1977) on the basis of new stratigraphical and palaeontological data from the Early Palaeozoic Bowers Group of northern Victoria Land (Antarctica), which indicates that its depositional basin was probably closely linked with the Dundas Trough of Western Tasmania. This alternative reassembly (Fig. 5B) has the added advantage of accommodating the South Tasman Rise with almost no overlap adjacent to the eastern coast of northern Victoria land; it also places the Casey region close to, and slightly east of, Esperance on the southwest coast of Western Australia. Both reassemblies show a substantial gap between the 2000 metre isobaths in the Great Australian Bight. Konig & Talwani (1977) have linked this area with a magnetic quiet zone and suggested that this zone represents non-oceanic subsided basement. If so, the southern boundary of the Australian plate should lie somewhere to the south of the southern boundary of the magnetic quiet zone, which would substantially account for the apparent under lap between the 2000 metre isobaths in the Great Australian Bight. The present study has indicated a remarkable similarity in the apparent 2 0 7 Pb/ 2 0 6 Pb zircon age ranges of -1600 to -3500m.y. determined by ion microprobe on rocks from both the Casey and Esperance regions. If these data are taken in conjunction with the added coincidence of a later — 1100m.y. old event recorded by conventional geochronological techniques in both areas, they provide further evidence in support of those Gondwana reconstructions, such as illustrated in Figure 5B, which place the Esperance region of southwestern Australia in close association with the Casey region of Wilkes Land in East Antarctica. However, without similar ion-microprobe data on zircons from rocks between Esperance and Cape Leeuwin in southwestern Australia it is not yet possible to make an unequivocal choice between the two reconstructions illustrated in Figure 5. These studies are in progress. Evolution of the Gondwana Precambrian Shield Whichever reassembly of Australia and Antarctica is chosen, it is clear that the rocks from both the Esperance and Casey regions have evolved within the Precambrian Shield core of Gondwana and their geochronological history should help to define the course of evolution of that shield. Although it is not yet possible to propose an unequivocal explanation of the observed 207pb/
202
J. F. LOVERING, G. A. TRAVIS, D. J. COMAFORD & P. R. KELLY
Fig. 5. Reconstructions of Australia and East Antarctica within Gondwana. A. Computer fit reassembly of continental margins after Sproll & Dietz (1969) and Smith & Hallam (1970). B. Reassembly based on reversal of magnetic lineation data (Griffiths, 1974) and geological matching between Tasmania and northern Victoria Land (Laird et al., 1977). Latitudes and longitudes refer only to present day Antarctica. Kanmantoo Group (South Africa) not shown. Continental margin at 2000 metre isobath. AG: Adelaide Geosyncline; SB: Smithton Basin; DT: Dundas Trough; AT: Adamsfield Trough; BG: Bowers Group.
GONDWANA SHIELD: ZIRCON DATING BY ION MICROPROBE 206pb ratio data on the uranium-poor sites in zircons from both the Esperance and Casey regions, one possible interpretation is that these particular zircons were derived f r o m Archaean crustal terrains of sialic composition which existed within, or in close association to, these areas 3400-3500 m.y. ago.
ACKNOWLEDGMENTS This study received m a j o r financial support from the Australian Research Grants Committee
203
and other support from the Australian Institute of Nuclear Science and Engineering by way of grants to one of us (J. F. Lovering). The Antarctic Division of the Department of Science and Technology provided two of us (J. F. Lovering and G. A. Travis) with vital logistic support to collect samples f r o m the Casey Station region during the 1977-78 austral summer. All ionmicroprobe analyses were made on the I M M A ion microprobe made available by J. R. Hinthorne and Applied Research Laboratories in their Sunland plant.
REFERENCES ARRIENS, P. A., 1975: The Precambrian geochronology of Antarctica. Abstr., 1st Aust. geol. Convention, 97-98. ARRIENS, P . A . , & LAMBERT, I. B . , 1 9 6 9 : O n t h e a g e a n d
strontium isotopic geochemistry of granulite-fades rocks from the Eraser Range, Western Australia and the Musgrave Ranges, Central Australia. Spec. Pubis geol, Soc. Aust., 2, 377-388. L . , 1 9 7 7 : ' T h e metamorphic geology of the Windmill Islands, Antarctica: a preliminary account. J. geol. Soc. Aust., 24,
BLIGHT, D . F . , & OLIVER, R .
239-262.
LAIRD, M . G . , COOPER, R . A . , & JAGO, J . B . ,
1977:
New data on the lower Palaeozoic sequence of northern Victoria Land, Antarctica, and its significance for Australian-Antarctic relations in the Palaeozoic. Nature, Lond., 265, 107-110. C., & DOEPEL, J . J . G., 1 9 7 4 : MalcolmCape Arid, Western Australia—1:250 000 Geological Series. Explan. Notes geol. Surv. West. Aust., SI51-7.
LOWRY, D .
RICHARDS, J . R . , BERRY, H . , & RHODES, J . M . ,
1966:
Isotopic and lead-alpha ages of some Australian zircons. J. geol. Soc. Aust., 13, 6 9 - 9 6 .
CAMERON, R . L . , GOLDICH, S . S . , & HOFFMAN, J . H . ,
1960: Radioactivity age of rocks from the Windmill Islands, Budd Coast, Antarctica. Stockh. Contr. Geol., 6, 1-6. COMPSTON, W . , & ARRIENS, P .
A.,
1968: T h e
Pre-
cambrian geochronology of Australia. Can. J. Earth Sci., 5, 5 6 1 - 5 8 3 . DOEPEL, J. J. G., 1975: Albany-Fraser Province; in Geology of Western Australia. Mem. West. Aust. geol. Surv.,2, 94-102. GEE, R. D., 1979: Structure and tectonic style of the Western Australian Shield. Tectonophys., 58, 327-369. GRIFFITHS, J. R., 1974: Revised continental fit of Australia and Antarctica. Nature, Lond., 249, 336-338. HINTHORNE, J . R . , ANDERSON, C . A . , CONRAD, R . L . ,
& LOVERING, J. F., 1979: Single-grain 2 0 7 Pb/ 2 0 6 Pb and U/Pb age determinations with a 10-^m spatial resolution using the ion microprobe mass analyzer ' ( I M M A ) . Chem. Geol., 25, 271-303. KONIG, M., & TALWANI, M . , 1977: A geophysical study
of the southern continental margin of Australia: Great Australian Bight and western sections. Bull, geol. Soc. Amer., 88, 1000-1014.
SMITH, A . G . , & HALLAM, A . ,
1970: T h e f i t o f
the
southern continents. Nature, Lond., 225, 1 3 9 - 1 4 4 . Morphological continental drift fit of Australia and Antarctica. Nature, Lond., 222, 345-348.
SPROLL, W . P . , & DIETZ, R . S . , 1969:
TUREK, A . , & STEPHENSON, N . C . N . , 1966: T h e r a d i o -
metric age of the Albany Granite and the Stirling Range Beds, southwest Australia. J. geol. Soc. Aust., 13, 449-456. TYRWHITT, D . S . , & ORRIDGE, G . R . , 1975: R e g i o n a l
geology and mineralization of the Fraser Range Oogenic Belt, Western Australia; in Knight, C. L. (Ed.) Economic Geology of Australia and Papua New Guinea. 1. Metals, 405-408. Australas. Inst. Min. Metall., Melbourne. WEBB, A . W . , MCDOUGALL, I . , & COOPER, J . A . , 1963:
Potassium-argon dates from the Vincennes Bay region and Oates Land; in Adie, A. J. (Ed.) Antarctic Geology, 5 9 7 - 6 0 0 . North Holland, Amsterdam. WETHERILL, G. W., 1956: Discordant uranium-lead ages, I. Trans. Am. geophys. Un., 37, 320-326.
GEOCHRONOLOGY AND EVOLUTION OF LATE ARCHAEAN GNEISSES IN NORTHERN LABRADOR: AN EXAMPLE OF REWORKED SIALIC CRUST Kenneth D. Collerson, 2 Andrew Kerr* and William Compston
2
Department of Geology, Memorial University of Newfoundland, St John's, Newfoundland, Canada 2 Research School of Earth Sciences, Australian National University, Canberra, Australia 1
ABSTRACT Several groups of late Archaean quartzo-feldspathic gneisses have been recognized in the Nachvak-Saglek-Hebron area of Northern Labrador. The field relations of one of these groups (the Kiyuktok gneisses) indicate derivation by partial anatexis of the earlier Uivak gneisses. The Kiyuktok gneisses yield Rb-Sr ages of ca 2800 Ma (Sr = 0.7086) but preserve a whole-rock PbPb age of ca 3500 Ma which is broadly comparable with Rb-Sr ages from the Uivak gneisses. Late Archaean orthogneisses derived from discrete intrusive bodies (Ikarut and Kammarsuit gneisses) also yield Rb-Sr ages of ca 2800 Ma (Sr = 0.7057-0.7059). A subordinate group of tonalite and granodiorite sheets in the Saglek area are of equivalent age but have a significantly lower initial ratio (0.7031-0.7037). However, they are petrologically and isotopically comparable to the chronologically equivalent Nuk gneisses in the Godthaab area of West Greenland, which are regarded as "juvenile" mantle-derived sialic crust. We consider the higher initial ratios displayed by the Ikarut and Kammarsuit gneisses to reflect interaction between preexisting sialic crust and ascending "juvenile" Nuk-type magmas. Such interaction may be a result of bulk magma mixing or alternatively contamination of ascending magmas by hydrothermal (crustal) fluids enriched in radiogenic Sr. This interpretation is consistent with our recognition of regionally reworked sialic crust in the form of the Kiyuktok gneisses and also with modern thought regarding the possible roles of magmatic and/or fluid mixing in the generation of calc-alkaline suites. We therefore suggest that processes of crust formation and differentiation in the North Atlantic Craton during the late Archaean, may be more complex than previously supposed. 0
0
INTRODUCTION The pre-3500Ma history of Archaean rocks in the North Atlantic Craton is now fairly well established (Black et at., 1971; Moorbath et al., 1972, 1975a, 1977; McGregor, 1973; Bridgwater et al., 1975; Bridgwater & Collerson, 1976, 1977; McGregor & Mason, 1977; Bridgwater et al., 1978; Hamilton et al.. 1978; Collerson & Bridgwater, 1979; Chadwick & Nutman, 1979; Collerson & Compston, unpublished data). However, these early Archaean enclaves are only of relatively limited extent and most of the North Atlantic, Craton consists of Late Archaean quartzofeldspathic gneisses, which range in age from 2600 to 3000 Ma (Pankhurst et al., 1973; Pidgeon & Hopgood, 1975; Moorbath & Pankhurst, 1976; Baadsgaard, 1976; Taylor & Moorbath, 1978). Rocks of this age are also abundant in other Archaean terrains which suggests that this time interval was a globally important period of crust formation.
Spec. Pubis geol. Soc. Aust, 7 (1981)
Isotopic studies of late Archaean quartzofeldspathic gneisses from West Greenland and Scotland by Moorbath et al. (1915b), Moorbath & Pankhurst (1976), Chapman & Moorbath (1977) and Hamilton et al. (1979) have supported their interpretation as new or juvenile sial. A number of examples of reworked Archaean sialic crust (i.e. crust which has experienced 'partial or complete melting leading to mobilisation and reconstitution as an essentially new rock", Moorbath, 1975a) have also been recently documented isotopically by Oversby (1975, 1978), Davies & Allsopp (1976) and Hickman (1978). Despite these, reworking as a viable, large-scale, crustforming or modifying process has generally been discounted from most Archaean crustal evolutionary models (see for example, Moorbath, 1975a, 1975Z?, 1976, 1977). .This paper presents results from several contrasting groups of late Archean gneisses in, the Nachvak to Hebron area of the Nain Province, 4
206
K. D. COLLERSON, A. KERR & W. COMPSTON Nachvak
Fiord
N
L LA BR A DOR
SEA LEGEND Early Archaean amphibolite facies gneisses Area affected by late Archaean events \.ooJ>
Ramah Group (lower Proterozolc) Churchill Province Structural trend- line
Hebron
0
L_
Fault
20 km i wc. m
Fig. 1.
Geological map of northern Labrador. The locations of the Churchill Province boundary and the Rarnah Group are taken from Morgan (1975).
Northern Labrador (Fig. 1). Field studies, petrology, and isotopic (Rb-Sr and Pb-Pb) evidence independently and collectively indicate that the formation and stabilization of this crust involved the addition of new material and reworking of pre-existing sial, and possibly mixing of both. The paper is in three parts. The first reviews field relations and petrology of the late Archaean rocks. The second section deals with whole-rock Rb-Sr and Pb-Pb isotopic systematics of these gneisses and the final part discusses aspects of late Archaean crustal evolution in the light of this evidence. FIELD RELATIONS AND PETROLOGY Characteristics of the Pre-3000Ma Gneiss Complex A knowledge of the lithology and evolution of the pre-3000 Ma gneiss complex in Northern Labrador is essential for the interpretation of late Archaean events. Aspects of this early history are discussed in Collerson et al. (1976), Bridgwater & Collerson (1976, 1977), Bridgwater et al. (1978)
and Collerson & Bridgwater (1979) and will not be treated in detail in this paper. However, a brief review of the major pre-3000 Ma rock-types fe presented below. Early Archaean rocks are best preserved within the amphibolite-facies area along the coast between Saglek Bay and Hebron Fiord (Fig. 1). Ifl this area, two lithostratigraphic sequences of different age are identified by the presence or absence of distinctive meta-diabase intrusions, termed Saglek dykes. Using this criterion, a detailed chronology has been established which is similar in most respects to that developed by McGregor (1973) in the Godthaab district of West Greenland (Table I). About 75% of the pre-3000 Ma gneiss complex is composed of early Archaean (pre-3500Ma) quartzo-feldspathic gneisses of trondhjemitic, tonalitic and granitic composition, termed the Uivak gneisses. These are distinguished from younger gneisses of similar appearance by the presence of Saglek dykes. In areas of low finite strain, intrusive relationships between these and .the Uivak gneisses are preserved.
G N E I S S E S IN N O R T H E R N
LABRADOR
207
TABLE I
Simplified table ofArchaean events in the North Atlantic Craton Date Deposi t ion of I sua supracrustal s and Aki lia . association
(ca 3750 Ma)
Intrusion of granitic-tonal itic parents of the AmT-tsoq gneisses
Deposition of Nulliak assemblage High-grade metamorphism (granulite facies)
(ca 3600-3700 Ma)
Emplacement of granitic-tonalitic parents of the Uivak gneisses Deformation and metamorphism
Emplacement of Fe-rich diorites and quartz monzoni tes
Emplacement of Fe-r.ich diorites and quartz monzonites
Deformat ion and metamorphi sm
(ca 3500-3600 Ma)
Intrusion of Ameral ik dykes
(ca 3200-3^00 Ma)
Intrusion of Saglek dykes
Extrusion and deposition of Malene supracrustals
(pre-3000 Ma)
Extrusion and deposition of Upernavik supracrustals
Emplacement of- layered gabbroic and anorthositic bodi-es into supracrustals
Emplacement of layered gabbroic and anorthositic bodies into supracrustals
Intercalation of Amitsoq gneisses, Malene supracrustal s, and anorthosites Intrusion of tonal i tes, granodior ites (Nuk gneisses) -
Deformation and metamorphism (retrogression of higher grade assemblages to amphiboli te facies)
Intercalation of Uivak gneisses and Upernavik supracrustals and metamorphism. Emplacement of ultramafic slices (ca 2700-3000 Ma)
Intrusion of tonalite-granodiorite sheets and plutons (Ikarut, Kammarsui t gneisses)
Deformation and granulite-facies metamorphi sm
(ca 2700-2900 Ma)
High-grade metamorphism and formation of Kiyuktok gneisses
Formation of post-tectonic, K-rich granites (Qar.qut granite) ,
(ca 2520 Ma)
Formation of post-tectonic, K-rich granites (Igukshuak granite)
The Uivak suite is dominated by polydeformed layered orthogneisses of tonalitic, trondhjemitic and granodioritic composition, termed the Uivak I gneisses, which are interleaved with, and locally cut by, subordinate amounts (5 to 10%) of more homogeneous augen gneisses (the Uivak II gneisses). These are more potassic than the Uivak I gneisses and exhibit a prominent iron-enrichment trend (Collerson & Bridgwater, 1979). Both sub-groups of the Uivak suite contain remnants of older supracrustal rocks (dominantly banded iron formation and amphibolites of ultrabasic to basic composition) termed the Nulliak assemblage. All the above rock-types were strongly deformed and metamorphosed before the intrusion of the Saglek dykes ca 3200 to 3400 Ma ago (Collerson, unpublished data). The remaining 20 to 30% of the 3000-3500 Ma old gneiss complex is formed by a varied group of rocks known as the Upernavik supracrustal suite. These are dominantly pelitic to psammitic metasediments (probably of volcanic derivation), with associated layered amphibolites and cut by subordinate meta-gabbroic and meta-leucogabbroic rocks. The Upernavik supracrustal suite does not contain Saglek dykes and is thus interpreted as a younger cover-sequence developed on a Uivak gneiss basement. Alternatively, it may represent a completely allochthonous assemblage which developed elsewhere and was subsequently interleaved with the Uivak gneisses by thrusting (cf. Bridgwater et al., 1974). Contacts between the two groups of rocks are commonly marked by disrupted ultramafic bodies, which locally pre-
serve garnetiferous assemblages indicative of upper mantle conditions (Collerson & Kerr, unpublished data). Microstructural and mineralogical evidence suggests that these high-pressure assemblages developed before emplacement, which indicates that the ultramafic rocks are allochthonous. Field Relations and Petrology of the Late A rchaean Gneisses The most obvious indication of late Archaean activity was the formation of a heterogeneous group of rocks, termed the Kiyuktok gneisses. The gneisses range from recognizable derivatives of pre-existing rocks, through structureless nebulites and ultimately to coarse-grained intrusive pegmatites and granitic sheets containing garnet and orthopyroxene, Other late Archaean gneisses recognized in the area are narrow sheets of finegrained tonalitic to granodioritic gneiss, and coarse-grained bodies of dominantly granodioritic composition (termed the Ikarut and Kammarsuit gneisses respectively). Relationships between the tonalitic and granodioritic gneisses and the Kiyuktok gneisses are somewhat equivocal although they probably formed penecontemporaneously. Granulite-facies mineral assemblages are widely developed in the Upernavik suite and are interpreted to have formed synchronously with the late Archaean gneisses. Significantly, mineral assemblages in the Upernavik supracrustal rocks from the area where early Archaean relations are preserved, in the vicinity of Saglek, indicate amphibolite-facies conditions. This suggests that
208
K. D. C O L L E R S O N , A . K E R R & W . C O M P S T O N
and albitic plagioclase are interpreted as zones of melting which rapidly crystallized soon after formation. The Kiyuktok Gneisses Development of Kiyuktok gneisses with an The Kiyuktok gneisses are a heterogeneous intrusive component takes place either by lit-pargroup of rocks which can be described in terms of lit injection or by net-veining and agmatite two components: (1) a relict component which formation. The border zones of the intrusive generally displays gneissic layering, and (2) a new pegmatites are commonly sites of in situ transcomponent of more massive appearance. The formation, which appears to be related to the relict component is macroscopically identical to emplacement of the pegmatites (Fig. 2e). The the Uivak I and, more rarely, to the Uivak II pegmatites are coarse-grained microcline-rich gneisses. It is preserved either as isolated patches rocks displaying an equigranular granoblastic and lenses or merely as a "ghost layering" (Fig. mosaic of largely strain-free grains. Plagioclase is 2d): present in two forms: (1) large strained grains The new or reworked component is massive which resemble those in the Uivak gneisses, and and structureless and contains distinctive clots of (2) smaller strain-free grains of more albitic comferromagnesian minerals (principally biotite) position (An _ ). Orthopyroxene and garnet are ranging up to 2 cm in diameter. Contacts between the principal ferromagnesian phases and are varithe new and the relict components are commonly ably retrogressed to aggregates of amphibole and diffuse and gradational, but are easily recognized biotite. due to the striking contrast between the two rockIn the system Qz-Ab-0r-H 0 at 7.0 kb (von types. Lithological and petrological features of Platen & Holler, 1966) they lie close to the terthe two components are summarized in Table II. nary minimum with Ab/An ratios of 2.0 to 4.0 Fragmented and partially resorbed Saglek (Fig. 3) and are similar to experimental melts of dykes are present in both components and display sialic rocks reported by Brown & Fyfe (1970) and intrusive relationships with ghost Uivak layering Busch et al. (1974). Ab/An ratios in this range (Fig. 2b and 2 f ) . These features strongly suggest are typical of both the Uivak gneisses and the that the Kiyuktok gneisses were at least partially metasedimentary rocks of the Upernavik suite. derived from the pre-existing Uivak suite by some The nebulitic rocks which dominate the Kiyukform of reworking. tok gneisses elsewhere are similar in most respects Most representatives of the Kiyuktok suite are to the above, but contain smaller amounts of dominated by the new component (80-90% by microcline and also biotite in addition to garnet volume) and preserve only small amounts of and orthopyroxene. In terms of their georelict (Uivak) material. These rocks are largely chemistry, the nebulitic rocks are intermediate in structureless (Fig. 2c and 2d) and are termed composition between the Uivak gneisses and the nebulites, in the sense of Mehnert (1968). The intrusive pegmatites (Fig. 3). progressive development of the Kiyuktok gneisses TABLE II can thus only be studied in areas where reworking Characteristics of relict and new components in the was incomplete and gradational transitions beKiyuktok gneisses tween Uivak gneisses and the younger nebulites are preserved. The relationship between the two groups of gneisses is best understood in the area around Kiyuktok Cove, which was investigated in detail by Kerr (1980), and also along the southern shore of Shuldham Island (Fig. 1). In these areas it is possible to discriminate between Kiyuktok gneisses which developed entirely in place and those whose development also involved the introduction of an externally derived intrusive component. In the first case, the new component occurs as irregular patches with diffuse contacts, completely enclosed by the host Uivak gneiss (Fig. 2a). The new component is microstructurally heterogeneous and retains some features typical grains. of the Uivak gneisses, such as strained plagioclase and amoeboid quartz grains. Prominent finegrained domains consisting of quartz, microcline the early Archaean terrain represents a higher crustal level.
5
10
2
Relict
(Uivak)
Component
New Component
v
F i n e - s c a l e l a y e r i n g d e f i n e d by m e l a n o c r a t i c and l e u c o c r a t i c layers.
No c o m p o s i t i o n a l
S t r o n g f o l i a t i o n p a r a l l e l to compositional layering: d e f i n e d by a l i g n m e n t o f b i o t i t e a n d / o r hornblende and e l o n g a t i o n o f q u a r t z and f e l d s p a r .
G e n e r a l l y massive and nonf o l i a t e d except where strongly a f f e c t e d by l a t e r events.
Mafic minerals evenly d i s t r i b u t e d throughout cratic layers.
melano-
M a f i c m i n e r a l s l o c a l i z e d into c l o t s , which are distributed i r r e g u l a r l y through the rock.
W e l l - p r e s e r v e d S a g l e k dykes with sharp, l o c a l l y discordant contacts with t h e i r host rocks.
S a g l e k dykes are cut by and p a r t i a l l y resorbed by the new component.
layering.
Sharp ( t e c t o n i c ? ) c o n t a c t s with post-Uivak supracrustal rocks.
L o c a l l y i n t r u s i v e contacts with s u p r a c r u s t a l rocks.
Large, strained p l a g i o c l a s e g r a i n s o f o l i g o c l a s e to andesine composition.
P l a g i o c l a s e present both as l a r g e , strained grains (An2o-An 3 o) and as smaller, s t r a i n - f r e e a l b i t i c (Ans-Anio)
Microstructures inequigranular.
typically
Microstructures typically e q u i g r a n u l a r and granoblastic.
than 20% modal
Up t o 60-70fc modal microcline.
Normally l e s s microcline.
GNEISSES IN NORTHERN LABRADOR
Fig. 2.
209
Typical field characteristics of the Kiyuktok gneisses. (a) Relict " U i v a k " layering with new component occurring as diffuse irregular patches, Kiyuktok Cove. (b) Saglek dyke with cuspate-discordant contact against extensively reworked gneiss, Ramah Bay. (c) Nebulitic Kiyuktok gneisses showing diffuse pegmatitic veinlets and distinctive ferromagnesian clots, Kiyuktok Cove. (d) Faintly visible " g h o s t " layering in Kiyuktok gneiss, Nachvak Fiord. •(e) Layered relict component in Kiyuktok gneiss cut by a vein of pegmatite which localizes the in situ obliteration of the early structures, Kiyuktok Cove. (f) Amphibolite boudin in thoroughly reworked gneiss, Shuldham Island.
K. D. COLLERSON, A. KERR & W. COMPSTON
210
chaean intrusive rocks as broadly synchronous. In terms of field relationships, petrology and geochemistry, these intrusive rocks are comparable to the temporally equivalent Nuk gneiss and the Ilivertalik granite in West Greenland (McGregor, 1973, 1979; Kalsbeek & Myers, 1973; Compton, 1978; Bridgwater etal., 1978). A younger limit for late Archaean igneous activity is provided by the Igukshuak suite, a group of post-tectonic granites, pegmatites and aplites which cut all the above rock types. They are exposed as pervasive stock works, sheets and narrow dykes, and contain zircons which define a discordia intercept age of ca 2520 Ma (Baadsgaard et al., 1979). The Igukshuak suite is correlated with the Qorqut granite in the Godthaab Fig. 3. Ternary normative Q-Ab-Or diagram showing district of West Greenland (McGregor, 1973). the compositional variability of the Kiyuktok gneisses. Fields of Uivak gneisses from Coller- Isotopic Results son& Bridgwater (1979). Experimental Methods. Rb-Sr and preliminary Pb-Pb isotopic analyses of representative samples of the different groups of late ArLate Archaean Orthogneisses of Intrusive chaean gneisses were undertaken at the AustraOrigin lian National University (A;N.U.). Analyses of Gneisses derived from intrusive rocks of late seven samples of Ikarut gneiss were subseArchaean age are widespread throughout the quently carried out at Memorial University gneiss complex and are volumetrically important (M.U.N.): Rb-Sr analyses at A.N.U. were southwest of Hebron and along the southern determined using a modified version of the isoshore of Nachvak Fiord (Fig. 1). They are less im- tope dilution method described by Compston et portant in the Saglek area but are represented by al. (1965). Full details of methods are described thin, fine-grained foliated sheets and dykes of in Page et al. (1976). Samples were loaded as tonalitic and granodioritic composition. chlorides with H G on triple rhenium filament Late Archaean gneisses along the inner reaches beads. Rb was analysed on the MSX (12.25 cm of Hebron Fiord, termed the Ikarut gneisses, radius; 90° sector) mass spectrometer and Sr form units up to 1 km in thickness which are sub- analyses were made either on a Nuclide Analyconcordant to the pre-Ikarut layering in the sur- sis Associates (30.5 cm radius; 60° sector) mass rounding Uivak gneisses. They are mostly augen spectrometer (NAA) or a single focusing (23 cm gneisses (Fig. 4a) regarded as derivatives of .por- radius; 60° sector) mass spectrometer (MSZ). phyrinic tonalites and granodiorites, with sub- At M.U.N., Rb/Sr ratios were obtained by preordinate mafic and felsic schlieren. cise X-ray fluorescence spectrometry and Sr/ In the Nachvak area, broadly equivalent rocks Sr ratios, were measured as unspiked loads in known as the Kammarsuit gneisses were de- a VG Micromass 30B mass spectrometer using scribed initially by Morgan (1975) but were not single tantalum filament beads. named. The Kammarsuit suite is a group of Sr analyses were corrected for variable mass coarse-grained porphyroclastic gneisses (Fig. 4b) discrimination by normalizing to an Sr/ Sr of granodioritic to granitic composition and is ratio of 8.3752. Analysis of the Eimer and lithologically equivalent to the Ikarut gneisses. Amend Sr carbonate standard on the NAA mass Thin sheets and dykes of foliated tonalite and spectrometer gives a normalized Sr/ Sr ratio granodiorite (Fig. 4c) which are well exposed on of 0.70813 ± 0.00004 (l a; Page etal., 1976). The Lister Island and Maidmonts Island (Fig. 1), are normalized 7Sr/ 6Sr ratio for the NBS 987 common throughout the gneiss complex, particu- strontium carbonate on the same mass spectrolarly in the Saglek area, which is regarded as a meter is 0.71035 ± 0.00006 (la; Page et al., higher crustal level. 1976). Spiked and unspiked analyses of the same All of the above rock-types display clear intru- standard gave mean normalized values of 0.71031 sive relationships with the pre-3000Ma old gneiss ± 0.00003 and 0.71029 ± 0.00006 respectively on complex, but their contact relations with the the MSZ mass spectrometer (Compston et al., Kiyuktok gneisses are rather equivocal. We At M.U.N, the same standard gave a value regard the Kiyuktok gneisses and the late Ar- of1977). 0.71024 ± 0.00003 (1 a). Replicate analyses of Q
2
87
86
88
87
8
8
86
86
GNEISSES IN NORTHERN LABRADOR
211
Fig- 4. Typical field appearance of the late Archaean orthogneisses. (a) Ikarut gneiss with characteristic augen texture and xenolithic fragment of more felsic composition, Hebron Fiord. (b) Kammarsuit gneiss with large micrpcline porphyroclasts and vein of meta-aplite, Nachvak Fiord. (c) Sheet of fine-grained tonalitic gneiss cutting outcrop of Uivak II gneiss with large xenolith of Nulliak amphibolite, Maidmonts Island.
K. D. COLLERSON, A. KERR & W. COMPSTON
212
V -208N>L Free-line Method 2753 * jo Ma yS* Sr 0 7086 /•75-208L - Reference Isochron / / 3700 Ma / - 3700 ' Ma-V/f / / / 074 2753 May // /S/ 75- 208 M 073 I (•^Free-line 75-238V^| ^ ^ MethodMa" 2995!t| K Earth Composition. . P71./ / Sr XT 75-208 F1 . i 1 r 1 1 107055 11 1
1
1
1
5
0
35051113 Ma
CL
75-208N
-Q
75-208L 75-208K
0 7 5
-
/
\
'
0 7 2
206/20414 Pb
4 9
0
4 0 0 0-2 0 4 0-6 0-8 10
Fig. 6. Pb-Pb plot of analytical data for the Kiyuktok gneisses. The specimen numbers are shown for comparison with the Rb-Sr data in Figure 5.
the standard potassium feldspar SRM 607 using the 85Rb-84Sr tracer (746895) by a number of analysts at A.N.U., including K.D.C., yielded a mean Rb/ 6Sr ratio of 24.132 ± 0.043 (standard error of the mean) and a mean normalized Sr/ Sr ratio of 1.19910 ± 0.00081 (standard error of the mean). This is equivalent to an age of 1413 Ma, assuming an initial Sr/ Sr ratio of 0.710. The analysis reported by the National Bureau of Standards compares well with this
determination, being equivalent to an age of 1412Ma. The Rb-Sr data were statistically analysed using the regression methods of York (1969) and Mclntyre et al. (1966), and also by the new regression technique of Cameron et al. (in press). Experimental errors assigned to the data are: Rb/ Sr 0.25%, Sr/ <>Sr 5 x 10 ~ (A.N.U.); and Rb/ Sr 0.5%, Sr/ 6Sr 5 X 10~ (M.U.N.). The decay constant for Rb used in this study is 1.42 x 10 - a - (Steiger & Jager, 1977). Pb-Pb whole-rock isotopic analyses were undertaken at A.N.U. using a modified version of the double spike method of Compston & Oversby (1969), and Oversby (1975). The decay constants used for U and 235jj were 1.55125 x 10 - a - and 9.8485 x 10 - a - respectively (Steiger & Jager, 1977). Error estimates for the data are 0.1% (la). Isotope Geochemistry of the Kiyuktok Gneisses, Sr isotopic results for the Kiyuktok gneisses are presented in Table III and plotted in Figure 5. Regression of the total population of data (n = 25) by the York (1969) method gives an age of 2775 ±168 Ma and an initial Sr/ Sr ratio (Sr ) of 0.7079 ± 0.0014. However it is obvious that the data exhibit considerable scatter, far greater than expected from the assigned experimental errors alone. This is reflected statistically by a mean square of weighted deviates (MSWD) value of 1278. Mclntyre et al. (1966) proposed statistical models to assess the geochemieal reasons for such scatter. The present data favour a Model III solution, in which the scatter is considered to be the result of variation in Sr , and largely independent of Rb/ ^Sr. This defines an age of 2715 ± 37 Ma and Sr of 0.7089 ± 0.0014. Independent studies (Compston & Collerson,
o
10
20
30
Rb/ Sr Fig. 5. Rb-Sr isochron plot for samples of Kiyuktok gneiss. 8T
87
87
86
8
86
87
TABLE
86
III
Rb-Sr isotopic data for the Kiyuktok Gneisses, Labrador Sr
(total) (ppm)
Sample No.
Rb (ppm)
KC-78-208A/I
5.59
177.21 178.49
0.120
0.71599 ± 2
KC-75-237
30.23
528.31
0.165
0,71293 ± 2
KC-75-238A
30.08
471.73
0.184
0.71419 ± 5
KC-78-239A
32.87
495.65
0.192
0.71374 ± 3
KC-78-238B
42.11
570.20
01213
0.71369 ± 4
594.65
0.216
0.71352 ± 2
707.41
0.287
KC-78-208A/2
1M
KC-75-238B KC-78-238A
70.14
87Rb/86Sr
8 7 S r / 8 e 'Sr
0.091
0.71601 + 2
KC-78-239B
49.17
466.01
0.305
0.71842 ± 3 0.72365 + 2
KC-76-348
137.82
1257.18
0.317
0.71878 + 2
66.60
358.97
0.537
0.72812 + 3
KC-75-2081
59-75
241.35
0.717
0.73855 + 1
KC-75-2080
110.82
328.34
0.979
0.75203
KC-75-208K
± 2
KC-78-208E
108.09
300.70
,1.043
0.75328 ± 2
KC-7^~53A/B/2
120.52
296.06
1.181
0.75664 ± 3
KC-7i»-53A/B
121.76
294.74
1.198
0.75702 ± 3
KC-75-208M
152.67
206.31
2.155
KC-78-208F/2
170.93
196.19
2.540
0.80702 ± 4
KC-78-208F/3
169.24
193.98
2.544
0.80723 + 8
KC-78-208F
173.15
194.62
2.594
0.80828 ± 6
4.202
0.87644 ± 4
0.79659 ± 4
KC-75-208L
180.35
125.97
KC-78-634A
222.65
125.67
5.219
0.91434 ± 4
KC-78-63AB
237-92
118.84
5.914
KC-75-20&N
190.35
89.34
6.301
0.94212 ± 7 0 . 9 5 4 5 1 ± 12
KC-75-208N/2
192.05
89.71
6.331 .
0.95624 ± 3
87
86
87
87
8
86
5
87
8
5
87
1 1
1
238
1 0
1
1 0
1
87
86
0
0
8
87
0
GNEISSES IN NORTHERN LABRADOR
1979; Collerson & Compston, unpublished data; Collerson et a I., unpublished data) have shown thft quartzo-feldspathic gneisses in the SaglekHebron area have responded to the effects of localized isotopic disturbance at ca 2500 and 1800 Ma. These broadly correspond to the age of the Igukshuak granite, and to the Hudsonian Orogeny. Cameron et al. (in press) have developed a new regression technique specifically designed for the analysis of disturbed isotopic systems. Application of the method requires a knowledge of the error structure in the primary data caused by the disturbance, obtained in the present study from the analysis of secondary isochrons in Uivak I gneisses. Using the Free-line model of this method, the most completely transformed samples (n = 15) correspond to an age of 2753 and have an initial Sr isotopic composition of 0.7086 (Fig. 5). For reference purposes the present-day Bulk Earth composition and a 3700 Ma (Sr = 0.7000) isochron are also shown in this figure. On an expanded scale in the same diagram are analyses of Kiyuktok gneisses with Rb/ Sr ratios less than 1.1. Samples in this compositional range have readily recognizable relict Uivak features. When these samples are included with those previously described the total population yields an age of -269 o of 0.7055. The lower Sr reflects the effect on the regression of samples which are less radiogenic than the typical nebulitic Kiyuktok gneisses, and which, as a result, preserve the Sr-isotopic systematics which evolved in relict Uivak material. Q
87
2 9 9 5
Ma
a n d
h a s
a n
86
S r
0
TABLE I V
Pb-Pb whole-rock analytical data for the Kiyuktok gneisses 207 20* 206p 20«, Samp 1e No . 13.56 KC-75-238A \2.kk 13.90 KC~75"208K 13.37 03 KC-75-208L 13-80 KC-75-208M H. 77 15.97 KC-75-208N 1^.77 16.65 Analyst: J.J. Foster b/
pb
pb/
pb
Preliminary 207pb/204pb - 206pb/ 204pb isotopic data for samples of Kiyuktok gneiss (Table IV) range from distinctly non-radiogenic to moderately enriched compositions. Regression by the least squares method of York (1969) yields an age of 3505 ± 112Ma (Fig. 6), Whole-rock Pb isotopic data from other Archaean areas are presented in Figure 7 for comparison with the results for the Kiyuktok gneisses. In terms of their isotope character the Kiyuktok gneisses are remarkably similar to results from the AmTtsoq gneisses
213
\
Kiyuktok gneisses
Arrrttsoq gneisses
cut>/cvq pk Fig. 7. Pb-Pb whole rock isochrons for different Archaean gneisses. The primary growth curve was calculated assuming a 238y/204pb j f 7.7. Comparative data were taken from the following sources: AmTtsoq gneisses (Black et al1971), Isua iron formation (Moorbath et al., 1975tf), Lewisian gneisses (Chapman & Moorbath, 1977). r a t
0
0
(Black et al., 1971). They are however significantly different to results from the Isua iron formation (which are slightly older and are interpreted to reflect derivation from a source region with a higher 238jj/ 204pb value) and also to the ca 2700-2900 Ma old Lewisian tonalitic gneiss complex from Scotland, which is regarded as a juvenile addition to the continental crust (Chapman & Moorbath, 1977; Hamilton etal., 1979). It is therefore apparent that the Kiyuktok gneisses were not derived from short-lived crustal precursors, but must instead be viewed as isotopically reworked crust which was at least 3500 Ma old. Preservation of the 3500Ma old isochron is interpreted to be the result of the severity of a Ufractionation event at this time, which virtually removed U from the Uivak gneisses. Isotope Geochemistry of the Late Archaean Orthogneisses of Intrusive Origin. Results for the late Archaean intrusive rocks are presented in Table V and plotted in Figures 8, 9 and 10. Despite the pronounced lithological similarity of these gneisses, Sr isotopic results strongly indicate that they comprise at least three discrete intrusive suites. A summary of results obtained by the regression methods of York (1969) and Mclntyre et al. (1966) are given in Table VI. Like the Uivak and Kiyuktok gneisses, data for these gneisses also exhibit scatter as a result of localized isotopic homogenization at ca 2500 and 1800Ma. Regression analysis using the Free-line method of Cameron et al. (in press) yields an age of 2813 ^ 190 Ma and an Sr of 0.7034 for the tonalitic and granodioritic sheets at Saglek. These results contrast with those from the Ikarut and Q
K. D. COLLERSON, A. KERR & W. COMPSTON
214 TABLE
V Bulk Earth Method 2939!™ Ma Sr. 0 7011
Rb-Sr isotopic data for Late Archaean intrusive rocks Sample No.
Rb (ppm)
Sr (total) (ppm)
87
86
Rb/ Sr
87
Sr/
e 6
Sr
F r e e - l i n e Method
Tonali t ic and Granodioritic Sheets - Saglek Area: KC-7^-72
67.78
483.18
0.406
0.72073 ± 2 0.72623 ± 3
KC-74-42D
61.35
357.25
0.497
KC-77-548B
89.77
463.99
0.560
0.72582 ± 2
KC-76-342C
57.21
281.33
0.588
0.72679 ± 6
KC-76-342C/2
58.59
281.59
0.602
0.72702 ± 4
KC-77-548A
110.30
397.95
0.803
0.73512 ± 7
KC-74-91A
152.97
375.72
1.180
KC-7^-91A/2
154.77
374.86
1.197
0.74622 + 3 0.74682 + 8
KC-74-44B
229-70
553.42
1.204
0.75218 + 1 0.77792 + 4
KC-74-153A/2
126.64
238.63
1.543
KC-74-153A/1
127.86
235.23
1.581
0.77881 ± 2
KC-76-343A
150.07
227.70
1.916
0.77707 ± 4
184.72
241.25
2.230
0.79511 ± 4
KC-7^-150B
185.82
240.67
2.249
0.79626 + 7
KC-76-431E
166.73
193.25
2.518
0.81905 + 6
KC-74-150B/2
0.80456 + 3
KC-77-5^1A
132.60
153.06
2.525
KC-74-45D
228.31
246.32
2.704
0.81193 ± 3
KC-74-45E
232.59
250.42
2.708
KC-77-548C/2
210.29
206.32
2.976
0.80806 + 7 0.82145 + I
KC-77-5A8C/1
208.25
202.79
2.998
0.82213 ± 3
KC-76-420
304.66
219.86
4.063
0.86481 ± 4
267.41
1.091
0.74904 ± 2
Kammarsuit Gneisses - Nachvak Area: KC-76-410A/2 KC-76-Al0A/1
100.60 100.75
262.95
1.111
0.74948 ± 6
KC-76-410B/2
107.12
258.64
1.201
0.75219 + 3
KC-76-410B/1
105.54
254.32
1.204
0.75225 ± 3
KC-76-410D
102.44
249.30
1.192
0.75115 ± 2
KC-76-410C/1
283.37
65.94
13.023
1.21362 ± 5
0.184
0.71333 + .00011
Ikarut Gneisses - Hebron Fiord: KC-78-631A
0.064
CO £ 0-80
KC-78-631C (1)
0.173
0.500
0.72575 ± .00003
KC-78-631C (2)
0.192
0.556
0.72748 ± .00003
KC-78-631B
0.227
0.658
KC-78-631D
0.333
0.966
0.74405 ± .00005
KC-78-631F
0.424
1.232
0.75641 ± .00004
KC-78-631E
0.714
2.083
0.78880 ± .00009
0.73379 ± .00010
Kammarsuit gneisses which yield ages of 2774 1 $ Ma (Sr0 = 0.7059) and 2689 i M a (Sr0 = 0.7057). The differences in initial 8 7 Sr/ 8 6 Sr ratios between the three sets of results are also apparent if the regression techniques of York (1969) and Mclntyre et al. (1966) are employed (Table VI), although the ages are basically equivalent. An estimate of the maximum protolith ages for these younger gneisses can be obtained by applying the Bulk Earth regression technique of Cameron et al. (in press). This yields the following estimates of maximum protolith age and Sr 0 ? s for the different groups of gneisses: Saglek area 2941 ± J ^ M a , Sr 0 = 0.7011; Ikarut gneisses 3301 ^349 Ma, Sr 0 = 0.7006; Kammarsuit gneisses 28771 f g M a , S r 0 = 0.701.
; |- 2813* jgo Ma
Sr. 0 7034
l
Bulk E a r t h Composition
10
0-5
15
20
8 7 R b / 8 6 S r
Fig. 8.
30
40
Rb-Sr isochron plot for samples of late Archaean tonalitic and granodioritic gneiss from the Saglek area.
The rationale behind the development of the Bulk Earth model to provide an estimate of precursor age is as follows. It has been shown by DePaolo & Wasserburg (1976a; 19766) that many oceanic and continental basic igneous rocks are derived from a mantle reservoir with a chondritic Sm/Nd ratio (viz. 0.31). The same authors, as well as O'Nions et al. (1977) and Hawkesworth et al. (1977) have also recognized a strong inverse correlation between ^ N d / i**Nd and 87 Sr/ «%; They have interpreted this to imply that Sm/Nd and Rb/Sr have fractionated coherently during mantle evolution, enabling the present-day Bulk Earth 87§ r /865 r to be estimated (viz. 0.7047). This ratio corresponds to the Sr isotopic composition of a mantle reservoir which evolved from ca 4550 Ma to the present day. The present-day Bulk Earth 8 7 Rb/ 86Sr ratio, which corresponds to an Rb/Sr ratio of 0.03, can thus be calculated (viz. 0.0856). Assuming that the mantle source region of basaltic rocks has a composition which approximates that of the Bulk Earth, it can be readily shown that all primary Rb-Sr isochrons derived from such regions will pass through a fixed point
F r e e - l i n e method
2774
IS? Mov
/
Fig. 9.
Rb-Sr isochron plot for the Ikarut gneisses.
GNEISSES IN N O R T H E R N L A B R A D O R
215
TABLE VI
Rb-Sr iso topic results for Late Archaean intrusive rocks, Northern Labrador Age Tonal it ic and G r a n o d i o r i t i e Sheets Saglek Area Ikarut Gneisses - Hebron A r e a
Kammarsuit Gneisses - N a c h v a k A r e a
(Ma)
Sr/86Srn
Model
2855 ± 191
0.7031 ± 0.0028
2815.± 190
0.7037 ± 0.0038
2782 ± 151
0.7059 ± 0.0012
2776 ± 156
0 . 7 0 5 9 ± 0.0016
2689 ± 12k
0.7057 ± 0.0023
2686 ± 136
0.7058 ± 0.0024
that corresponds to this composition. Although studies of ancient and modern basaltic rocks detect that the mantle is geochemically heterogeneous (e.g. Hanson, 1977; Hofmann & Hart, 1978; Sun et al., 1979; Wood, 1979), Sm/Nd isotopic results (DePaolo & Wasserburg, \916a and Hamilton et al., 1978) support an interpretation that i43Nd/ i44Nd and Sr/ 6Sr evolution in the Archaean mantle approximated that of the Bulk Earth. If it is assumed that the plutonic precursors of Archaean tonalitic gneisses formed by partial melting of short-lived mantle-derived mafic protoliths whose mean initial Sr/ ^Sr ratios equalled that of the contemporary Bulk Earth, then a primary isochron defined by these gneisses would yield an estimate of the age and initial Sr/ Sr ratio of these protoliths if forced to pass through the Bulk Earth composition (0.0856; 0.7047). If, on the other hand, the gneisses were derived by the metamorphism of anatectic sialic melts, or if there has been opensystem redistribution of Rb or radiogenic Sr on a large scale, then the assumptions behind the Bulk Earth method are no longer valid. The age and Sr of the late Archaean orthogneisses in the Saglek area are broadly comparable to those presented by Moorbath & Pankhurst (1976) from the temporally equivalent and
Skk
York
(1969)
M c l n t y r e et al. 37
York
36
York
(1966) M o d e l IV
(1969)
M c l n t y r e e t al.
(1966) Model IV
(1969)
M c l n t y r e et al.
(1966) M o d e l 11
lithologically similar Nuk gneisses from the Godthaab area of West Greenland. However, gneisses with high initial Sr/ ^Sr ratios similar to those displayed by the Ikarut and Kammarsuit gneisses are as yet unreported from West Greenland. The calculated values for Sr for the precursors of the Saglek orthogneisses and for the Kammarsuit gneisses, each 0.7011, are closely comparable to the Sr/ Sr value obtained by Hart & Brooks (1977), which is believed to reflect the Sr/ 6Sr composition of the mantle source region at that time. However, the value obtained for the Ikarut gneisses (0.7006) is distinctly different, suggesting that earlier Archaean material has been involved in the pedogenesis of this suite. DISCUSSION Kiyuktok Gneisses Preservation in the Kiyuktok gneisses of a o r 3500 Ma Pb-Pb age strongly suggests that the source area of these rocks experienced an early Archaean high-grade metamorphic event which severely fractionated U from Pb as suggested by Black et al. (1973) and Moorbath (1976). This is fully consistent with compelling field evidence (discussed earlier) which indicates that the Kiyuktok gneisses were derived by partial melting and remobilisation of the Uivak suite. In a plot of Rb against Sr (Fig. 11) there is a considerable degree of overlap between the fields for the Kiyuktok gneisses and their postulated precursors. Since the ratio of Uivak I to Uivak II gneisses in the early Archaean enclave between Saglek and Hebron is approximately 9:1, it is reasonable to view the Uivak I gneisses as the predominant protolith. However, in terms of Rb-Sr chemistry, the fractionation trend of the Kiyuktok gneisses (Fig. 11) is similar to that displayed by the Uivak II gneisses. As field relationships indicate that the bulk of the Kiyuktok gneisses were derived from the Uivak I suite, we conclude that the similarity in Rb-Sr chemistry reflects the reworking process, rather than the isochemical of the Uivak II suite. We thereFig. 10. Rb-Sr isochron plot for the Kammarsuit metamorphism fore tentatively ascribe the trend either to fracgneisses. 87
86
0
8
0
8
87
87
87
8
87
86
87
8
K. D. COLLERSON, A. KERR & W. COMPSTON
216
are contained in Collerson et al. (unpublished data), Cameron et al. (in press) and Collerson & Compston (unpublished data) and in an earlier publication by Hurst et al. (1975). The most recent estimates of age and Sr (Collerson & Compston, unpublished data) are 3659 +||]Ma and 0.70019 by the Free-line method and 3664 + ^ M a and 0.70014 by the Bulk Earth method of Cameron et al. (in press). The mean Rb/Sr ratio for the Uivak I gneisses obtained from the centroid of the Free-line isochron and also estimated from Figure 11 is 0.22. This corresponds to an 87Rb/ 86Sr ratio of 0.705. Using the information presented above, it is possible to model the temporal growth of radiogenic Sr in the postulated protoliths of the Kiyuktok gneisses. In Figure 12 it is clear that by 2750Ma, Uivak gneisses with an average Rb/Sr ratio of Fig. 11. Log-log plot of Rb against Sr showing the 0.22 would exhibit an 87s /86s ratio of ca fields for the Uivak gneisses and data for the 0.7086, the value obtained by the Free-line regresisotopically analysed Kiyuktok gneisses. sion of the most nebulitic, and presumably most tional melting of the Uivak I gneisses, or to mix- reworked group of gneisses. Kiyuktok gneisses ing in various proportions between a minimum with low R b / Sr ratios (less than 1.1) generally display a greater amount of relict (Uivak) melt and Uivak I gneiss. Further support for this hypothesis is provided material than gneisses with higher ratios. These by modelling the Sr evolution of the Uivak I high R b / 86Sr ratios may be the result of partial gneisses. Rb-Sr isotopic data for these gneisses reworking of the Uivak II suite (e.g specimens 0
87
500
1000
2000
Sr (p.p.m.)
r
87
r
86
87
0-730
0725
Q
O oc in <0 00 fjfi
0720
0 715
T
/
1 1 1Rb/Sr A Mantle evolution trends Al subcontinental A2 oceanic islands 0 03 A3 mid-oceanic ridge B Ikarut gneiss 0 12 G Uivak gneiss 0 22 D fSaglek tonalitic and 0 27 E : (granodioritic sheets 0-42 F ., 0 41 Q Kammarsuit gneiss Q ^Q
$/
g / e / / / /
/ /> / //
/>
0 710
oo [ dominantly nebullticKIYUKT0K J o G N E I S S \ with abundant 1 relict component • "E 0 7 0 5 _
-Mixing
0 700 4
5
4
0
3-5
3.0
25
20
15
10
0-5
0
Age Ga Fig, 12. Sr evolution diagram showing the average Sr growth lines for the Uivak I gneisses and the late Archaean S l S ' ^ i ^ meaMired initial 8 7 / 8 6 ratios for different populations of Kiyuktok gneisses® also plotted and lie along the S r / S r growth line for the Uivak gneisses Sr
87
86
Sr
GNEISSES IN NORTHERN LABRADOR
KC-78-634A & B) or, alternatively, they may reflect addition of Rb from an external source, either in the form of intrusive pegmatites or via a fluid phase (c./. Collerson & Fryer, 1978). Kiyuktok gneisses with low 8 7 Rb/ 86Sr ratios are therefore regarded as partially reworked Uivak gneisses. In these gneisses Sr isotopic systems were only partially re-set and no significant addition or loss of 87Sr or 87 Rb is believed to have occurred. Inclusion of these dominantly relict gneisses in the regression of data for the total population results in a bias towards a lower initial 87 Sr/ 86Sr composition. Nevertheless, this Sr 0 , when plotted against age in Figure 12, falls virtually on the Uivak I gneiss evolution line. When 87 Sr/ 86 Sr is plotted against \i mole 8^Sr (Fig. 13), the majority of the Kiyuktok gneisses fall within the fields of the Uivak I and Uivak II gneisses. However, the Kiyuktok gneisses with high 87 Rb/ 86Sr ratios show considerable enrichment in 87Sr and lie along the trend defined by the Uivak II gneisses. Although this could be interpreted as reflecting the isochemical metamorphism of the Uivak II gneisses, it is considered unMkely for the reasons discussed above. A more plausible explanation is that it is the result of sampling of volatile-bearing partial melts which are enriched in 87 Rb or 87Sr and which are mixed in various proportions, with members of the Uivak I suite.
070 0
0-2
04
0-6
0-8
10
1-2
1-4 1-6
p mole 8 6 S r
Fig. 13. Mixing diagram defined by plotting 8 7 Sr/ 8 6 Sr vs moles 8 6 Sr, showing data for the Uivak suite and the Kiyuktok gneisses.
217
Late Archaean Orthogneisses of Intrusive Origin Consideration of Rb-Sr isotopic results for the late Archaean orthogneisses, and for comparison, the Kiyuktok gneisses and Uivak I gneisses presented above, shows that the younger orthogneisses are broadly comparable in age (ca 2700 to 2800 Ma) and have significantly different initial ratios (0.7034 to 0.7059), which indicate differences in the isotopic composition of their source regions or interaction with crustal strontium. The tonalitic and granodioritic gneiss sheets from the Saglek area show low initial ratios which are comparable with those reported from the Nuk gneisses of the Godthaab region (Moorbath & Pankhurst, 1976). Accordingly, they are viewed as probable derivatives of short-lived precursors of mafic composition ( e.g. the mafic component of the Upernavik supracrustal suite?), an explanation which is consistent with the petrology and geochemistry of these rocks and also with current petrogenetic models for the genesis of tonalitic magmas ( e.g. Arth & Barker, 1976; Wyllie, 1977; Glikson, 1979). The higher initial ratios displayed by the Ikarut and Kammarsuit gneisses (0.7057-0.7059) suggest input from a more radiogenic source region, such as pre-existing sial. Since these rocks are dominantly granodioritic in composition, their derivation by anatexis of pre-existing sial is feasible (Wyllie, 1977). Xenoliths within the gneisses may represent samples of the source. However, estimates of protolith age for the different groups of late Archaean gneisses indicate that the Ikarut gneisses have been derived from significantly older precursors than the other two gneissic groups. The age of the protoliths of the Saglek sheets and Kammarsuit gneisses is ca 2900Ma and the Sr 0 for this source region is ca 0.7011 which agrees well with estimates by Hart & Brooks (1977) of the Sr isotopic composition of the mantle at that time. To evaluate the applicability of this model to these gneisses, the Sr isotopic evolution of the three suites of orthogneisses is depicted in Figure 12. The Sr isotopic evolution of the gneisses from Saglek and the Kammarsuit gneisses is modelled firstly using Sr 0 values and average Rb/Sr ratio estimates derived from the Free-line regression. When plotted in terms of Sr 0 and age, and extrapolated back in time, the growth curve intersects the mantle evolution line at ca 3000 Ma. Then, using the postulated Sr 0 of the protoliths and average Rb/Sr ratios from the Bulk Earth regression and modelling forward development, it is clear that the measured Sr 0 values (by the Freeline method) would be reached in ca 200 Ma,
218
K. D. COLLERSON, A. KERR & W. COMPSTON
which would correspond to the age given by the chemistry to Archaean tonalitic gneiss complexes (Windley& Smith, 1976). isochrons. Results presented for the Ikarut gneisses and Implications for Crustal Development plotted in Figure 12 indicate that these gneisses This study has demonstrated beyond reasonmust have evolved independently. Firstly, the source must have had a different Rb/Sr ratio able doubt that the formation and stabilization of late Archaean crust in Northern Labrador inthan that represented by the average used in this study, as the extrapolated Sr evolution line back volved substantial regeneration of pre-existing to the mantle growth curve intersects it at caAXOO sialic material. In one group of gneisses (the Ma, which is considered to be too old. This there- Kiyuktok gneisses) the sialic precursors were at fore indicates that multistage processes were in- least 3500 Ma old, and are probably represented volved in the petrogenesis of the Ikarut gneisses by the early Archaean Uivak gneisses. The reworking of this early Archaean sial was such as: (1) anatexis of ca 3650 Ma old granulitefacies (low Rb/Sr) crust or (2) magmatic mixing accomplished by in situ partial melting which was initiated by solidus depression as a result of the or isotopic exchange between material derived from a ca 3200-3000 Ma old basaltic (low Rb/Sr) passage of volatiles (dominantly water) from lower-crustal (and possibly upper-mantle) levels. source (Upernavik mafic gneisses) and Kiyuktoktype nebulitic (high Rb/Sr) melt. It is also pos- This resulted in the development of a zone within sible that basic magma at ca 3000 Ma was in- the crust where P Q = P , and Hp-saturvolved in such mixing and melting processes (c/. ated melting was possible under the ambient geoPatchett, 1980). Similar petrogenetic schemes therm. Operation of this process is fully consisbased on mixing or isotopic contamination be- tent with estimates of pressure and temperature tween mantle and crustal components have been from metamorphic assemblages in the suprainvoked to explain the petrology and geo- crustal units (6.5-9kb and 750-850°C). These chemistry of certain Phanerozoic calc-alkaline results suggest a late Archaean geothermal gradiplutonic and volcanic associations (Eichelberger, ent in this part of the North Atlantic Craton of 1975; Armstrong et al., 1977; Francis et aL, 1977) approximately 30-35 °C/km, which is slightly which are broadly analogous in terms of geo- higher than the gradients estimated by O'Hara Ui
Total
Fig. 14. Schematic diagram showing possible crustal processes during the late Archaean based upon evidence From Northern Labrador.
GNEISSES IN NORTHERN LABRADOR
(1977) and Wells (1979) for Archaean gneisses in Scotland and Greenland (25-30 °C/km). The intrusive pegmatites which form part of the Kiyuktok suite are regarded as the transport mechanism for water and other volatiles liberated by melting and dehydration of the Upernavik metasediments and the Uivak gneisses under granulite-facies conditions. The model envisaged is shown schematically in Figure 14. The more varied characteristics of the late Archaean intrusive rocks can be explained in a number of ways: (1) Partial melting of a relatively short-lived basaltic source material (e.g. Upernavik mafic gneiss). (2) Partial melting of a ca 3000 Ma old tonalite or trondhjemite parent derived by the above mechanism. (3) Mixing of melts produced by mechanism (1) with melts derived by anatexis of early Archaean sial (e.g. the Kiyuktok gneisses). Model (1) is most applicable to the tonalitic sheets and dykes in the Saglek area, which are similar to the Nuk gneisses of the Godthaab district. However, the composition and isotopic characteristics of the Ikarut and Kammarsuit gneisses are explicable by either model (2) or model (3) as shown schematically in Figure 14. In view of the close temporal and spatial association between the Kiyuktok gneisses, we favour model (3), which invokes large-scale mixing and isotopic exchange. Such mixing could be accomplished by hybridization of magmas or by contamination of juvenile magma with crustally derived fluids, enriched in radiogenic Sr. Isotopic homogenization of Sr must then have occurred between the juvenile and contaminating component (cf. Briqueu & Lancelot, 1979). Although rocks corresponding to the Kiyuktok gneisses have not yet been reported elsewhere in the North Atlantic Craton, the late Archaean intrusive rocks described in this paper have clear lithological equivalents in the Nuk-Ilivertalik plutonic association in West Greenland. However, the initial Sr isotopic composition of these intrusive rocks from Greenland is generally lower and similar to the Sr isotopic composition of the
219
contemporary mantle. This is used as evidence by Moorbath (1975a, 1976, 1977) that these gneisses, and other Late Archaean and Early Proterozoic intrusive suites, are juvenile mantle-derived additions to the crust. Although we agree in general with these conclusions, we wish to point out that the Late Archaean evolution of the Labrador portion of the North Atlantic Craton also involved substantial reworking of Early Archaean crust to form the Kiyuktok gneisses. In terms of areal extent, reworked Early Archaean rocks in Northern Labrador are more important volumetrically than "juvenile" material. Indeed, even the rocks which resemble typical Nuk gneisses, in terms of lithology and geochemistry, appear to reflect derivation from a higher Rb/Sr sialic source, or alternatively, substantial Sr-isotopic contamination with crustal material. Accordingly, we suggest that sialic crust can be reworked on a large scale (cf. Moorbath, 19756, 1977) and reworking may well be a more important process in late Archaean crustal development than has previously been acknowledged, occurring synchronously with the emplacement of "juvenile" material into the crust. Its recognition, however, is strongly controlled by the presently exposed structural level and by the intensity of subsequent deformation and metamorphism. ACKNOWLEDGMENTS Financial support for this study was provided by an NSERC Operating Grant XA-8694) to Collerson between 1974-1980, and a research grant from the Department of Indian and Northern Affairs to Memorial University in 1977 and 1978. Kerr was supported during his graduate programme from these sources and also by a Postgraduate Fellowship from Memorial University of Newfoundland. Isotopic work was mostly undertaken while K.D.C. was a Visiting Research Fellow in the Research School of Earth Sciences at A.N.U. during 1978-79. Technical assistance from all members of the Isotope Geology Group, especially L. P. Black, J. J. Foster, R. W. Page and D. Millar is gratefully acknowledged.
REFERENCES ARMSTRONG, R . L . , TAUBENECK, W . H . , & HALES,
P. O., 1977: Rb-Sr and K-Ar geochronometry of Mesozoic granitic rocks and their Sr-isotopic composition, Oregon, Washington and Idaho. Bull,
geol. Soc. Am., 88, 3 9 7 - 4 1 1 . ARTH, J . G . , & BARKER, F . , 1 9 7 6 : Rare earth partitioning between hornblende and dacitic liquid and implications for the genesis of trondhjemitic-tonalitic magmas. Geology, 4, 5 3 4 - 5 3 6 .
BAADSGAARD, H . , 1976: Further U-Pb dates on zircons
from the early Precambrian rocks of the Godthaabsfjord area, West Greenland. Earth planet.
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ARCHAEAN SEDIMENTARY AND VOLCANIC ROCKS
ARCHAEAN EPICLASTIC SEDIMENTS DERIVED FROM MAFIC VOLCANICS, NORTH POLE, PILBARA BLOCK, WESTERN AUSTRALIA John S. R. Dunlop & Roger Buick
Department of Geology, University of Western Australia, Nedlands, Western Australia 6009 ABSTRACT Silicified arenitic sediments from an early Archaean chert-barite unit at North Pole, Pilbara Block, Western Australia contain clasts which are texturally identical to underlying mafic volcanics. Most grains were derived locally from pillow margins and flow tops, although some scoriaceous grains may have come from local, highly vesicular basalts, or from distant felsic or mafic subaerial eruptions, transported as floating clasts. The volcanic arenites were deposited in shallow sublittoral sand bodies in a tidally dominated, moderately high energy environment. The sand bodies formed shoals upon which intertidal and supratidal lutites and evaporites were deposited. The shoaling sequence was only formed when general basin subsidence and volcanic eruption was temporarily and locally inactive.
INTRODUCTION pillowed flows of mafic and minor ultramafic Well-preserved, early Archaean sedimentary metavolcanic rocks; felsic volcanics and terrirocks of low metamorphic grade are rare but im- genous sediments are very rare and only occur near the top of the sequence. The volcanic pile portant. Knowledge of the early evolution of life has many intercalated chert beds, mostly lenticuand the environmental conditions which con- lar and less than 10m thick, but the larger beds, trolled the rate and style of this initial biological such as the chert-barite unit, are up to 50 m thick radiation depends upon studies of such sedi- and persist laterally over 20 km. ments, Until recently, however, the only well- In the massive mafic metavolcanics enclosing documented examples of particularly ancient the chert-barite unit, actinolite metalow-grade sediments were from the 3.5-3.3Ga morphic mineral present, whileisinthethemain more perSwaziland Supergroup of southern Africa. meable pillowed mafic rocks, prehnite Lately * rocks from a similarly old chert-barite pumpellyite are common. This suggests that and the unit at North Pole in the Pilbara Block of massive rocks equilibrated at higher metaWestern Australia have been shown to contain morphic temperatures than the pillowed rocks, probable evidence of life in early Archaean i.e. about greenschist-facies conditions compared shallow-water sediments (Dunlop, 1976, 1978; with prehnite-pumpellyite-facies conditions. Dunlop et al., 1978; Lambert et al., 1978; Barley Brittle deformation of the shallowly-dipping et at., 1979; Walter et al., 1980), thus greatly rocks (less than 45 °) is predominant and the area broadening the range of information available to exhibits only low strain. students of these problems. This paper discusses ages obtained by different methods volcanic epiclastic sediments from North Pole Isotopic various Warrawoona Group rocks suggest which place important constraints on palaeo- from deposited between about 3.4 and environmental modelling of the sedimentary that theyagowere (Sangster & Brook, 1977; Pidgeon, basin in which the chert-barite unit was 3.5Ga 1978; Hamilton et al., 1981). A model lead age of deposited. about 3.42Ga has been obtained from galena North Pole (119° 28 'E: 21° 07/S) is about within hydrothermally remobilized barite at 160 km south of Port Hedland and 50 km west of Marble Bar (Fig. 1). The chert-barite unit is the North Pole (J. Richards, pers. comm., 1978). lowermost concordant sedimentary unit exposed METAVOLCANIC ROCKS in the North Pole Dome (Hickman, 1973), lying The chert-barite unit at North Pole is enclosed near the base of the Warrawoona Group, the predominantly metavolcanic lower part of the within metamorphosed basalts which probably represent a tholeiitic suite (c/. Barley, 1980); Archaean supracrustal succession of the Pilbara Block. In the North Pole Dome, the Warrawoona there is no indication of contemporaneous felsic volcanism in the area. The metavolcanics are Group is dominated by interbedded massive and Spec. Pubis geol. Soc. Aust., 7 (1981)
J. S. R. DUNLOP & R. BUICK
226
(
Mafic-Ultramafic
Metavolcanic Rocks
Lowermost Chert - Barite Unit 0 t h e r
North
S t r a t i f o r m
C h e r t
Pole Adamellite
Units
I
I
E l m
E l
Fig. 1. Geological sketch map of the eastern flank of the North Pole Dome.
composed of a complex of interlayered massive and pillowed flows and sub-volcanic sills, each several metres to tens of metres thick. In pillowed flows, structures range from small (less than 1 m diameter) subspherical bun-shaped pillows to large (greater than 2 m diameter) mattress-shaped and mega-pillows (terminology after Dimroth et al., 1978). Discrete pillows are rare; most small structures are buds from larger structures or parts of tubular forms. Where well developed, tubes often branch and entwine to form entrail-like masses and their surfaces have flow striations. Where three-dimensional exposure of lava tubes allows the flow direction to be determined, it was from west to east, but because such exposures are scarce, this direction may not be general. Thin layers of hyaloclastic breccia surround pillows and also overlie pillowed flows. Many pillows have a concentric, marginal vario l i c layer, about 5-10 cm from the pillow surface. The varioles are 2-10 mm in diameter and, in the middle of the layer, overlap to form 90% of the rock. Both massive and pillowed flows are commonly vesicular (or amygdaloidal), with vesicu-
larity increasing upward within each flow. In one pillow immediately underlying the chert-barite unit, vesicularity is about 40% with a mean vesicle diameter of 3 mm, indicating eruption into a subaqueous environment, probably with less than 100 m water depth (Jones, 1969). No hyalotuffs have been identified in the metavolcanics, suggesting that water depths were great enough to prevent phreatic explosion. In surface exposures, the metavolcanics have been heterogeneously altered (cf. Barley, 1980) and deeply weathered. Some relict igneous mineral phases are present in drill core, mostly clinopyroxene and ilmenite. However, even in the most highly altered material, the crystal forms of the original minerals are often preserved as pseudomorphs. Plagioclase has been albitized, olivine replaced by chlorite, quartz, carbonates, and less commonly epidote and pumpellyite, clinopyroxene within varioles by quartz and albite or, in non-variolitic rock, by quartz, albite, and chlorite, and ilmenite by leucoxene and various iron oxides. Much of the silicification and carbonation is post-metamorphic. Because of the pseudomorphic mimicry of primary mineral phases, many igneous and devitrification textures are preserved: 1. Spinifex textures (Fig. 3a) occur in quenched, more-magnesian, massive flows. Clinopyroxene was the main primary mineral phase. Crystals ranged from small (5 mm), acicular, skeletal forms at the upper and lower margins of flows to large (15 cm) dendritic types near the flow top. The groundmass was composed of radiating clusters of 1 mm clinopyroxene crystals. Cumulate olivine was present toward the base of some flows. 2. Subophitic textures (Fig. 3c) occur throughout less-magnesian, massive flows that have not been quenched. Some highly carbonated, non-variolitic pillows also appear to have subophitic textures developed throughout. Usually, however, only pillow cores are subophitic, some with scattered pseudomorphs after olivine phenocrysts. 3. Hyaloclastic textures are developed in the outermost of the three textural zones of pillow margins. Very fine grained devitrified glass (now chlorite-silica-magnetite) forms a 2-10 mm thick crust on the exterior of pillows, and is the source for the rectangular blocks of calcite-cemented hyaloclastic material interstitial to pillows. 4. Perlitic textures (Fig. 3i) form between the hyaloclastic and variolitic zones of pillow margins. Lenticular and subspheroidal bodies (1-3 mm in diameter) mainly composed of
EPICLASTIC SEDIMENTS FROM NORTH POLE, PILBARA
silica but often with epidote-pumpellyite cores and concentric brownish bands of irontitanium oxides, chlorite and albite are bounded by anastomosing perlitic fractures some of which contain prehnite. Farther away from the pillow margin, these structures grade into 1 mm thick monomineralic laminae of chlorite, magnetite or silica. The perlitic zone is usually 1-3 cm thick. 5. The variolitic layer is the innermost textural zone of pillow margins and begins about 5 cm from the outer edge of most pillows. It has gradational contacts with the outer perlitic zone and the subophitic pillow core with the number of varioles rapidly decreasing in both directions. The varioles, which are defined only by a colour difference resulting from a lesser abundance of chlorite, have overgrown an original olivine-phyric (Fig. 3e) and felted microlitic texture (Fig. 3g) causing no textural disturbance. Some relict clinopyroxene is present in the microlitic groundmass and prehnite or silica fills small vesicles.
THE CHERT-BARITE UNIT The chert-barite unit is restricted to the eastern flank of the North Pole Dome by an apparent structural discontinuity between the low-strain, brittly deformed eastern sector and the more folded and strained western side. The contact between the underlying metavolcanics and the chert-barite unit has chertified sediments draping over highly silicified pillows and hyaloclastic breccia; no erosional features are visible. The upper contact is also concordant and nonerosional, and therefore the changes in depositional style were probably not marked by hiatuses* Four main sedimentary types make up the chert-barite unit: 1. Silicified carbonates occurring mainly as futites and intraclast conglomerates (Dunlop, 1978), - 2. Baritized or silicified evaporative gypsum of both primary and diagenetic origin (Lambert etal. 1978; Barley etal., 1979), 3. Primary chemical sediments, now banded red, white, black and grey cherts, 4. Green arenitic sediments, which are interpreted here as silicified volcanic epiclastic sediments derived from pillowed basalts and scoriaceous rocks. The chert-barite unit was deposited in shallow water, indicated by the presence of evaporites, rapid lateral and vertical facies changes within the unit, vesicle size and abundance in the enclosing metavolcanics and the prevalence of littoral features, such as intraclast conglomerates, flaser, f
227
wavy and lenticular bedding and bimodal flame structures. Desiccation phenomena (breccias of curved mud flakes) suggest periodic exposure to the atmosphere. The discovery of stromatolites (Dunlop & Groves, 1978; Walter et al., 1980) and carbonaceous microspheroids which have been interpreted as microfossils (Dunlop et al., 1978) supports a shallow-water depositional environment, if the benthic microbiota was photosynthetic. EPICLASTIC VOLCANIC SEDIMENTS Silicified epiclastic sediments with grains interpreted to be of volcanic origin are one of the commonest rocks in the chert-barite unit. They have a characteristic green colour, attributable to the relatively high chlorite content of the clasts and will be referred to as Volcanic arenites" (as used by Folk, 1968) to distinguish them from other arenites from different provenances. The volcanic arenites are usually found in thick (1-10 m) beds near the top of the stratigraphic sequence. Significant beds (1-2 m thick) occur near the base of the sequence in some localities and thin (less than 0.5 m thick) beds occur throughout the sequence. The lateral continuity of the beds is uncertain because outcrops are restricted to hilltops and are commonly fault bounded. It is suspected, however, that the beds are lenticular, as are those of other rock types where they can be traced laterally. The volcanic arenites are underlain by almost all rock types present in the chert-barite unit and their basal contact is usually non-gradational. Where the arenite beds overlie lutite, the contact is locally erosional, with massive lutite grading into intraclast conglomerate or breccia composed of tabular or, rarely, blocky clasts with an arenitic matrix. The frequency of lutite clasts decreases rapidly upward, leaving a well-sorted arenite. Some contacts with underlying lutites are stylolitic. The usual rock overlying the volcanic arenites is massive or finely laminated grey silicified lutite, thought to have been originally calcilutite, locally containing 1-2 mm silica pseudomorphs after gypsum. This upper contact is commonly gradational over 0.2 to 0.5 m, from grain-supported arenite with no matrix, through matrix-supported arenite to pure lutite. In a few places, this transition is marked by coarsely interlayered bedding, from arenite through flaser bedding, wavy bedding, and lenticular bedding to lutite. Where other rocks overlie the arenite beds, the contact is always sharp. Graded bedding is almost universally developed in the volcanic arenites, often on a number of scales. The arenite beds have a finingupward tendency, over 2-5 m, often forming 4
228
J. S. R. DUNLOP & R. BUICK
repeated arenite-to-lutite cycles. Within these matrix-support fabric is developed. Most arenites coarse cycles, individual laminae are also graded. have 40-50% void space (now filled with silica) Most volcanic arenite beds are cross-bedded on but this may be reduced to about 20% by coma small scale, with sets composed of graded paction, which leads to face—rather than pointlaminae. Trough cross-beds have gently curved contact between grains. sets at moderate angles of inclination, composed A typical sample of trpugh cross-bedded of graded laminae (1-2 cm thick) in which the arenite has grain sizes ranging from 4 mm to grains fine upward from 4mm to 0.05 mm. 0.5 mm, with a mode of 0.74 mm, a median of Tabular cross-beds also occur, with sets at low 0.64mm and a mean of 0.62mm (n = 318). All angles of inclination, composed of only one or values were derived from the graphical statistical two graded laminae, up to 10cm thick. The formulae of Folk & Ward (1957) and there is a grain-size range is similar to that of the trough slight bias toward finer grain size because the cross-beds. Measurement of foreset dip azimuths measured grains were in a polished slab. and the scale of the cross-beds is limited by the Although in bulk samples the sorting is poor intense surface silicification on outcrops. (grain size range—7 0 units, inclusive graphic Flat laminated beds are also composed of standard deviation—1.6), at any level within a repeated fining-upward laminae in which grain graded lamina a size range of only 2 0 units is size ranges from 5 mm to 0.01 mm. Where flaser, present. Therefore on a microscopic scale, the wavy and lenticular bedding are developed in arenite is well-sorted. The grain-size distribution zones transitional into overlying lutites, the is very close to normal (inclusive graphic skewarenitic laminae (1-3 cm thick) rarely show ness = 0.09; graphic kurtosis = 1.06) and a logreverse graded bedding. probability plot produces an almost perfect logThe volcanic arenites have a grain-support normal curve which is only slightly convex upfabric, except in the uppermost parts of grada- ward (Fig. 2). Usually the grains are angular, but tional contacts with overlying lutites, where a some grain types (discussed below) are very well rounded. All grain types are moderately spherical. Bulk sample measurements suggest that these arenites are submature (indicated by the absence of a clay-sized fraction and poor sorting), but at lamina scale the texture is mature. The grains have internal textures which closely resemble many of those seen in the enclosing metavolcanic rocks. Ten varieties have been identified: 1. Clasts (Fig. 3b) with textures similar to those of the groundmass of the finer-grained, upper portions of spinifex-textured flows, with radiating clusters of 1 mm long pseudomorphs after clinopyroxene (Fig. 3a). These clasts are moderately common, but there are no recognizable sedimentary products of the more coarse grained interiors of these flows. 2. Clasts with a microcrystalline silica groundmass and acicular ilmenite grains (Fig. 3d), which have previously been interpreted as rutilated quartz grains (e.g. Dunlop, 1978; Lambert et al., 1978; Barley et al., 1979). The texture of these clasts is identical to those seen in some subophitic metabasalts (Fig. 3c). 1 1 1 -2 - 1 0 1 2 3 4 5 3. Lenticular to subcircular clasts with concentric brown or black laminae (0.1-0.3 mm) i defined by differing amounts of very fine Fig. 2. Log-probability plot of grain-size distribution of silicified volcanic epiclastic arenite from the grained opaque minerals (Fig. These chert-barite unit at North Pole: 318 grains clasts often have a dark, ovoid core of silica, were measured at the intersections of a 1 mm chlorite and opaque minerals which is more grid on a polished slab. coarse grained than the outer laminae. The
EPICLASTIC SEDIMENTS FROM NORTH POLE, PILBARA
. 3.
229
Photomicrographs of altered metavolcanics enclosing the chert-barite unit (a, c, e, g, 1) and texturally similar grains from arenites within the chert-barite unit (b, d, f, h, j). Other grains in the aremtes (k, 1, m, n) are not directly comparable to any textures observed in the metavolcanics, but may come from the same source. Magnification: a, b, c, i, j, m x 13.3, d x 23.4, e, f, g, h, k, 1, n x 33.4.
230
J. S. R. DUNLOP & R. BUICK
mineral content makes this origin improbconcentric laminae usually mimic the outline able. It is more likely that the clasts are of the clast (Fig. 3j), but some have more ireroded amygdales, with the internal circular regular laminae. This clast type can be comstructures and concentric laminae being freepared with the lenticular and subcircular space growth features. structures in the perlitic zone of pillow margins (Fig. 3i). Though the clasts do not 10. Clasts with a cellular structure (Fig. 3m), often with more than 50% intragranular possess the epidote-pumpellyite cores seen in void space now filled with silica. The voids the least-altered perlitic metabasalts, the within grains are usually circular, but some dark, granular central bodies may be relics are lobate, elliptical or dumbbelLshaped, of these features. and vary in size from 2mm-0.1 mm. Most 4. Clasts with flat or irregular laminae which voids are lined with a very thin layer of are truncated by the clast margins. This type opaque minerals. The original very fine is intergradational with the more irregular grained groundmass has been replaced by members of the preceding clast variety. The microcrystalline silica. Some of the smallest clasts have textures similar to those seen in grains are shard-like, void-free, bounded by the laminated part of the perlitic zone of two or three concave surfaces and are compillow margins. posed entirely of silicified groundmass. 5. Clasts which contain structures similar to These possibly formed by in situ break-up of pseudomorphs after large skeletal olivine larger clasts. The clasts usually make up crystals (Fig. 3f). Although olivine-phyric 1-10% of the grains in the volcanic arenites, textures are common in the interior part of but some thin (less than 20 cm thick) beds are the variolitic zone of pillows (Fig. 3e), clasts composed entirely of these extremely with similar textures are rare in the arenites. angular grains. Such rocks are very poorly 6. Clasts with microlitic textures (Fig. 3h) sorted with grains ranging from 10 mmwhich resemble the felted microlitic texture 0.1 mm in diameter, have submature textures in the variolitic zone of pillow margins (Fig. and lack matrix, being cemented by micro3g). crystalline silica which appears to have 7. Clasts containing small (0.1-0.5 mm), replaced an earlier coarse-grained blocky sharply defined, circular structures filled cement. No evidence of pressure solution or with coarse silica crystals set in a groundcompaction can be seen. The clasts are idenmass of very fine grained, mottled silica and tical in texture to scoriaceous volcanic rocks. chlorite with minor scattered subhedral opaque grains (possibly magnetite) (Fig. 3n). The predominant mineral of all clast types is The circular structures which make up silica, with variable carbonate (up to 50%), up to 4-20% of the clast, appear identical to silici- 10% opaque minerals (mostly iron and titanium fied vesicles seen in the variolitic z@ne of oxides), up to 10% chlorite and rare pumpellyite. pillow margins. The ratio of the minerals present seems to be 8. Clasts containing diffusely margined, sub- determined more by the state of alteration than circular structures (0.1-0.5 mm) filled with by the original mineral assemblages of the grains. coarse silica crystals set in a groundmass There is usually no intergranular matrix similar to that of type 7. These structures present, although in the uppermost sections of may be poorly preserved vesicles or olivine beds which grade into lutite there may be a grey, pseudomorphs. This is the most common silt- to clay-sized matrix which is now silicified, clast variety (Fig. 31). but was perhaps originally carbonate (Dunlop, 9. Circular clasts (0.5-1 mm diameter) with a 1978). This matrix composes up to 50% of the thin concentric surface crust of opaque rock. Instead, most volcanic arenites have been minerals and internal coarse-grained and cemented by material which is now entirely silica. cryptocrystalline silica, carbonate and minor It is uncertain whether primary silica cement was chlorite (Fig. 3k). Some have a variably posior a precursor carbonate cement was tioned secondary internal circular structure, developed during late diagenesis (Dunlop, 1976). and may have poorly defined internal con- silicified In some sediments, the grains had a fringe of centric laminae of chlorite and opaque fibrous cement, in which crystals minerals. The origin of these clasts is were normal to clast surfacestheandoriginal about the same unknown. It appears that they are not accre- length. The fringe is defined by a layer of opaque tionary lapilli, having no volcaniclastic core deposited at the cessation of this phase or circumferentially disposed microcrystal- minerals of cementation. Subsequently, the remaining lites. They are possibly ooids, but their pore space has been filled by coarser grained irregular internal structure and mafic blocky cement. Most sediments, however, appear
EPICLASTIC SEDIMENTS FROM NORTH POLE, PILBARA
to have undergone only the latter stage of cementation. The early lithification of these sediments is indicated by the occurrence of tabular clasts (5 cm x 1 cm) of volcanic arenite within polymictic conglomerates conformably and gradationally overlying some arenite beds. Stylolites subparallel to bedding are present in some samples. A few are responsible for the emplacement of other lithologies against the arenites and some arenitic clasts have been embayed during the pressure-solution event. Because silicified calcilutites display much greater stylolitization than primary silica sediments in the chert-barite unit, pressure-solution probably occurred before the major episode of silicification. The alteration history of these rocks is very complex. Subsequent to deposition, the arenites have undergone low-grade metamorphism, accompanied by moderate carbonation and a high degree of silicification. The sequence of these events is uncertain. Silicification and carbonation were not restricted to single episodes. CONCLUSIONS Most grains in the volcanic arenites were derived from basalts; both the preserved textures of the clasts and their mineral assemblages are consistent with this source. The scoriaceous clasts may be derived from the rare scoriaceous flow tops within the mafic metavolcanic sequence, but may also be a product of felsic volcanism. In other parts of the Pilbara Block, felsic volcanism occurred throughout the deposition of the Warrawoona Group (Barley et al., 1979) and, although subaerial eruption which produced pumice was rare in the older felsic volcanic units of the Warrawoona Group (Barley, 1980), it is conceivable that pumice from distant eruptions floated long distances before beaching (cf. Binns, 1972) or settling in the depositional basin of the North Pole volcanic arenites. The type 9 spherical clasts may represent ooids or eroded amygdales from a felsic or mafic source. A predominantly singlesource origin of the grains is supported by the near log-normal distribution of grain sizes from a typical bulk sediment sample (Fig. 2). The textures of most grains are similar to those of marginal zones of pillows or the upper margins of flows; fragments of coarse spinifex textures from flow interiors are not present and grains with olivine-phyric textures from the inner part of the pillows are rare. Subophitic-textured grains may be derived from the margins of the rare unquenched flows and pillows. The submature texture of whole-rock samples and the angularity of the grains suggest that their source was not far from the site of deposition.
231 The dominant bedforms developed during the deposition of the volcanic arenites were probably current megaripples (Reineck & Singh, 1975) of various morphologies, which produced the trough and tabular cross-beds. The maximum grain size (5 mm) indicates that current velocity was probably in the high part of the megarippleproducing range and that the less abundant plane bedding was produced when current velocities fell short of the megaripple range. Current velocity probably fluctuated rhythmically to produce the repeated graded laminae in the cross and plane beds. The absence of clay-sized sediment at the top of the graded laminae suggests that if any very fine grains were deposited during the latest stage of current waning, they were eroded by the initial high energy phase of the succeeding current pulse. Larger arenite-to-lutite cycles (2-5 m thick) indicate that the currents fluctuated over longer periods as well, gradually waning with a sudden increase in current velocity at the start of a new cycle. The poor sorting, angularity of grains and lognormal grain-size distribution of bulk sediment samples indicates that the material now in the sediment has not been reworked. This, accompanied by the similar orientation of cross-beds in large samples, suggests that the volcanic arenites were deposited by unidirectional currents, although a bimodal current regime with one much weaker current may produce similar features. Visher (1969), in studies of modern arenites, concluded that sediments with sorting and grainsize parameters similar to those displayed in the log-probability grain-size distribution of Figure 2 are most likely to have been deposited in fluvialor tidal-dominated environments. The sedimentary structures present in the volcanic arenites and the current regimes which could have produced them are also best explained by a fluvial or tidal environment of deposition. The absence of many typical fluvial features such as channel lag deposits, pebbly channel- and point-bars, filled channels, associated levee and floodplain deposits and the lack of diverse ripple crossbedding suggests that fluviatile deposition is the less likely alternative. A tide-dominated environmental model explains many of the unusual features of the .volcanic arenites. The two scales of graded bedding could have been produced by an ebb-flood cycle (graded laminae) and a spring-neap cycle (arenite-to-lutite cycles) of longer periodicity. The variability of the sedimentary structures may be due to overprinting of a regular tidal pattern by storm activity or changes in current patterns produced by altering sea-floor topography. Time-
J . S. R . D U N L O P & R .
232
BUICK
equal to that of deposition. However, during the velocity asymmetry of the tide system could proformation of the shoaling sequences of volcanic duce the apparent single direction of current arenites and lutites, basin subsidence must have flow, as is common in modern restricted basins ceased locally. (Klein, 1977). Although rapid erosion of mafic volcanics and Previous studies have provided compelling evidence for a shallow-water environment of deposition of epiclastic volcanic sediments occurs deposition for the chert-barite unit at North Pole when mafic flows are subjected to the erosive (Dunlop, 1976; Dunlop, 1978; Dunlop et al., rigours of a shallow-water environment (e.g. For1978; Lambert et al., 1978; Barley et al., 1979; nari et al., 1979), epiclastic sediments derived Walter et al., 1980). These studies suggested that from the mafic volcanic sequence are rare in the the unit represents the sedimentary products of a North Pole Dome (except in the chert-barite unit) series of small, shallow evaporative basins and have not been reported from elsewhere in the bordering a larger shallow marine basin in an Pilbara Block. This suggests that volcanic eruparea of low topographic relief. If a tidal environ- tion was almost continuous and the volcanic pile ment of deposition of these sediments is accepted, developed rapidly, with few hiatuses marked by the volcanic arenites probably represent a shallow erosion. Lowe (1980) considered that this was subtidal sand body. Shoaling eventually pro- general in the mafic parts of volcanic sequences duced conditions in which features of intertidal in early Archaean greenstone belts, resulting in environments were developed at the top of some the development of clastic-starved sedimentary arenite beds (e.g. flaser, wavy and lenticular bed- environments. Thus, the chert-barite unit at North Pole, with ding). Commonly, the volcanic arenites are overits large amounts of volcanic epiclastic sediment lain by lutites which have features suggestive of subaerial exposure (Barley et al., 1979) and are of mafic provenance, appears to be a product of interpreted as intertidal to supratidal mudflat unusual localized environmental conditions. Perdeposits at the top of a shoaling sequence. The haps the other remarkable features of this unit, presence of beds composed entirely of type 10 i.e. the development of major Archaean evaporscoriaceous clasts gives added support to inter- ite deposits and the preservation of very ancient tidal-supratidal conditions at some stages during life-forms also depended upon the presence of the deposition of the chert-barite unit. As the these rare environmental circumstances. If so, it clasts were probably able to float, most would would seem that detailed environmental investihave been deposited on beaches above normal gations of Archaean sediments would provide more worthwhile information about such interhigh water mark. The Warrawoona Group succession at North esting subjects than has been gained from the rather haphazard approach which has previously Pole is reported to be approximately 12 km thick (Hickman & Lipple, 1975). Barley et al. (1979) prevailed. The authors acknowledge receipt of Commonhave demonstrated that shallow-water conditions prevailed over much of the period of deposition wealth Postgraduate Research Awards, and of the Warrawoona Group throughout the thank M. E. Barley, R. G. Brown, M. G. eastern Pilbara Block. Reconciliation of these Dunlop, J. E, Glover, D. I. Groves, and D. M. observations requires basin subsidence at a rate McConchie for diverse assistance. REFERENCES
BARLEY, M. E., 1980: Evolution of Archaean calcalkaline volcanics. Ph.D. Thesis, Univ. West. Aust. [unpublished]. BARLEY, M . E . , DUNLOP, J . S . R . , GLOVER, J . E . , &
GROVES, D. I., 1979: Sedimentary evidence for'an Archaean shallow-water volcanic-sedimentary facies, eastern Pilbara Block, Western Australia.
Earth planet. Sci. Lett., 43, 74-84. BINNS, R. E., 1972: Composition and derivation of pumice on postglacial strandlines in northern Europe and the western Arctic. Bull, geol Soc Am., 83, 2303-2324. DIMROTH, E . , COUSINEAU, P . ,
LEDUC, M . , & SANS-
CHAGRIN, Y., 1978: Structure and organization of Archaean subaqueous basalt flows, Rouyn- Noranda area, Quebec, Canada. Can. J. Earth Sci 15 902-918. " '
DUNLOP, J. S. R., 1976: The geology and mineralization of part of the North Pole barite deposits, Pilbara Region, Western Australia. B.Sc. (Hons) Thesis, Univ. West. Aust. [unpublished]. , 1978: Shallow-water sedimentation at North Pole, Pilbara Block, Western Australia. Pubis
geol. Dep. & Extension Service, Univ. West. Aust,, 2, 30-38. DUNLOP, J . S . R . , & GROVES, D . I . , 1978: Archaean
evaporitic sulphates from the North Pole barite deposits, Pilbara Block, Western Australia. Geol.
Assoc. Can., Mineral. Assoc. Can., geol. Soc. Am. Abstracts with Programmes, 10, 393. DUNLOP, J . S . R . , M U I R , M .
D . , MILNE, V. A . , &
GROVES, D. I., 1978: A new mierofossil assemblage from the Archaean of Western Australia. Nature,
Lond., 174, 676-678.
EPICLASTIC SEDIMENTS FROM NORTH POLE,
FOLK, R. L., 1968: Petrology of Sedimentary Hemphill's Book Store, Austin, Texas.
PILBARA
233
Rocks.
LAMBERT, I. B . , DONNELLY, T . H . , DUNLOP, J . S. R . , &
FOLK, R. L., & WARD, W . C . , 1957: Brazos River bar: a
GROVES, D. I., 1978: Stable isotopic compositions of early Archaean sulphate deposits of probable evaporitic and volcanogenic origins. Nature, Lond., 276, 808-811.
study in the significance of grain size parameters. J. sediment. Petrol., 27, 3-26. FORNARI, D . J . , MOORE, J . G . , & CALK, L . , 1979: A
large submarine sand-rubble flow on Kilauea volcano, Hawaii. J. Volcanol. Geotherm. Res., 5, 239-256. HAMILTON, P . J . , EVENSEN, N . M . , O'NIONS, R . K . , GLIKSON, A . Y . , & HICKMAN, A . H . , 1981: S m - N d
dating of the North Star Basalt, Warrawoona Group, Pilbara Block, Western Australia. Spec. Pubisgeol. Soc. Aust., 7, 187-192. HICKMAN, A. H., 1973: The North Pole barite deposits, Pilbara Goldfield. Ann. Rep. geol. Surv. West. Aust. for 1972, 57-60. HICKMAN, A . H „ & LIPPLE, S. L . , 1975: E x p l a n a t o r y
notes on the Marble Bar 1:250000 Geological Sheet, Western Australia. Rec. geol. Surv. West. Aust., 1974/20. JONES, J. G., 1969: Pillow lavas as depth indicators. Am. J. Sci., 267, 181-195. KLEIN, G. DE V., 1977: Tidal circulation model for deposition of clastic sediment in epeiric and mioelinal shelf seas. Sediment. Geol., 18, 1-12.
LOWE, D. R., 1980: Archaean sedimentation. Rev. Earth planet. 5c/., 1980 (8), 145-167.
Ann.
PIDGEON, R. T., 1978: 3450m.y. old volcanics in the Archaean layered greenstone succession of the Pilbara Block, Western Australia. Earth planet. Sci. Lett., 37, 421-428. REINECK, H . E . , & SINGH, I. B . , 1975:
Sedimentary Berlin.
Environments.
Depositional
Springer-Verlag,
SANGSTER, D. F., & BROOK, W. A., 1977: Primitive lead in an Australian Zn-Pb-Ba deposit. Nature, Lond 270, 423 . VISHER, G. S., 1969: Grain size distributions and depositional processes. sediment. Petrol., 39, 1074-1106. WALTER, M . R . , BUICK, R . , & DUNLOP, J . S. R . , 1980:
Stromatolites 3400-3500 Myr old from the North Pole area, Western Australia, Nature, Lond., 284, 443-445.
ARCHAEAN PLATFORM-TO-TROUGH SEDIMENTATION, EAST PILBARA BLOCK, AUSTRALIA* K. A. Eriksson
Programs in Geosciences, University of Texas at Dallas, Richardson, Texas 75080, U.S.A. Present address: Department of Geological Sciences, Virginia Polytechnic Institute, Blacksburg, Virginia24061, U.S.A. ABSTRACT Fades patterns and palaeocurrent determinations in the east Pilbara Block have led to the recognition of two separate areas of contemporaneous platform (alluvial) to trough (submarine fan) sedimentation in the upper sedimentary interval of the Archaean greenstone sequence. The shallow- to deeper-water transition in both areas was abrupt to the exclusion of any shallow-marine shelf facies. In the Gorge Creek-Pilgangoora-Soanesville belt the transition was from the north and northeast to the southwest whereas in the Budjan Creek-Mosquito Creek belt it was from the north to the south and southeast. Although it cannot be stated unequivocally that sedimentation of these two belts was contemporaneous, it is proposed that a submarine trough extended at least around the western and southern margins of an east Pilbara protocraton during deposition of the terrigenous clastic sedimentary interval. The absence of shallow marine facies in the east Pilbara implies continental to marine transition along a narrow continental shelf. This sedimentation style is similar to that in the Fig Tree and Fig Tree-Moodies Group transition in the Barberton Mountain Land for which a rift control has previously been postulated, and is also analogous to that recognized in the Canadian Archaean. In the Barberton Mountain Land a steep-rift margin was succeeded by the development of a stable continental shelf through outbuilding of a turbidite wedge, probably associated with a eustatic rise in sea-level. On that shelf the extensive tidal-flat, deltaic and barrierbeach sediments of the Moodies Group accumulated.
INTRODUCTION Detailed documentation of lithologies, primary sedimentary structures and vertical and lateral interrelationships of facies in Archaean sedimentary sequences allows interpretation of depositional processes and thereby recognition of the most ancient sedimentary environments. In so far as the composition of the source area, the nature of terrestrial to marine transitions and the morphology of the depository can be defined, sedimentological studies of this type significantly constrain models of Archaean crustal development. Such an approach in the Barberton Mountain Land, South Africa has led to the recognition of a mixed but notably potash-rich source area supplying sediment to a continental margin (Eriksson, 1980), thus requiring modification of previously postulated crustal models (e.g. see Anhaeusser, 1973; Hunter, 1974; Tarney et at., 1976). Sedimentological studies are being undertaken at an early stage in the investigation of the Pilbara Block in Western Australia and will be
important in constraining future models of crustal development based on petrographic, geochemical, structural* geochronologic and geophysical criteria. Barley et al. (1979) have already established, on the basis of a sedimentological study of intercalated cherty sediments, that shallow-water conditions existed throughout the development of the ca 3500 Ma basal volcanic Warrawoona Group in the eastern Pilbara Block (Fig. 1). This paper focusses on the dominantly terrigenous sedimentary interval overlying the volcanics, which is variously termed the Gorge Creek, Budjan Creek and Mosquito Creek Group (Fig. 1). In most areas no major time-break is recognizable between the basal volcanic and upper sedimentary intervals, suggesting that the terrigenous sediments are not much younger than 3500 Ma. PALAEOENVIRONMENTS IN THE EAST PILBARA BLOCK Braided alluvial and submarine-fan deposits predominate in the upper sedimentary interval in
* Contribution No. 382, Programs in Geosciences, The University of Texas at Dallas. Spec.
Pubis geol. Soc. Aust., 7 (1981)
236
K. A. ERIKSSON PT| Proterozoic P7I Gorge Creek-Budjan CreekMosquito Creek Groups vj Warrawoona Group F H Granites-Gneisses ! Boobina Porphyry 10
0
Kilometres
-22°00
Fig.
Locality map and general geology of the east Pilbara Block. G: Gorge Creek; LR: Lalla Rookh; P: Pilgangoora; S; Shaw Gorge; Sn: Soanesville; BC: Budjan Creek; MC: Mosquito Creek.
the east Pilbara Block and respectively overlie lacustrine deposits and deeper-water, basin-plain banded iron formations. The alluvial and submarine-fan sequences are generally separated in space such that braided alluvial sediments occur in the Gorge Creek and Budjan Creek areas, and submarine-fan deposits in the PilgangooraSoanesville and Mosquito Creek areas (Fig. 1). Vertical stacking of sediments assigned to these two depositional environments is rare. ALLUVIAL DEPOSITIONAL ENVIRONMENTS Upper Braided Alluvial Plain Sediments of upper braided alluvial-plain origin comprise conglomerates with subordinate sandstones and occur in the Lalla Rookh, Shaw Gorge and Budjan Creek areas (Fig. 1). The conglomerates are framework-supported and contain pebbles up to 40 cm in diameter which are moderately to well rounded and of diverse composition (Table I). Associated sandstones, including the matrix to the conglomerates, are arkosic. Complex vertical and lateral interrelationships of conglomerates and sandstones consist of overlapping and mutually erosive lenses (Fig. 2a). The conglomerates are interpreted as longitudinal bar deposits. Bars of this type develop from a mid-
channel nucleus which forms during the early stages of waning flood discharge when the sediment load exceeds the carrying capacity of the stream (Boothroyd & Ashley, 1975). Such bars are coarsest on their upstream edges and often sandy downstream where bar-edge sand wedges develop (Miall, 1977). The intercalated planar cross-beds in Figures 2a and 3 are apparently of this origin. With continued waning in discharge, a stage is reached when pebbles are no longer in motion. Instead, sand is transported across bar surfaces as upper flow-regime plane-beds and, at TABLE I
Percentage clast compositions Black chert
1
2
3
4
.39
36
6
34
14
45
' 33 ' 1-0 '''
2
4
10
6
White chert
21
Banded chert Vein quartz
1
Acid- volcanic Mafic volcanic
14
*
Tuff
2
Banded iron formation
4
Porphyry Quartz arenite
'
65 '
37
1. Lalla Rookh 2. Shaw Gorge 3. Lower Conglomerate: Budjan Greek 4. Upper Conglomerate: Budjan Creek
7
7
:
I ;
PLATFORM-TO-TROUGH SEDIMENTATION, PILBARA La I la Rookh -Budjan Creek
Gorge Creek Shaw Gorge
m
237
Budjan Creek
4
b
l ° 0 ° l Conglomerates ^ Plane-Beds Trough Cross-Beds Y/A Planar Cross-Beds
Fig, 2. Depositional sequences for: (a) upper and (b and c) mid to lower braided alluvial environments. lowest discharge, in interbar channels as lower flow-regime dunes. These processes lead, respectively, to the formation of plane-beds and trough cross-beds (Fig. 2a). Lateral switching of bars and channels, a characteristic of braided alluvial plains, produces interbedded conglomerates and sandstones. This association is common in the geological record and is characterized as the Scott model (Miali, 1977). Mid to Lower Braided Alluvial Plain Sandstones with sporadic thin conglomerates are the dominant lithology in the Gorge Creek area where they comprise upward of 2 km of stratigraphic thickness. Sediments of this type also occur in the Lalla Rookh, Shaw Gorge and Budjan Creek areas (Fig. 1). The sandstones are arkosic in all areas whereas associated conglomerates contain a mature assemblage of pebbles dominated by vein quartz and chert. Immediately above the Warrawoona Group at Budjan Creek the pebbles are immature in rounding and composition, reflecting local derivation from the underlying felsites. Various types of vertical sequences of lithologies and sedimentary structures are evident in these braided alluvial sandstones. With minor variations, two characteristic sequences (Fig. 2b and c) are considered to reflect differing depositional processes on the mid to distal alluvial plain. The sequence in Figure 2b is analogous to that documented from the South Saskatchewan River by Cant (1978). With waning discharge vertical aggradation occurs within channels via dune migration producing thick trough crossbedded sandstone intervals which commonly overlie basal pebble lags (c/. Fig. 2b). Sand flats are produced upon emergence of channel bars and are environments where small-scale bed-
forms migrate at low discharge. The upper 50 cm represented in Figure 2b was probably developed in this way. In contrast to the pronounced aggradation inferred for the sequence in Figure 2b, that shown in Figure 2c formed by downstream accretion of trains of linguoid or transverse bars, often across channel lag gravels. The process is typified by the Platte River in Colorado where stacked planar cross-bed sets are produced during single flood cycles (Miall, 1977). The palaeocurrent pattern for planar cross-beds at Budjan Creek (Fig. 2c) similarly reflects unidirectional accretion of bars, in this instance to the south. The small-scale cross-beds capping the sequence probably developed on bar tops during the final waning stages of floods. Fine-grained terrigenous clastic sediments are rare in these Archaean fluvial sediments and only one unit of siltstone-shale was noted. Braided rivers are characteristically high-energy environments: as a result fine-grained sediments can only accumulate on abandoned reaches of the alluvial plain or, more commonly, are flushed out of the fluvial system into receiving basins. LACUSTRINE DEPOSITIONAL ENVIRONMENT At the base of the Gorge Creek and Budjan Creek Groups in the type areas, sediments of probable lacustrine origin are developed as isolated bodies from a few to 150 m in thickness with strike lengths of up to 4 km. The dominant sedimentary structures are wave ripples preserved either as symmetrical or asymmetrical forms in plan-view or, as at Budjan Creek, in crosssection. The latter are characterized by complex internal organization of cross-laminae and
238
K. A. ERIKSSON
Fig. 3.
Fluvial conglomerate containing an intercalated planar cross-bed set which developed on margin of longitudinal bar.
Fig. 4.
Wave ripples showing complex internal organization of cross-laminae and irregular lower bounding surfaces.
PLATFORM-TO-TROUGH SEDIMENTATION, PILBARA
239 characteristic facies association is of lensoid conglomerate-sandstone bodies from 200-1500 m wide and 25-500 m thick enclosed within shale and banded iron formation with subordinate limestone (Fig. 5). The coarse-grained terrigenous clastic lenses are randomly dispersed, increase in size and coarsen upward through a stratigraphic thickness of 2.5 km. Associations of this type characterize submarine fan channel and interchannel environments (Ingersoll, 1978; Mutti & Ricci Lucchi, 1978). The stratigraphic development can be related to progressive channel infilling, abandonment through lateral switching, and envelopment of channel complexes in fine-grained terrigenous clastic and chemical sediment. The multitude of channel sequences suggests the mid fan as the most likely depositional setting, being an environment with numerous bifurcating channels off the trunk inner-fan channel (Mutti & Ricci Lucchi, 1978). The upward-coarsening of successive channel complexes indicates progradation through time of the mid fan. Internally, the channel sequences consist of erosively-based, channelized conglomerate-sandstone couplets up to 6m thick (Fig. 5). ConSUBMARINE FAN DEPOSITIONAL glomerates in the uppermost channels reach 4 m ENVIRONMENTS in thickness and contain both well-rounded extraMid Fan basinal and angular intrabasinal rip-up clasts Sediments interpreted as being of mid-fan (Fig. 6). The conglomerates display both normal and inverse grading as well as imbrication. origin occur in the Pilgangoora area (Fig. 1). The
irregular lower bounding surfaces (Fig. 4). Also present are second-order ripples in the troughs of the symmetrical forms, while associated planebedded surfaces with primary current lineations and wrinkle marks suggest shallow water. Evidence of wave processes by themselves, to the exclusion of tides, is one characteristic of lacustrine deposits (Collinson, 1978). However, the lack of more definitive characteristics such as fauna! or chemical makes the lacustrine interpretation debatable. Notwithstanding, an environment of low energy is indicated Four upward-coarsening, progradational depositional cycles occur in the inferred lacustrine interval at Budjan Creek. Braided alluvial sediments of the type shown in Figure 2c cap wave-rippled sandstones and siltstones in each of the cycles. A fifth cycle, which is overlain by thick braided alluvial conglomerates and sandstones, has thinly-bedded turbidites in place of the wave-rippled fades indicating that, with time, the depositional interface in the lake descended below wave base. The significance of this is reviewed in the palaeogeographic interpretation.
200-1500m
o o 3
ES9
Conglomerate
E l
MM
Sandstone
E H J Horizontal
Shale
EH03 P l a n e - B e d s
E23 Banded Iron Formation
E 3
Starved Ripples
I S
Convolute Lamination Flame Structures
Lamination ES
Sole Marks
Graded Beds
Fig. 5. Morphology and internal characteristics of a mid-fan channel deposit enclosed within basin-plain shales and banded iron formations.
240
K. A. ERIKSSON
Fig. 6.
Flg
7
Mid-fan conglomerate containing rounded extrabasinal and angular intrabasinal rip-up clasts.
^ n t o o n e i 0 n t h r ° U g h p o r t i o n o f o u t e r - f a n depositional lobe showing upward increase in proportion of
PLATFORM-TO-TROUGH SEDIMENTATION, PILBARA Associated sandstones are composed of graded 1-3 cm depositional units. The channel complexes fine outward and upward into more sandy facies, where well-developed upward-fining depositional sequences vary from l - 3 m in thickness, and consist of pebbly, massive and plane-bedded sandstone, often with flame structures at the base of beds. Separating the coarse-grained channel complexes from the enclosing fine-grained sediments are well-developed levee deposits consisting of horizontally-laminated shale with thin graded siltstone laminae and starved ripples. Softsediment deformation is common in the form of small-scale flame and ball-and-pillow structures. Sedimentation in submarine channels of the type discussed here takes place from highconcentration turbidity currents a n d / o r grain flows (Mutti & Ricci Lucchi, 1978). The Archaean channel sediments accumulated in response to pulsating gravity flows as indicated by the abundant evidence of waning velocities. The interchannel sediments are totally devoid of coarse clastics, implying that high-energy gravity flow processes were confined to the channels. The interchannel muds developed as a result of normal background suspension sedimentation. Associated chemical sediments, most notably banded iron formation, developed in those parts of the depository removed from clastic input.
241
Outer Fan-Basin Plain Upward- and inward-coarsening lensoid depositional sequences, consisting of interbedded sandstone and shale occur in the Soanesville and Mosquito Creek areas (Figs 1 and 7). Individual lenses vary from 100-300 m in width and 25100m in thickness (Fig. 8). According to the submarine-fan model of Mutti & Ricci Lucchi (1978), these lensoid bodies can be interpreted as outer-fan depositional lobes. A complex internal arrangement of Bouma intervals is recognizable in these lenses (Fig. 8). In the centre and especially in the upper half of the lenses, shallow channels contain " a - e " and a few complete Bouma sequences. Flute casts are common at the base of beds. Individual sequences average 60cm but are up to 2 m in thickness. Sandstone predominates with the argillaceous " e " intervals generally less than 2cm thick. These channel-fill deposits interfinger laterally and are interbedded with " b - d - e " a n d / o r "b-cd-e" sequences in which thicker " e " intervals are developed. An interchannel origin is inferred for these incomplete Bouma sequences. The channel and interchannel sediments were deposited, respectively, from high- and low-concentration turbidity currents. The stratigraphic sequences in the Soanesville and Mosquito Creek areas consist of overlapping lensoid bodies of the type described above. Active
Fig. 8. Morphology and internal characteristics of outer-fan depositional lobe showing upward and inward increase in proportion of sandstone.
K. A. ERIKSSON 242 progradation of outer fan lobes followed by Warrawoona Group. Depositional processes were abandonment and lateral switching of the deposi- dominated by weak wave reworking but evidence tional locus characterizes outer fans and accounts for foundering of the floor of the rifts is present for the overlapping of depositional sequences. In in the Budjan Creek area where the numerous general, the size and coarseness of individual stacked progradational cycles indicate progreslenses increases upward through the stratigraphy, sive deepening of the lakes. With continued riftwith the uppermost lenses often overlain by ing, possibly associated with spreading, a subcoarse-grained to pebbly sandstone intervals (Fig. marine, deeper-water trough developed at least 8). There are a few conglomerates which pinch- around the southern and western margins of the out laterally over 50-100 m. An upward-fining more stable east Pilbara protocraton. Thereafter and upward-thinning of graded depositional units contemporaneous platform (alluvial) to trough is developed in these intervals which are inter- (submarine fan) sedimentation prevailed in the preted as lower mid-fan channel deposits formed east Pilbara. On the protocraton, braided-river sediments rapidly filled the early grabens and during progradation across the outer fan lobes. The outer-fan lobe deposits coarsen upward were supplied to the trough, where submarine and inward from shale with a few thin, but later- fans prograded across basal banded iron formaally persistent sandstone horizons. The sandstone tions. The chemical sediments are considered to beds vary from 5-20cm in thickness, contain have developed on the floor of the troughs in well-developed groove casts at their base, and response to waning volcanism. The inferred consist of graded "b-c-d" and "c-d" Bouma spatial interrelationships of the various alluvial intervals. Thin banded iron formation layers are and submarine fan subenvironments are outlined frequently intercalated within shale in the Soanes- on Figure 9. The platform-to-trough transition ville area. This association of facies characterizes was to the south and southeast in the Budjan the basin plain which is an environment sur- Creek-Mosquito Creek belt and to the southwest rounding, and extending basinward of the outer in the Gorge Creek-Pilgangoora-Soanesville belt fan lobes and is marked by suspension sedimenta- (Fig. 1). Gee (1979) similarly proposed the existtion interrupted infrequently by dilute turbidity ence of troughs in the Mosquito Creek area and currents and precipitation of chemical sediments. between the East and West Pilbara regions and an alluvial to submarine-fan transition is also probable off a west Pilbara protocraton. PALAEOGEOGRAPHIC INTERPRETATION Palaeocurrent determinations and facies patterns in the upper sedimentary interval in the east DISCUSSION Pilbara Block indicate that sediment was derived The platform-to-trough transition in the east from source areas to the north and northeast. The Pilbara Block was abrupt to the .exclusion of any composition of the source terrain has significance shallow marine facies. A similar style of sedimennot only with respect to the broader palaeo- tation for which a rift control has been postulated geography, but especially in contributing to our prevailed in the Barberton Mountain Land, knowledge of the early crust. Clasts in the upper South Africa, during the development of the Fig sedimentary interval indicate reworking of the Tree Group and Fig Tree-Moodies Group transiWarrawoona Group, and include volcanic and tion where braided alluvial sediments prograded chert pebbles from the Taiga Taiga and Salgash across earlier submarine fan deposits (Eriksson, Subgroups, quartz arenite pebbles from reworked 1980). In the Barberton Mountain Land, howfelsic pyroclastics in the Warrawoona Group, ever, the initial steep-rift margin was succeeded porphyry pebbles from the Boobina Porphyry, by the development of a shelf-rise margin with a and banded iron formation clasts from the top of well-developed continental shelf. This morphothe Warrawoona Group and possibly the base of logical change was achieved through outbuilding the upper sedimentary interval. Granitic clasts of a turbidite wedge and, by analogy with Holoare absent but the arkosic composition of all cene coastlines (Gorsline, 1978), a eustatic rise in sandstone indicates that potassic granite was sea-level. On the continental shelf, extensive widespread in the source area. deltaic, barrier beach and tidal flat sediments The terrestrial and marine depositional en- accumulated. vironments in the upper sedimentary interval in the east Pilbara Block can be related in a geotectonic model which recognizes an early intra- ACKNOWLEDGMENTS cratonic rift-stage evolving into a rifted continenThe author benefitted from discussions with tal margin. During the earliest stages of rifting, Ganapathy Shanmugan on submarine fan lakes developed in local topographic lows on the models. H. Wilhelmij assisted in the field. The
PLATFORM-TO-TROUGH SEDIMENTATION, PILBARA
Fig. 9.
Palaeogeographic model showing the contemporaneous submarine-fan depositional environments.
co-operation of the Geology Department, University of Western Australia is acknowledged. Financial assistance for the project was provided
243
development of braided alluvial and ,
by an Earth Science Section National Science Foundation Grant, NSF Grant No. EAR 7842307.
REFERENCES R., 1 9 7 3 : The evojutiqn of the early (Eds) Glaciofluvial and Glaciolacustrine Sedimentation. Spec. Pubis Soc. Econ. Paleont. and Precambrian crust of South Africa. Phil. Trans, Mineralog., 23, 193-222. ft. §gf. London A 273, 3 1 9 - 3 8 8 . BARLEY, M . E . , D U N L O P , J . S . R . , G L O V E R , J . E . , & C A N T , D. J., 1978: Development of a fades model for GROVES, D . I . , 1 9 7 9 : Sedimentary evidence for an sandy braided river sedimentation: comparison of South Saskatchewan River and the Battery Point Archaean shallow-water volcanic-sedimentary Formation; in Miall, A. D. (Ed.) Fluvial Sedifades, eastern Pilbara Block, Western Australia. mentology. Mem. Can. Soc. petrol. Geol., 5, Earth planet. Sci. Lett., 43, 7 4 - 8 4 . 627-639. BOOTHROYD, J . C . , & A S H L E Y , G. M., 1 9 7 5 : Processes, COLLINSON, J. D,, 1978: Lakes; in Reading, H. G. (Ed.) bar morphology, and sedimentary structures on Sedimentary Environments and Fades, 61-79. braided outwash fans, northeastern Gulf of Blackwell Scientific Publ., Oxford. Alaska; in Jopling, A. V., & McDonald, B. C. ANHAEUSSER, C .
K. A. ERIKSSON
244
K. A., 1980: Transitional sedimentation styles in the Moodies and Fig Tree Groups, Barberton Mountain Land, South Africa: evidence favouring an Archaean continental margin. Precamb. Res., 12, 141-160. GEE, R. D., 1979: Structure and tectonic style of the Western Australian Shield. Tectonophys., 58, 327-369. ERIKSSON,
GORSLINE, D. J., 1978: Anatomy of margin basins. J.
sediment. Petrol., 48, 1055-1068. HUNTER, D . R . , 1974: Crustal development in the Kaapvaal Craton, 1. The Archaean. Precamb. Res., I, 259-294.
INGERSOLL, R. V., 1978: Submarine fan fades of the
Upper Cretaceous Great Valley sequence, northern and central California. Sediment. Geol 21 205-230. MIALL, A . D . , 1977: A review of the braided river depositional environment. Earth-Sci. Rev 13 1-62. " ' M u m , E., & RICCI LUCCHI, F., 1978: Turbidites of the northern Apennines. Internat. geol. Rev., 20 125166.
TARNEY, J . , DALZIEL, I. W . D . , & DE WIT, M . J., 1976:
Marginal basin 'Rocas Verdes' complex from S. Chile: a model for Archaean greenstone belt formation; in Windley, B. F. (Ed.) The Early History of the Earth, 131-146. Wiley, London.
CHEMICAL SEDIMENTATION IN ARCHAEAN-EARLY PROTEROZOIC GREENSCHIST BELTS OF THE DHARWAR CRATON, INDIA
S. M. Naqvi, P. K. Govil, & J. J. W. Rogers National Geophysical Research Institute, Hyderabad 500 007, India Department of Geology 029A, University of North Carolina, Chapel Hill, North Carolina 27514, U.S. A, 1
1
2
1
2
ABSTRACT Three to four successive types of greenschist belt occur in the Dharwar craton (mostly Archaean) of southern India. Volcanic rocks change from primarily ultramafic (komatiitic) in the older belts to tholeiitic and calc-alkaline in the younger belts. Chemical sediments also change from an abundance of high-Al, high-Mg sediments of uncertain origin in the older belts to earbonate and chert sequences in younger belts; fuchsite-bearing quartzites and ironstones (including banded magnetite quartzites) are found in belts of all ages. Abundances of trace metals are related to the types of volcanism, with oxides predominant in older belts and sulphide minerals abundant in the youngest belt. Limestones are characterized by high Na/K ratios and low Sr contents relative to younger limestones. Cherty and iron-rich sediments have Ni/Cr ratios similar to those postulated for the primitive Archaean mantle. These features are probably best explained by volcanogenic (exhalative) processes, which occurred throughout the change in crustal structure caused by the development of sialic (cratonic) areas of gneisses and tonalite-trondhjemite bodies, about 3300m.y. ago. Before development of these cratonic elements, belts were ensimatic; after, both plat formal and deep-basin varieties of belts were formed.
INTRODUCTION A large variety of chemical sediments occurs in the greenschist belts of the Dharwar craton of southern India. The Dharwar craton, of dominantly Archaean age, consists largely of granitegneiss terrain surrounding mafic belts (broadly referred to as the "Dharwars") which apparently range in age from more than 3500 m.y. to less than 2200m.y. Rocks in the schist belts range from greenschist to amphibolite fades and show great structural and stratigraphic complexity. The Dharwar craton apparently contains three to four successive sequences of greenschist belts (Table I and Fig. 1). Different sediments are confined to different age groups. Chemical properties of the sediments, however, show some broadly similar features that permit interpretation of the history of the craton. Both exogenic and exhalative-volcanogenic sediments are apparently present. This paper provides a preliminary interpretation of the significance of these sediments. GEOLOGICAL SETTING OF CHEMICAL SEDIMENTS IN GREENSCHIST BELTS Regional correlation of the different greenschist belts of the Dharwar craton is given in Spec. Pubis geol. Soc. Aust., 7 (1981)
Table I. The oldest group is represented by the Holenarsipur schist belt and its homotaxial units, such as Nuggihalli, Krishnarajpet, Hadnur, Ghatti Hosahalli and Banavara (Fig. 1). All phases of the Peninsular gneisses appear to intrude this group. The second phase of intrusive gneissic activity is dated at about 3360m.y. (Beckinsale et al., 1980). Holenarsipur and related belts are composed dominantly of metaultramafites interbedded with a variety of chemical sediments, including: (1) high-Al, highMg rocks constituted by mineral assemblages of chloritoid-magnetite or kyanite-garnet (± staurolite, muscovite, biotite and minor graphite); (2) fuchsite quartzite, interbedded with barite in the Ghatti Hosahalli belt; and (3) banded magnetite quartzite (± garnet, grunerite, actinolite), also referred to as "ironstones". The Javanahalli schist belt (Fig. 1) appears to be younger than the Holenarsipur group, but no age data are available. The Javanahalli belt consists mainly of metasediments such as paraamphibolites and quartzites. Current-bedded meta-arkoses are present at the base of the sequence, and chemical sediments including magnetite quartzite, fuchsite quartzite, chert, carbonates, and calc-silicate rocks abound. All
246
S. M. NAQVI, P. K. GOVIL & J. J. W. ROGERS
Fig. 1. Simplified geological map of the southern low-grade Dharwar era-ton (modified after Swami Nath et al 1974, 1976). The map shows the distribution of different successive greenschist belts (greenstone belts and geosynclinal piles of Naqvi et al1978c). Sample locations are also shown.
sediments are metamorphosed to amphibolite fades. The youngest member of the group is fuchsite quartzite (Naqvi et al., 1980). The mafic schists associated with the fuchsite quartzites and ironstones have lower A1 0 , Na 0, and K 0 contents and higher CaO content than amphibolitic rocks of the Holenarsipur group and may be metasediments. The Bababudan group of schist belts rests on the gneisses above an unconformity marked by mature, oligomictic, quartz-pebble conglomerate (Viswanatha, 1968). This group consists mainly of basic metavolcanic rocks, argillaceousarenaceous detrital sediments (with current bedding and ripple marks), and banded magnetite quartzite. Magnetite quartzite is the dominant chemical sediment, and some layers are rich in Mn minerals. Fuchsite quartzite and meta-ultramafites are reported from a few locations. 2
3
2
2
The Chitradurga group consists mainly of metavolcanics, greywackes, and metashales. Pyritiferous chert is the most abundant chemical sediment, and ironstones also contain sulphides. Limestones and clays, commonly Mn-rich, are also present. Ramakrishnan et al. (1976) regard the Chitradurga group as younger than the Bababudan group, with the basal Chitradurga unconformity represented by the Talya conglomerate. Naqvi et al. (1978a, 19786), however, regard the Talya conglomerate as a greywacke-type, not representative of an unconformity; they believe that the Bababudan group is a marginal platform facies of the time-equivalent, deeper-water, greywacke facies of the Chitradurga group. PETROLOGY The chemical sediments of the various Dharwar greenschist belts can be grouped into several
TABLE
I
Correlation of schist belts of Dhar war Craton Kr i shnarajpet
Holenarsipur
Nuggihal1i
Ghatt i Hosahal1i
Javanahal1i
Bababudan
Ch i tradurga 2.6 - 2.2 b.y. Granitic activity. Chert, sericitic phyllite Metavolcan i cs Chlorite schist Greywacke conglomerates (K.M. Kere), greywacke Greywacke conglomerate (Aimangala) Arkose-gri t Greywacke conglomerate (Talya).
Kaldurga Conglomerates Quartzite (BMQ) Argi11i te schist Mafic flows Mafic flows with interlayered amphibolegarnet schist Chlorite schist Orthoquartz i te (fuchsitic in places) Basal conglomerate Unconformi ty Dykes Quartz veins Pegmatites ' Gran i tes Fuchsite quartzite Ultramafic schist Ortho-amphiboli tes Para-amphibolites i nterbedded wi th carbonate, calcsi1icates etc. Para-gnei sses. TONAL ITIC AND TRONDHJEMITIC ACTIVITY
Dun i te Amphi boli te Fuchs i te quartz i te Hornblende sch i s t Tremoli teact i noli teschist.
Ti tani ferous magnet i te Dun i te, Dun i te Fuchs i te quartz i te Fuchsite quartzite Fuchs i te quartz i te Bari te High A1, Mg, sediments Hornblende schist Hornblende schist Tremoli te-act i nol te- Tremolite-actinolite- Tremoli te-act inoli teschists, serpent inite, schi,st, serpentin te, schists, serpentinamphi boli te. ite, peridotite. peri dot i te.
60-70% of the area 95% of the area consi its is made up of basic/ ultrabasic rocks of basicultrabas ic rocks BASEMENT ????????
80% of the area i made up of ultrabas ic rocks
80% of the area is made up of ultrabasic rocks
90% of the area is made up of metasediments
More than 65% of the area is made up of volcanics
40% of the area is made up of sediments
248
S. M. NAQVI, P. K. GOVIL & J. J. W. ROGERS
categories. The high-Al, high-Mg rocks of the Holenarsipur and other older belts have been described by Naqvi (1978). This paper concentrates on the fuchsite quartzites, ironstones (magnetite quartzites), cherts, and carbonate rocks. Fuchsite Quartzites Fuchsite quartzites are more abundant in the Holenarsipur group than in younger belts. Fuchsite quartzite is the youngest member of the Javanahalli belt, occurs sparsely in the Bababudan group, and is present at only one location in the Chitradurga group. Petrographically, the fuchsite quartzites consist of euhedral to subhedral grains of quartz with fuchsite inclusions. The quartzites do not contain polycrystalline fragments. At Banavar, in the Holenarsipur group, Ramiengar et al. (1978) have reported chromite, apatite, zircon, arid tourmaline in the fuchsite quartzites, but the present writers have been unable to observe or separate the zircon grains from the rock. Ramiengar et al. (1978) have not yet described the morphology of the zircons observed by them. In general, impure layers contain garnet, kyanite, and muscovite. Apatite, tourmaline, garnet, kyanite and chromite occur as well-developed euhedral crystals and thus indicate their metamorphic origin. Garnet generally and kyanite locally contain inclusions of fuchsite. Quartz boundaries in the fuchsite quartzites are not sutured, as in clearly detrital quartzites. Barite is interbedded with the quartzites only in the Holenarsipur group (in the Ghatti Hosahalli schist belt). The presence of thin alternate layers of barite and fuchsite quartzite resting over the ultramafic volcanic rocks is very strong field evidence in favour of exhalative rather than detrital origin of the fuchsite quartzites of the Holenarsipur group. Ironstones (Magnetite Quartzites) Ironstones consist of magnetite and quartz ( ± garnet, grunerite, and actinolite). The Javanahalli and Bababudan groups contain banded magnetite quartzites, with minor grunerite in the Bababudan group and grunerite, actinolite and anthophyllite in the Javanahalli group. The Holenarsipur group contains minor amounts of rocks consisting of large garnets with interstitial quartz, magnetite, grunerite, and actinolite. Ore grades of magnetite quartzites occur in the Bababudan group. Iron-rich rocks in the Chitradurga group are more pyrite-bearing rather than magnetite-bearing (Naqvi etal., 1977).
Cherts Cherts are recognized only in the Chitradurga group and are extremely fine grained (not recrystallized). Different localities contain varieties such as hematitic (jasper), pyritiferous, chromiferous, and carbonaceous. Sulphide horizons contain pyrite, pyrrhotite, chalcopyrite, sphalerite, galena, cobaltite, and linnaeite (Naqvi et al., 1977). Some horizons contain siderite. Manganese (as psilomelane) is mined in some chert and underlying clay horizons (Mn is also present in some limestones). Carbonate Rocks Carbonates are generally absent from the Holenarsipur and Bababudan groups. In the Javanahalli group, carbonates range from pure calcite to pure dolomite; associated calc-silicates contain calcite, dolomite, quartz, amphibole, epidote, and/or diopside in various localities. The carbonates of the Chitradurga group are mostly high-calcium varieties (cement grade in some places). SEDIMENTARY COMPOSITION The measured compositions of Early Precambrian chemical sediments can be controlled by four processes: (1) exogenic (sedimentary) processes, including both chemical precipitation from ocean water and inclusion of clastic materials derived from existing crustal sources; (2) volcanogenic (exhalative) processes, in which constituents are added directly to the sediments by essentially hydrothermal processes, possibly from mantle sources; (3) metamorphism after original formation of the sediments; and (4) more recent weathering. The effects of these various processes cannot always be distinguished, particularly in preliminary studies such as the present one. Available chemical data are summarized in Table II and Figures 2-7. Table II contains average compositions of those suites for which the data are homogeneous enough to permit calculation of valid averages. Additional points are plotted on the figures either for single analyses of samples or for suites for which the data are too variable for averaging. On the figures, compositions of limestones and fuchsite quartzites are shown without distinction for geographic location or age because there seems to be no variation in composition of single samples with age or location. The Javanahalli cherts (non-pyritiferous) and Chitradurga cherts (pyrite-bearing) are plotted separately. Compositions of ironstones are variable with age and location and, therefore, are
C H E M I C A L SEDIMENTATION, INDIA
io-O r
1 0-1
•
o
Fig. 2.
Na vs K relationships in chemical sediments of the Dharwar greenschist belts. Symbols in this - and all other diagrams are as follows: O limestone (all belts) • fuchsite quartzite (all belts) barite x chert (Javanahalli belt) A pyritiferous chert (Chitradurga belt) ironstone • Holenarsipur belt Javanahalli belt © Bababudan belt ® Chitradurga belt
0-1
1-0 %
10-0
K
TABLE
249
plotted separately for different suites. Symbols for the various suites are shown in the caption for Figure 2. Figure 2 shows the relationships between Na and K (recalculated from their oxide percentages). The fuchsite quartzites have N a / K ratios less than 1.0, which is consistent with their content of mica. Limestones, however, have N a / K ratios greater than 1.0; this high ratio contrasts with the ratio of about 0.5 which appears to characterize more recent limestones (Heier & Billings, 1970a, 19706). The ironstones also have N a / K greater than 1.0, probably indicating a small amount of detrital component such as clays; comparison with the data of Bayley & James (1973) for U.S. iron formations indicates that the N a / K ratios of ironstones depend highly on mineralogy. Figure 3 shows the relationships between Rb and Sr. Rubidium contents are low in all rock types, similar to the low concentrations of K. The Sr concentrations in barite show expected high values, but the Sr contents of the limestones are unexpectedly low. The summary of Sr in limestones prepared by Veizer (1978) shows normal concentrations in the range of hundreds to several thousands of ppm Sr. The 8 7 Sr/ 8 6 Sr ratio of 0.706 available for one limestone from the Chitradurga belt is somewhat above the general curve of strontium-isotope evolution shown by Veizer & Compston (1976) in Figure 4. II
Average abundance of selected elements in chemical sediments of Dharwar greenschist belts FUCHSITE QUARTZITE
Javanaha111 belt
Number .. of • Samples ppm Co 16 Cr 35 Ni 11 Go • 5 in. 114 Pb 36 m 15 |r 9*7 % Mn Ti , P Na K
Ghitradurga belt
Javariahal l i belt
13 83 36
22
21
230
414 87
Pyritif- Pyritiferous erous (Chitra- (Chi tradurga durga belt; Javanabel t; m unmi ner- i ner- hal 1 i alized) belt alized)
8
10
6
158 38 9 54
182 24 19
210
48
43 152 . 144 10 90 13 5
10
10
10
0.220
83
10
0.270
0.270
0.161
0.186
0.35
0.085
0.06
0.608
0.71 0.546
0.108
0.248
0.072
0.620
0.708
4.82
0.415 0.113
Ca Mg
Ghatti Hosahal 1 i belt
2625 45 25
19 269 99
All values are in parts per mill ion or percent by weight. 1 From Naqvi et al. (1977) -
0.731 O.'l 0.194 0.04
17 225
72 .
10
6
2
l4o 125
205 4
35 11
0.175
1 pur be 11
403 221 129 40 66
10
30 <5
6700
<10
0.106
0.256 0.068
0.09 0.197 1.49 0.675 0.162
Ghatt i Hosahal 1 i belt
0.042 0.581 0.169
0.025 0.370 0.1 0.424 0.548
Javanahal 1 i . belt
Chitradurga belt
42
41 185 167 33 92 18 <5
278 86 26
102 19 <5 <10 0.242 0.234 0.026 1 -135 0.1 1.493 0.414
<10
0.240 0.182 0.062 1.087 0.1 .0.114 0.011
250
S. M. NAQVI, P. K. GOVIL & J. J. W- ROGERS ioV
cP o
10
10
10
S r (ppm) -
10
Fig. 3. Rb vs Sr relationships. INDEX OF CRUSTAL EVOLUTION
0-715
0 0
0-7 10 .rCHITRADURGA ^LIMESTONES
en cn
CD oo oo
0-705 ^Ppei
*
mm
SP°ntl
e
0-700
3 b.y.
Q T| —M |'-ols-cl/3 P-
IB.S.I G.B Mc A. I Mt.B sn. \aci. 8G «T. aB.G.xaB.
is.R.a \ B#L.. I Sw. o.ash.
Fig. 4. Distribution of *7Sf/86Sr carbonates (after Veizer & Compston, 1976). Position of Chitradurga limestone is shown. Explanations: Q—Quaternary; T & M—Tertiary and Mesozoic; P-D-PermianDevonian; S-C-Silurian-Cambrian; A & P-Adelaidean and Pertatataka Fm; B.S. & B.N. & G.—Bitter Springs Fm. and Nonesuch Shale and Bangemall Group; G.B. & CL & B-Grenville Belt and Clarke Series and Belt Series; McA & G.-McArthur and Gunflint Fms; Mt.B & T—Mt Bruce Supergroup and Transvaal sequence; S.R. & O. & Sh—Steep Rock Lake and Ontario greywackes and Shamvian System; B.L.—Bulawayan Lm; Sw—Swaziland sequence. r a t i o
i n
Relationships between Cr, Co, and Ni are shown in Figures 5 to 7. The constant-ratio line shown on each diagram is the presumed ratio for the Archaean upper mantle determined by Sun & Nesbitt (1977). Cobalt concentrations are slightly abovq these lines (high values), but the absolute
Co concentrations are very low and subject to considerable error. Observed Cr/Ni ratios, however, are commonly close to this mantle ratio for all rocks except the barite rocks. Considering the diversity expected for the wide variety of rock types investigated, the relationships are surpris-
CHEMICAL SEDIMENTATION, INDIA
O^
A•
/ +uA-+ •
/
•
X
Fig. 5. Co vs Ni relationships. Constant-ratio line is presumed Archaean upper mantle ratio (from Sun & Nesbitt, 1977).
ingly close between observed Co/Ni, Co/Cr, and Cr/Ni ratios and the presumed Archaean uppermantle ratios. The fuchsite quartzites, as expected, contain high Cr concentrations. These rocks generally have high contents of Al, Mg, .Cr, Ni, Co, Zn, Pb and Cu relative to detrital quartzites studied by Hussain (1980). Lowe & Knauth (1977) found volcanic zircons in ''green cherts" of similar composition, thus presumably indicating a nondetrital source. Interbedded barite is regarded as also of chemical origin, in distinction to the proposed formation of detrital barite by Heinrichs & Reimer (1977). Preliminary data of oxygenisotope ratios 6( O) on two samples of fuchsite quartzite (from Holenarsipur schist belt) which contain fairly large numbers of quartz veins are + 11.6 and +12.5. These values when compared with those of quartz of granodiorites ( + 7.0 to + 10.00, Taylor, 1978) and Early Precambrian cherts from a Rhodesian greenstone belt (+11.0 to +18.7, Perry & Tan, 1972) and Greenland (+12.9 to +20.4, Perry et al., 1978) show that
251
these quartzites were not derived from the quartz of plutonic granitic rocks. It seems that the minerals like fuchsite, chromite and kyanite and mixing of the vein quartz have lowered the oxygen isotopic ratio as compared to pure Archaean cherts of Greenland or Rhodesia. Any modification of the chert tends to lower the 6(180) value (Perry, 1967). Abundances of Ni and Cr are different in the ironstones of different belts (Figure 7). In general, the Ni and Cr contents of the ironstones are in the range of 100 to 300ppm, similar to the concentrations in the mafic rocks of the belts (Naqvi & Hussain, 1973). The highest concentrations of Ni and Cr are in the pyritiferous rocks of the Chitradurga belt. The analyzed samples from the belt were chosen to avoid major sulphide layers, which contain minerals such as pyrite, pyrrhotite, chalcopyrite, galena, cobaltite, and linnaeite; some areas contain Cu in ore grade, with abundances of as much as 18 percent Cu and 2200 ppm Co (Naqvi et al., 1977). Thus, it is likely that the Ni and possibly the Cr concentrations in the analyzed Chitradurga rocks are the result of sulphide mineralization rather than primary constituents. io r 3
ls
10 rr
J io' 10 _ Cr (ppm)—Fig. 7. Cr vs Ni relationships. Constant-ratio line is presumed Archaean upper mantle ratio (from Sun & Nesbitt, 1977).
I
io
1
SUMMARY AND CONCLUSIONS The abundance of elements in the sediments of the Dharwar greenschist belts can be summarized as follows: (1) concentrations of K and Rb are lower than would be expected in more modern C r (ppm) sediments except where some specific process Fig. 6. Co vs Cr relationships. Constant-ratio line is (e.g. the formation of fuchsite mica) caused local presumed Archaean upper mantle ratio (from concentrations; (2) abundances of Cr and Ni are high in most sediments, with concentrations Sun & Nesbitt, 1977).
252
S. M. NAQVI, P. K. GOVIL & J. J. W. ROGERS
similar to those in associated mafic rocks and an abundance ratio similar to that proposed for the Archaean upper mantle; (3) concentrations of Sr are low, particularly in limestones that would be expected to contain concentrations at least one order of magnitude higher. Because of the variety of possible processes controlling element abundances (discussed earlier), a complete explanation of these compositional features must be very tentative. As the unusual composition of the rocks appears to be associated with the age of the sediments, it is likely that neither metamorphism nor later weathering have greatly influenced the compositions. Discrimination, then, must be made between exogenic (sedimentary) and exhalative (volcanogenic) processes. The importance of volcanic, exhalative, processes in the Archaean has been stressed by a number of writers. Goodwin (1973) ascribed the Michipicoten iron ores of the Canadian Shield to volcanogenic activity because of the relationships between iron ore and underlying volcanic rock. Lowe & Knauth (1977, 1978) have shown that volcanic activity (accretionary lapilli) can cause the formation of rocks that have been regarded as typical sediments (in that case, ooids) by other workers. The possible importance of exhalative processes in the Dharwar chemical sediments is indicated by the three features of their composition summarized above. The low K and Rb in quartzites and low Sr in limestones are not properties of more " n o r m a l " sedimentary rocks. These low contents are present even in rocks formed after the development of a sialic crust, which apparently took place about 3300m.y. ago. Before that, the anorthosite-norite-troctolite rich mafic crust previously proposed by Naqvi (1978) would not have been a significant source of lithophile elements, but such elements should have been present in sedimentary source rocks for the Javanahalli and younger schist belts (see correlations in Table I). The abundances of Cr, Co, and Ni are not normally regarded as affected by volcanogenic processes (e.g., summary of Hutchinson, 1973). The similarity of Cr-Co-Ni abundance ratios in the sediments to those in presumed Archaean upper mantle, however, suggests direct mantle derivation. Furthermore, sedimentary processes would be expected to separate Cr from Co and Ni. A mantle that was capable of producing hightemperature, komatiitic lavas may also have been capable of producing exhalative fluids rich in refractory elements, and that process seems to be the most likely explanation for the observed compositions of the Dharwar sediments.
The varieties of chemical sedimentation described above can be correlated with the types of volcanic activity that characterize the different schist belts. Naqvi (1979) has shown a progressive change from ultramafic (komatiitic) volcanism in the Holenarsipur and other older belts to tholeiitic and calcalkaline volcanism in the younger belts. Siliceous volcanic rocks are also progressively more abundant in younger belts. A similar change occurs in the dominant mineralogy of the ore metals. Sulphides predominate in the younger belts (especially Chitradurga), and oxides are more abundant in the older belts. Although individual samples of chemical sediments in the Holenarsipur group contain about the same amount of Ni and Cr as younger sediments, the large quantity of Holenarsipur chemical sediments causes the group as a whole to be rich in Ni and Cr. The Cr ( - 1000-3000ppm), Ni (-1000 ppm) and Co ( - 5 0 ppm) content of the fuchsite quartzite resting directly on ultramafic volcanic rocks in the Ghatti Hosahalli and Banavara belts (Narayana & Naqvi, 1980) is difficult to explain by detrital processes. This Ni and Cr abundance is probably related to komatiitic volcanism and associated exhalation products. Similarly, chalcophile elements occur in greatest quantity in the Chitradurga belt because of the large volume of pyritic rocks. These preliminary data for the chemical sediments of the Dharwar schist belts are consistent with earlier proposals concerning the evolution of the craton (Naqvi, 1978, 1979). Before about 3300m.y. volcanism and sedimentation took place in an essentially oceanic environment. About 3300m.y. ago, development of sialic gneisses and trondhjemites-tonalites formed a thickened sialic crust that provided at least a partial source for younger sediments. Ultimately, the Bababudan group was deposited in a platform environment, and water in the adjacent ocean basins was deep enough for the sulphiderich Chitradurga group to be deposited. ACKNOWLEDGMENTS We are grateful to Dr S. Balakrishna, Director of National Geophysical Research Institute, Hyderabad, for his support for this work and permission for its publication. Dr B. P. Radhakrishna's and Dr Hari Narain's keen interest and guidance in the project are also acknowledged. Members of the Geochemistry Group of N.G.R.I., helped in various ways in the preparation of this paper. Dr P. A . Mueller, University of Florida, provided the Sr isotope analysis. This work has been conducted under a Joint Indo-US project on the Precambrian of South India funded by NSF Grant INT78-17128.
CHEMICAL SEDIMENTATION, INDIA
253
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BAYLEY, R . W . , & JAMES, H . 68, 9 3 4 - 9 5 9 . BECKINSALE, R . P . ,
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Precambrian Iron and Manganese deposits, 2 3 - 3 4 . Proceedings of the Kiev Symposium 2 0 - 2 5 August, 1970, UNESCO. HEIER, K . S . , &
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, 1978: The oldest marine carbonate ooids re. interpreted as volcanic accretionary lapilli, Onverwacht gfoup, South Africa. J. sediment. Petrol., 48, 709-722. M., 1 9 7 8 : Geochemistry of Archaean metasediments; evidence for prominent anorthositenorite-troctolite (ANT) in the Archaean basaltic primordial crust; in Windley, B. F., & Naqvi, S. M. (Eds) Archaean Geochemistry, 343-360. Elsevier, Amsterdam.
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,.,. 1979: Distribution of elements in the crust and mantle during the Archaean—evidence from the Indian shield. Chem. Geol., 24, 1-23. NAQVI, S . M . , & HUSSAIN, S . M . ,
1973: R e l a t i o n
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tween trace and major element composition of the Chitaldrug metabasalts, Mysore, India, and the Archaean mantle. Chem. Geol., 11, 1 7 - 3 0 . NAQVI, S . M . , HANUMANTHA R A O , T . , NATRAJAN,
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SAIN, S. M., 1977: Mineralogy, geochemistry, and genesis of the massive base metal sulphide deposits of Chitradurga (Ingaldhal), Karnataka, India. Precamb. Res., 4, 361-386. NAQVI, S . M . ,
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NARAIN,
1978a: The primitive crust—evidence from the Indian shield. Precamb. Res., 6, 323-345.
NAQVI, S . M . , DIVAKARA RAO, V . , RAMA RAO, P . , H U S SAIN, S . M . ,
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19786: Stratigraphy of Dharwar formations— a tentative suggestion based on field data and conceptual models/Workshop on ''Standardization of stratigraphic nomenclature of the Precambrian formations of South India", Geol. Soc. Ind, Bangalore. NAQVI,
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A H M A D , S . M . , & UDAY R A J , B . ,
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and geochemistry of paragneisses from the Javanahalli schist belt, Karnataka, India. J. geol. Soc. Ind., 21, 577-592. NARAYANA, B . L . , & NAQVI, S . M . , 1 9 8 0 : Geochemistry of spinifex-textured peridotitic komatiites from Ghatti Hosahalli, Karnataka, India. J. geol. Soc. Ind,
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PERRY, JR., E. C., 1967: The oxygen isotope chemistry of ancient cherts. Earth planet. Sci. Lett., 3, 6 2 - 6 6 . P E R R Y , JR., E. C., & T A N , F. C., 1972: Significance of oxygen and carbon isotope variations in early Precambrian cherts and carbonate rocks of southern Africa. Bull. geol. Soc. Am., 83, 647-664. PERRY J R . , E . C . , AHMAD* S . N . , & SWULIUS, T .
M.,
1978: The oxygen isotope composition of 3800 m.y. old metamorphosed chert and iron formation from Isukasia, West Greenland. J. Geol., 86, 223-239. RAMAKRISHNAN, M . ,
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N., &
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NATH, J., 1976: Basement-cover relationships of Peninsular gneiss with high grade schists and greenstone belts of southern Karnataka. J. geol. Soc. Ind., 17, 97-111. RAMIENGAR,
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S.,
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M . N . , CHAYAPATHI, N . , & RAMAKRISHNAN,
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1978: Banded chromite-fuchsite quartzite in the older supracrustal sequence of Karnataka. J. geol. Soc. Ind., 19, 577-582. SWAMI N A T H , J . , RAMAKRISHNAN, M . , & VISWANATHA,
M. N., 1974: The cratonic greenstone belts of southern Karnataka and their possible relation to the charnockite mobile belt [Abstract]; in Tectonics
and Metallogeny of Southeast and Far East Asia, 38-39. Geol. Surv. Ind., Calcutta. , 1976: Dharwar stratigraphic model and Karnataka craton evolution. Rec. geol. Surv. Ind., 107, 149-175. S U N , S . - S . , & NESBITT, R . W . , 1 9 7 7 : Chemical heterogeneity of the Archaean mantle—composition of the earth and mantle evolution. Earth planet. Sci.
Lett., 35, 4 2 9 - 4 4 8 . TAYLOR, H. P., 1978: Oxygen and hydrogen isotope studies of plutonic granitic rocks. Earth planet. Sci. Lett., 38, 177-210. VEIZER, J., 1978: Strontium abundance in common sediments and sedimentary rock types; in Wedepohl, K. H. (Ed.) Handbook of Geochemistry, Chap. 38K, Springer-Verlag, Berlin. VEIZER, J . , & COMPSTON, W . , 1976 : 8 7 S r / 8 6 S r c o m p o s i -
tion in Precambrian carbonates as an index of crustal evolution. Geochim. cosmochim. Acta, 40, 905-914. VISWANATHA, M . N . , 1968: Rep. geol. Surv. Ind. [unpublished].
EVIDENCE FROM RARE-EARTH ELEMENTS FOR THE CHEMICAL COMPOSITION OF THE ARCHAEAN CRUST S. R. Taylor & S. M. McLennan Research School of Earth Sciences, Australian National University, Canberra, ACT. 2500, Australia
ABSTRACT The present-day composition of the upper continental crust is reasonably well constrained by large-scale sampling programmes and studies of rare-earth elements (REE) in sedimentary rocks. Its composition approximates to granodiorite and has remained unchanged during postArchaean time. REE patterns in Archaean sedimentary rocks differ from their post-Archaean counterparts having no Eu-depletion and generally lower EREE and La/Yb. The average pattern resembles that of modern calc-alkaline andesites, and indicates that the exposed Archaean crust was different to the present-day exposed crust. Archaean sedimentary REE patterns are best modelled by a two-component system involving Archaean mafic volcanics and felsic igneous rocks (tonalites-trondhjemites, felsic volcanics). Using such a model, estimates of the exposed Archaean crust can be made, and indicate a composition considerably more mafic than the post-Archaean upper crust. The Archaean upper crust is similar in composition to the present-day total crust .except that it may be enriched in Ni and Cr.
INTRODUCTION There has recently been an active interest in the nature of the Earth's early crust (Bridgwater & Fyfe, 1974; Hargraves, 1976; Taylor, 1977, 1979; Collerson & Fryer, 1978; Fyfe, 1978; O'Nions & Pankhurst, 1978; Young, 1978; Taylor & McLennan, 1981a, b for example). To date however, there has been no attempt at actually estimating the composition of that ancient entity (with the exception of the rare-earth elements—see below). This is remarkable when one considers the constraints such an estimate would impose on all models of crustal evolution. In this paper, we estimate the composition of the Archaean exposed crust, compare it to the composition of the present-day continental crust, and comment on some implications for the evolution of the Earth's crust. THE MODERN CONTINENTAL CRUST The present-day continental crust comprises about 0.43% of the mantle-crust system, by mass. Despite its small mass, more than 10% of K, Ba, Cs, Sr, Th and U are concentrated in the continental crust when compared with the total mantle-crust system. The present composition of the upper continental crust is reasonably well constrained from large-scale sampling programmes (Shaw et al., 1967, 1976; Fahrig & Eade, 1968; Eade & Fahrig, 1971, 1973) and
Spec. Pubis geol. Soc. Aust., 7 (1981)
studies of the sedimentary rocks (Goldschmidt, 1954; Taylor, 1964, 1977, 1979; Condie, 1967; Young, 1969; McLennan et at., 1979; Taylor & McLennan, 1981a, b). The sedimentary approach has recently been most concerned with the REE distribution in sedimentary rocks as an index of upper crustal composition (Wildeman & Haskin, 1973; Jakes & Taylor, 1974; Nance & Taylor, 1976, 1977; Taylor, 1977, 1979; McLennan et al., 1979; McLennan & Taylor, 1980; Taylor & McLennan, 1981a, b). This approaches particularly effective since the REE do not concentrate in sea water, have very short residence times, and are not strongly affected by most sedimentary or meta-. morphic processes. Post-Archaean fine-grained sedimentary rocks have REE patterns typical of granodiorites (La/Yb ~9; Eu/Eu* = 0.64 ± 0.05). Such composition agrees perfectly with the sampling programs which indicate a majorelement (and some minor-element) composition approximated by granodiorite. In fact, if one were to estimate upper crustal composition by simply choosing the igneous rock most similar to the sedimentary REE pattern, the estimate would be very close to that obtained by large-scale sampling. This observation becomes important when we look at the Archaean crust. The composition of the upper continental crust can be measured, but the lower-crust composition is model-dependent. A viable model for the cornN
S. R. TAYLOR & S. M. McCLENNAN
256
position of the total crust is that it approximates the average composition of calc-alkaline, islandarc-type volcanics and associated igneous rocks (i.e. andesite; Taylor, 1967, 1977; Taylor & McLennan, 1979). It should be noted that this model for crustal growth by island-arc volcanism can not be confidently extrapolated back beyond the Upper Proterozoic. The compositions thus refer to late additions to the crust. Possibly much of the bulk of the continental crust was derived from crustal growth in late Archaean time (30002500Ma). Whether the mechanism for this sharp episodic increase in continental volume was due to continuation of Archaean bimodal basalttonalite igneous activity, to conventional islandarc volcanism, or to a third mechanism is unclear at present. Such questions are important for estimating the crustal abundances of Ni and Cr especially. If we assume that the upper crust represents about 1/3 of the total crust, the model predicts a lower crust of more basic composition. Is there any evidence which supports this prediction? Obtaining an estimate of average lower-crustal
La
Ce
Pr
Nd
Sm
abundances is intrinsically more difficult than upper-crustal abundances since the lower crust is probably lithologically complex and there are no natural sampling devices such as sedimentation. However, the few studies made seem to agree essentially with the "andesite" model as described most recently by Taylor & McLennan (1981 b) (Rogers, 1977; Taylor & McLennan, 1979, 1981 a; Mueckeetal., 1979). The generation of the large Eu depletion (Eu/Eu* —0.64) in the upper crust is almost certainly an intra-crustal process. No common mantle-derived rock can be characterized by an Eu anomaly of any sort. Depletion of Eu is probably due to plagioclase remaining as a residual phase during partial melting. Since plagioclase is not stable below about 40 km (10 kb), such Eu depletion must be caused by relatively shallow (intra-crustal) melting events. ARCHAEAN SEDIMENTARY REE PATTERNS Archaean sedimentary rocks have REE patterns which differ substantially from post4
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Fig. 1. Rare-earth element (REE) patterns, normalised to chondritic meteorites (C-l, volatile free) of Archaean and post-Archaean sedimentary rocks from Australia. Also shown are estimates of the average Archaean Australian sedimentary rock (AAS) and average post-Archaean Australian sedimentary rock (PAAS). Note the differences in the average patterns and small overlap for the fields. AAS resembles typical andesites and is thought to represent exposed Archaean crust. PAAS is thought to be parallel to the pastArchaean exposed crust, but about 20% higher in EREE on account of low REE-bearing sedimentary rocks, such as carbonates and evaporites, which are abundant during this time. The sedimentary REE patterns indicate a granodioritic upper crust throughout the post-Archaean (reproduced with permission of Elsevier, from Taylor & McLennan, 1981a).
C H E M I C A L COMPOSITION OF A R C H A E A N
Archaean sedimentary rocks. Figure 1 compares REE patterns, normalized to chondrites, for Archaean and post-Archaean sedimentary rocks from Australia. The average Archaean pattern has lower (La/Yb) , lower EREE, and no significant Eu-anomaly. Taylor & McLennan (1981a, b) have discussed at length the origin of these REE ! patterns and conclude that they probably represent a reasonable sampling of the exposed crust during the Archaean. The absolute REE abundances are probably very similar to the Archaean upper crustal abundances since sedimentary rocks of crustal origin with low REE abundances (e.g. evaporites, carbonates) were not abundant at that time. Solely on the basis of REE, Taylor & McLennan (198la) modelled the Archaean upper crust using the two-component system of Archaean mafic volcanics and trondhjemite-tonalite granitic rocks. Other lithologies such as andesites, felsic volcanics, ultramafics and K-rich granitic rocks of intracrustal origin were considered of lesser importance. Recent work in the Yellowknife district of northern Canada (Jenner 0 al., in press) indicates that, locally, felsic volcanics can play an important role in the origin of Archaean sedimentary REE patterns. Taylor & McLennnan (1981a) determined that a model involving 30-50% Archaean trondhjemite and fonalite granitic rocks mixed with typical Archaean tholeiitic basalts could explain the REE patterns found in Archaean sedimentary rocks.
257
CRUST TABLE I
Crustal abundances
3
k
Archaean . upper
2 PostArchaean total
57-.4 0,.9 1 5 . .6 9..5 5..2 7.• 3 3. . 1 0-• 9
58.0 0,.8 18,.0 7.• 5 3.• 5 7.• 5 3..5 1 .• 5 '
66.0 0.6 16.0 4.5 2.3 3.5 3-8 3.3
65.1 0.50 16.0 4.3 2.3 3-4 4.1 2.70
Rb Ba Sr Pb
25 240 300 7
42 350 400 7
110 700 350 15
790 410 17
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
1 2 . ,.6 ft 26.,8 3. , 1 13-,0 2. 78 0..90 2.85 0.,48 2. 93 0. 63 1 . 81 0. 26 1 . 79 0 . 28
19 38 4. 3 16 3. ,8 1 .. 1 3. .6 0.,64 3. 0.,82 2. 3 0. 32 2. 2 0. 30
30 64 7.1 26 4.5 0.88 3.8 0.64 3-5 0.80 2.3 0.33 2.2 0.32
Y Th U Zr Hf Nb"
15 2. 9 0. 75 100 3 5
22 4. 8 1 . 25 100 3. 0 , 11
22 10.5 2.5 240 5.8 25
Mn Cr V Sc Ni Co Cu Zn
1300 140 150 25 90 30 80 100
1 100 55 1.75 30 30 25 60
600 35' " 60 10 20 10 25 52
N
Si 02 tlo2 AI2O3 FeO-p MgO CaO Na20 K2O
PostArchaean upper
LateArchaean upper
PPm
1
.
-
-
-
• •
-
-
9.7 1.2 -
620 88 -
26 -
33 55
THE ARCHAEAN-PROTEROZOIC TRANSITION Detailed studies of Lower Proterozoic sedimentary sequences in North America and Australia (McLennan et al., 1979; McLennan & Taylor, 1980) strongly indicate a rapid evolution of sedimentary "REE patterns associated with the Archaean-Proterozoic boundary. It has been suggested that the large volumes of K-rich granitic rock intruded into the upper crust near the end of the Archaean were responsible for the change in sedimentary REE patterns. The rapidly changing REE patterns documented in Lower Proterozoic sequences represent increasing unroofing of these rocks at the end of the Archaean (McLennan et al., 1979; McLennan & Taylor, 1980). age REE patterns, as sampled by the sedimentary rocks, thus provide the key which enables us to COMPOSITION OF THE EXPOSED identify the most common, or average, igneous ARCHAEAN CRUST rock making up the crust. Table I lists our The determination of the composition of the estimate of the Archaean upper crust. Major Archaean upper crust is constrained by the REE elements, Ba, Co, Cu and Zn were determined patterns in sedimentary rocks which provide the from the average exposed Archaean volcanic rock overall crustal average. The crust, however, is from the Canadian Shield (Goodwin, 1977). Sr, primarily formed from igneous rocks. The aver- Zr, V, Ni, Cr, Sc and Pb were derived from 9
1)
Estimate of the early Archaean ( p r e - 3 - 0 x 1 0 y r ) exposed c r u s t . M a j o r e l e m e n t s , B a , C o , Cu and Zn derived from the average Archaean volcanic rock from the Superior P r o v i n c e , Canada (Goodwin, 1977). REE from A v e r a g e Archaean Sedimentary R o c k , AAS (Taylor & M c L e n n a n , 1981a). C r , S r , Z r , V , N i , Sc and Pb derived from tholeiite/tonalitetrondhjemi'te mix (data from Sun & N e s b i t t , 1977., 1978; G l i k s o n , 1979). Rb from K/Rb = 3 0 0 , U from K/U = .10*., Th from Th/U = 3 - 8 .
2)
Estimate of the bulk composition of the p r e s e n t - d a y conticrust pred i cted by the a n d e s i t e m o d e l . From Taylor (1977, 1979) with m i n o r modifications as described by Taylor & McLennan (1981a, b). Ni data revi sed from 20 to 30 ppm.
3)
Estimate of the bulk composition of the present-day upper continental c r u s t . From T a y l o r (1977, 1979) with modifications as discussed by T a y l o r & McLennan (19812?). The most significant m o d i f i c a t i o n is the lowering of REE. abundances by 20% (see text for d i s c u s s i o n ) . Ni value from Shaw et al. (1976).
4)
Estimate of the composition o f the exposed Archaean crust in the Canadian Shield.- From Fahrig 6 Eade (1968); Eade & Fahrig (1973).
258
S. R. TAYLOR & S. M. McCLENNAN
1-0
10
pp m
100
0.1
POST-ARCHAEAN UPPER C R U S T
1.0
o/
10
/ o
Hg. 2, Comparison of the compositions of the Archaean and post-Archaean exposed crusts. The Archaean exposed crust is enriched in ferromagnesian elements and depleted in the large-ion lithophile (LIL) elements, when compared to the post-Archaean upper continental crust. This indicates a significantly more mafic exposed crust during the Archaean. Note the particular enrichment of Ni and Cr in the Archaean.
mixing average Archaean tholeiite and average tonalite-trondhjemite granitic rocks in the proportion of 1:1 (Sun & Nesbitt, 1977, 1978; Glikson, 1979). The REE were taken from the average Archaean sedimentary rock of Australia (AAS; Taylor & McLennan, 1981a). The values for Rb, U and Th were derived from the canonical ratios of K/Rb = 300; K/U = 10 and Th/U = 3.8. Figure 2 (also see Table I) compares the Archaean upper crust to the post-Archaean upper crust. The Archaean is considerably more mafic, being enriched in the ferromagnesian elements and depleted in REE and LIL elements. Ni and 4
Cr are enriched to a greater extent than are the other ferromagnesian elements. This trend is probably real for two reasons, Archaean basalts are slightly, but significantly (20%) enriched in Ni and Cr compared to MORB (Sun, pers. comm.), so that crustal averages based on tholeiite-tonalite models will be high for these elements. The post-Archaean upper crustal averages are based on andesite models for crustal growth, and so may be biased toward low Ni and Cr values. Previous estimates of the exposed Archaean crust (Fahrig & Eade, 1968; Eade & Fahrig, 1973) appear to have included a large volume of late
CHEMICAL COMPOSITION OF ARCHAEAN CRUST Archaean K-rich granitic rocks. This results in a composition which is only marginally different to the post-Archaean upper crustal composition (Table I). These estimates can only be regarded as representative of the latest Archaean. SPECULATIONS O N T H E A R C H A E A N LOWER CRUST It is intrinsically very difficult to calculate the composition of the Archaean lower crust since we are uncertain about many important parameters such as crustal thickness, precise mechanism of crust formation and degree of intracrustal melting. On the other hand, knowledge of the composition of the exposed Archaean crust does
Fig. 3.
259
provide an important boundary constraint on composition and allows speculation, in a qualitative manner, on the lower Archaean crust. Figure 3 compares the composition of the Archaean upper crust to the composition of the present-day total crust as calculated by Taylor & McLennan (19816). The comparison is very close, except that Ni and Cr are higher in the Archaean crust. This is attributed to a change in crustal growth from basalt-tonalite igneous activity to andesites in post-Archaean times (Taylor & McLennan, 1981a, b). On close inspection, the Archaean upper crust appears to be slightly enriched in most ferromagnesian and slightly depleted in most LIL elements, when compared to the modern total
Comparison of the compositions of the Archaean exposed crust to the post-Archaean total continental crust (see Table I). The Archaean upper crust compares favourably to the present day bulk continental crust, although it is slightly enriched in most ferromagnesian elements and slightly depleted in most LIL elements. Note the significant enrichment of Ni and Cr in the exposed Archaean crust.
S. R. TAYLOR & S. M. McCLENNAN 260 crust, indicating a slightly more mafic composi- marked the end of the Archaean in mosi places (Taylor & McLennan, 1981a, b). tion, consistent with the above models. The weight of the evidence would seem to indiThe important influence of the mantle in the cate that the Archaean crust was geochemically derivation of many early Archaean igneous rocks has been noted by several workers (O'Nions & less differentiated than the present continental Pankhurst, 1978, for example). Movement of crust. As we have noted, the Archaean upper elements from the lower crust by metamorphic crust was slightly more mafic than the present processes (Heier, 1973, 1978; Collerson & Fryer, total crust. It is reasonable to suggest that the 1978) is difficult to evaluate although it probably Archaean lower crust was locally more mafic occurred to some extent. Large-scale geochemical than the Archaean upper crust in regions where differentiation through intracrustal melting is intra-crustal melts had been extracted. also documented throughout the Archaean (for example Taylor & Hallberg, 1977; Fryer & A C K N O W L E D G M E N T S Jenner, 1978) but appears to be a relatively minor We thank C. Neagle and G. Stewart for assisphenomenon before the granite production which tance in the preparation of the manuscript. REFERENCES , 1978: The distribution and redistribution of heat-producing elements in the continents. Phil crust: fact, fiction, fantasy. Geosci. Can., 1, 7-11. Trans. R. Soc. Lond., Ser. A, 288, 393-400. COLLERSON, K. D . , & FRYER, B. J . , 1978: The role of fluids in the formation and subsequent develop- JAKES, P., & TAYLOR, S. R . , 1974: Excess europium content in Precambrian sedimentary rocks and conment of early continental crust. Contrib. Mineral. tinental evolution. Geochim. cosmochim. Acta, 38 Petrol., 67, 151-167. 739-745. CONDIE, K. D., 1967: Geochemistry of Early Precambrian greywackes from Wyoming. Geochim. JENNER, G . A . , FRYER, B. J . , & MCLENNAN, S. M., in press: Geochemistry of the Archaean Yellowknife cosmochim. Acta, 31, 2135-2149. Supergroup. Geochim. cosmochim. Acta. EADE, K. E . , & FAHRIG, W . F . , 1971: Geochemical MCLENNAN, S. M . , FRYER, B. J . , & YOUNG, G. M., evolutionary trends of continental plates—a pre1979: Rare earth elements in Huronian (lower Proliminary study of the Canadian Shield. Bull. geol. terozoic) sedimentary rocks: composition and Surv. Can., 179. evolution of the post-Kenoran upper crust. Geo, 1973: Regional lithological and temporal varichim. cosmochim. Acta, 43, 375-388. ation in the abundances of some trace elements in MCLENNAN, S. M . , & TAYLOR, S. R., 1980: Rare earth the Canadian Shield. Pap. geol. Surv. Can., 72-46. elements in fine-grained sedimentary rocks and FAHRIG, W . F., & EADE, K. E . , 1968: The chemical ores from the Pine Creek Geosyncline; in evolution of the Canadian Shield. Can. J. Earth Ferguson, J. & Goleby, A. B. (Eds) Uranium in the Sci., 5, 1247-1252. Pine Creek Geosyncline, 175-190. Internat. Atomic Energy Agency, Vienna. FRYER, B. J . , & JENNER, G. A., 1978: Geochemistry and origin of the Archaean Prince Albert Group MUECKE, G . K . , PRIDE, C . , & SARKAR, P . , 1979: Rarevolcanics, western Melville Peninsula, Northwest earth element geochemistry of regional metaTerritories, Canada. Geochim. cosmochim. Acta, morphic rocks. Phys. Chem. Earth, / / , 449-464. 42, 1645-1654. NANCE, W . B., & TAYLOR, S. R., 1976: Rare earth eleFYFE, W . S., 1978: The evolution of the earth's crust: ment patterns and crustal evolution—I. Australian modern plate tectonics to ancient hot spot tecpost-Archaean sedimentary rocks. Geochim. tonics? Chem. Geol., 23, 89-114. cosmdchim. Acta, 40, 1539-1551. GLIKSON, A. Y., 1979: Early Precambrian tonalite, 1977: Rare earth element patterns and crustal trondhjemite sialic nuclei. Earth-Sci. Rev 15 evolution—II. Archaean sedimentary rocks from 1-73. Kalgoorlie, Australia. Geochim: cosmochim. Acta, 41, 225-231. GOLDSCHMIDT, W. M., 1954: Geochemistry. Oxford O'NIONS, R. K . , & PANKHURST, R. J . , 1978: Early Univ. Press, Oxford. Archaean rocks and geochemical evolution of the GOODWIN, A. M., 1977: Archaean volcanism in Earth's crust. Earth planet. Sci. Lett., 38,-211-236. Superior Province, Canadian Shield. Spec. Pap. ROGERS, N . W . , 1977: Granulite xenoliths from Lesotho geol. Ass. Can., 16, 205-241. kimberlites, and the lower continental crust. HARGRAVES, R. B., 1976: Precambrian geologic history Nature, Lond., 270, 681-684. Science, N. Y., 193, 363-371. SHAW, D . M . , DOSTAL, J . , & KEAYS, R. R . , 1976: AddiHEIER, K. S., 1973: A model for the composition of the tional estimates of continental surface Precambrian deep continental crust. Fortschr. Miner 50 shield composition in Canada. Geochim. cosmo174-187. chim. Acta, 40, 73-83. BRIDGWATER, D . , & FYFE, W . S., 1974: The pre-3b.y.
CHEMICAL COMPOSITION OF ARCHAEAN CRUST
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SHAW, D. M . , REILLY, G . A . , MUYSSON, J . R . , PATTEN-
TAYLOR, S. R . , & HALLBERG, J . A . , 1977: R a r e - e a r t h
DEN, G. E., & CAMPBELL, F. E., 1967: An estimate
elements in the Marda calc-alkaline suite: an Archaean geochemical analogue of Andean type volcanism. Geochim. cosmochim. Acta, 41,
of the chemical composition of the Canadian Precambrian Shield. Can. J. Earth 5c/., 4, 829-853. SUN, S.-S., & NESBITT, R. W., 1977: Chemical heterogeneity of the Archaean mantle, composition of the Earth and mantle evolution. Earth planet. Sci. Lett., 35, 429-448. - . 1978: Pedogenesis of Archaean ultrabasic and basic volcanics: evidence from rare earth elements. Contrib. Mineral. Petrol65, 301-325. TAYLOR, S. R., 1964: Trace element abundances and the chondritic earth model. Geochim. cosmochim. Acta, 28, 1989-1998. ,1967: The origin and-growth of continents. Tectonophys., 4, 17-34. , 1977: Island arc models and the composition of the continental crust. Am. geophys. Un., Ewing Series, 1, 325-335. ^
, 1979: Chemical composition and evolution of the continental crust: the rare earth element evidence; in McElhinny, M. W. (Ed.) The Earth: its Origin Structure . and Evolution, 353-376. Academic Press, London.
1125-1129. TAYLOR, S. R . , & MCLENNAN, S. M . , 1979: D i s c u s s i o n
on "Chemistry, thermal gradients and evolution of the lower continental crust" by J. Tarney & B. F. Windley. J. geol. Soc. Lond., 136, 497-500. — , 1981a: The rare earth element evidence in Precambrian sedimentary rocks: implications for crustal evolution; in Kroner, A. (Ed.) Precambrian Plate Tectonics, 527-548. Elsevier, Amsterdam. 19816; The Composition and evolution of the continental crust: rare earth element evidence from sedimentary rocks. Phil. Trans. R. Soc. Lond., Ser. A, 301, 381-399. WILDEMAN, T. R., & HASKIN, L. A., 1973: Rare earths
in Precambrian sediments. Geochim. Acta,
cosmochim.
37, 4 1 9 - 4 3 8 .
YOUNG, G. M., 1969: Geochemistry of early Protero- zoic tillites and argillites of the Gowganda Formation, Ontario, Canada. Geochim. cosmochim. Acta,
33, 4 8 3 - 4 9 2 .
, 1978: Some aspects of the evolution of the Archaean crust. Geosci. Can., 5, 140-149.
R E L A T I O N S B E T W E E N V O L C A N I C ROCKS IN T H E W A R R A W O O N A G R O U P : C O N T I N U O U S OR CYCLIC EVOLUTION? M. E. Barley
Division of Mineralogy, CSIRO, Private Bag, P.O. Wembley, Western Australia 6014. Formerly: Department of Geology, University of Western Australia, Nedlands, Western Australia 6009 ABSTRACT The Archaean (O7 3500 to 3300 Ma) Warrawoona Group in the eastern Pilbara Block contains some of the oldest known sequences of well-preserved volcanic rocks. Studies of volcanics in the Kelly Belt and McPhee Dome which have experienced minimal strain and low-grade metamorphism indicate the processes involved in the development of this volcanic terrain. Thick ( > 2 k m ) sequences of basic to intermediate calc-alkaline lavas-, subvolcanic porphyry intrusions and coarse pyroclastics, represent local centres of calc-alkaline volcanism. Else. where, extensive sheets of intermediate volcanic debris (largely pyroclastic) extend from the elevated calc-alkaline volcanic centres and, together with epiclastic volcanic sediments and cherts, are interlayered with dominantly basaltic sequences of tholeiitic lavas. Volcanics in both the calc-alkaline and tholeiitic sequences display coherent textural, mineralogical and chemical trends. The chemical variation in the calc-alkaline sequences reflects a spectrum of high-level fractional crystallization trends involving the observed phenocryst phases at high fluid pressures. The chemical variation and evolution of the tholeiitic sequences can be quantitatively explained in terms of low-pressure fractional crystallization trends involving the observed phenocryst phases and relating more evolved lavas to parental (komatiitic) magmas. Sequences in which sheets of largely dacitic, calc-alkaline, volcanic debris interfinger with tholeiitic lavas locally produce the impression of volcanic cycles, widely accepted as a typical feature of Archaean volcanism. In these sequences, there is no consistent trend towards less mafic compositions which can be interpreted in terms of fractionation of a single parent magma. When the observed lateral and vertical zonation within the Warrawoona Group is taken into account, it appears that a range of compositions has been locally available for eruption throughout the history of this volcanic pile. The volcanic stratigraphy of the Warrawoona Group has been produced by the interlayering of lavas and "pyroclastics" from independently fractionating tholeiitic and calc-alkaline magmas.
INTRODUCTION The Archaean terrain in the eastern Pilbara Block is characterized by batholithic granitoid and gneiss domes with intervening greenstone domes and synclinoria (Fig. 1). The greenstone sequence (Pilbara Supergroup) can be divided into the lower, predominantly volcanic, Warrawoona Group and the overlying, predominantly clastic, sedimentary Gorge Creek Group (Fig. 2). The Warrawoona Group consists of basalts interlayered with cherty sediments and sequences of intermediate and acid volcanics. Recent age determinations (Sangster & Brook, 1977; Pidgeon, 1978; Richards, 1978; Hamilton et al.f 1981) indicate that the age of the Warrawoona Group is between 3300 and 3500 Ma. The pattern of facies variation in the Gorge Spec, Pubis geol. Soc. Aust.,
7 (1981)
Creek Group (Eriksson, 1981) indicates that the emplacement of major batholiths into their present positions post-dates the deposition of the Gorge Creek Group and that the greenstone sequence was deposited as a continuous cover over a large portion of the eastern Pilbara (e.g. Hickman, 1975b; Barley etal., 1979). Greenstone synclinoria are the result of folding related to the emplacement of the domal batholiths and do not directly reflect the shape of original basins of deposition. Metamorphism of the greenstone sequence ranges from prehnite-pumpellyite to amphibolite facies and has been accompanied by varying degrees of strain. The structural, metamorphic and-geochemieal evolution of the granitoid and gneiss domes are discussed elsewhere in this volume (see Davy & Lewis, 1981; Bettenay et at., 1981).
264
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• E3 G3
m Fig. 1.
Proterozoic
1
1 Pilbara
Post- tectonic Granitoids Other
Granitoids
Calc - a l k a l i n e Volcanic s
*
Supergroup
Major
Folds
Major
Faults
3 4 5 0 Ma (Pidgeon,1978 !
Simplified geological m a p of part of the eastern Pilbara Block, modified f r o m Hickman & Lipple (1975) and Hickman (1975a).
Most previous studies of the Warrawoona Group (Hickman & Lipple, 1975; Hickman, 1975a, 1981; Glikson & Hickman, 1981) have been at a regional scale and there is thus only limited understanding of processes involved in the development of volcanic sequences. This paper attempts to constrain interpretations of the nature of Warrawoona Group volcanism. Field studies of volcanic sequences in the Kelly Belt and McPhee Dome (Fig. 3), an area which has experienced minimal strain and low-grade metamorphism, provide an indication of the types of volcanism, eruptive mechanisms and depositional
environments represented. Petrographic and geochemical data, and detailed observations in other parts of the eastern Pilbara, indicate that the complex volcanic sequences in the Warrawoona Group result from variations in styles and sites of volcanism with time (e.g. Fig. 2) and that there are lateral facies variations, particularly in the intermediate and acid volcanic units. Regional reconnaissance geochemical studies are insufficient to unravel the evolution and genesis of tike Warrawoona Group. A better understanding of processes involved in the development of the volcanic sequences is prerequisite to such studies.
m c C l a s t i c Sediments _—I
Fig. 2.
Ferruginous Sediments
Calc-^Alkaline V o l c a n i c s (R = R h y o l i t e ) Tholeiitic V o l c a n i c s ( — = Cherty Sediments)
Composite lithostratigraphic section of the Archaean greenstone sequence developed in the eastern Pilbara Block. M c P h e e Creek and North Pole represent examples of sections dominated by calc-alkaline and tholeiitic sequences respectively. Section A — A embraces all lithologies in the sequence and is similar to the regional stratigraphy of Lipple (1975) or Hickman (1981).
THE WARRAWOONA GROUP IN THE KELLY BELT AND McPHEE DOME In the Kelly Belt and McPhee Dome, the Warrawoona Group is between 5 and 8 km thick in apparently unfaulted sections, and consists of sequences of mainly intermediate to acid calcalkaline volcanics interlayered with mainly basaltic sequences of tholeiitic lavas (Fig. 3). Steeply dipping (>60°) volcanic units in the Kelly Belt face to the east as shown by volcanic and sedimentary structures. In the centre of the McPhee Dome volcanic units dip shallowly (<30°). On the western flank of the dome, dips increase to 70° whereas on the northern flank dips rarely exceed 40°. A major subvertical NNE-trending fault occurs in the axial region of the syncline between the Kelly Belt and McPhee Dome. Preliminary Bouger anomalies indicate that the McPhee Dome has a granitoid core (Hickman, 1975a). Metamorphism of the Warrawoona Group in the area ranges from prehnite-pumpellyite fades to upper greenschist facies (Barley, 1980). Thick sequences of intermediate volcanics, dominated by subaqueous pyroclastics, have been mapped as the Duffer Formation throughout the eastern Pilbara Block (Hickman & Lipple, 1975; Hickman, 1975a; Lipple, 1975). However, the intermediate and acid volcanism responsible for these sequences was restricted to discrete centres. In the McPhee Creek region of the McPhee Dome (Fig. 3) massive and pillowed flows of low-iron, plagioclase-phyric basalts are interlayered with
andesitic and dacitic lavas, subvolcanic intrusives, coarse pyroclastics, laharic breccias and minor epiclastic sediments. Bedding in volcanoclastic and epiclastic units in this area is irregular, and the association of basic and intermediate lavas with coarse fragmental units apparently represents a proximal facies or centre of calcalkaline volcanism. Elsewhere in the McPhee Dome and in the Kelly Belt, thick (up to 2 km) sequences of intermediate volcanics are composed of subaqueous pyroclastic units interlayered with tuffaceous sediments, epiclastic volcanic sandstones and conglomerates (e.g. Barley et al., 1979). These sequences contain minor fine-grained cherty units and some are capped by laminated cherty horizons. Bedding within the sequences is remarkably uniform. Subaqueous pyroclastic units are poorlysorted, massive to crudely graded (Figs 6A & B) and were apparently deposited as subaqueous gravity flows. Individual units may have either been the direct result of explosive phreatomagmatic eruptions or subsequent mudflow deposits which originated on the slopes of an elevated volcanic centre. Various sedimentary structures in tuffaceous and epiclastic sedimentary units throughout the area indicate deposition and reworking of volcanic debris in a shallow-water to sub-aerial high-energy environment (Barleyetal., 1979). In the McPhee Dome, sequences of tuffaceous
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and epiclastic sediments, which include more abundant cherty sediments and sediments deposited by turbidity currents, are developed at greater distances (10-20 km) from McPhee Creek. Sequences of this type, capped by moderately thick (10 to 20 m) laminated cherty sedimentary units, underlie thick basaltic sequences east of Gobbos and at Copper Gorge (Fig. 3). The cherty sediments include grey and white laminated cherty sediments, silicified epiclastic volcanic sediments, black carbonaceous cherty sediments, graphitic shale, silicified carbonate, intraclast breccia and conglomerate (e.g. Barley, 1978; Barley et al., 1979). Only cherty sediments of volcanic origin are described in this paper. One kilometre east of Gobbos (Fig. 3) epiclastic volcanic sediments which reflect deposition from turbidity currents, as evidenced by the development of Bouma A and B divisions, overlie subaqueous pyroclastics and epiclastic volcanic sandstone and conglomerate. These units are between 10 and 15 cm thick, massive or internally graded, and grade upward into sequences of turbidites in which B or C, D and silicified finegrained E divisions are developed (Fig. 6C), and laminated grey and white cherty sediments with well-developed complete or partial (2 to 5 cm thick) Bouma sequences. These grey and white
Fig. 3.
cherty sediments contain microgranular chlorite, sericite and carbonate, which distinguishes them from most of the other cherty sediment types. Flame structures and load casts are developed at the base of some depositional units. Sediments of this type at the same stratigraphic level at Copper Gorge (Fig. 3) are part of a diverse assemblage of cherty sediments, many of which indicate deposition and reworking in shallow water (Barley, 1978). On the flanks of the McPhee Dome and in the Kelly Belt (Fig. 3) sequences of intermediate pyroclastics and associated sediments are overlain by a thick, dominantly basaltic sequence with minor interlayered cherty sediments and intermediate subaqueous pyroclastics and epiclastic sediments. Lavas in the dominantly basaltic sequence occur as massive flows with brecciated or pillowed flow tops or as compound pillowed flows, and range in thickness from 5 to 50m. The presence of hyaloclastic breccias at both the base and top of compound pillowed flows, and the commonly observed vertical transition from massive to pillowed lava indicates that lava flows were mainly subaqueous (e.g. Ballard & Moore, 1977; Dimroth et al., 1978). Widespread massive and pillowed flows containing between 20 and 40% vesicles (by volume) indicate that lavas were
Simplified geological map of the Kelly Belt and McPhee Dome, with localities mentioned in text. Q > Gobbos, CG = Copper Gorge, CH = Copper Hills, E = Emu Creek, B - Budjan Creek.
VOLCANIC ROCKS IN WARRAWOONA GROUP
commonly erupted into shallow-water (i.e. less than 200m: Jones, 1969). A single thick ( - 5 0 m ) flow can be traced for more than 5 km in the Kelly Belt (Barley, 1980) and the lateral continuity of thin (<5m) cherty sedimentary horizons for over 10 km indicates eruption in an environment with subdued topography (Barley et at., 1979). A variety of sediment types occur as cherty sediments in the dominantly basaltic sequence, including silicified intermediate to acid subaqueous pyroclastics, epiclastic volcanic sediments, carbonaceous cherty sediments, silicified carbonates and silicified hyaloclastic detritus. Unsilicified subaqueous pyroclastics and epiclastic volcanic sediments locally underlie cherty sedimentary horizons. On the northern flank of the McPhee Dome and in the Kelly Belt the dominantly basaltic sequence is overlain by rhyolite lava flows, subaqueous pyroclastics and epiclastic volcanic sediments (mapped as the Wyman Formation by Hickman & Lipple, 1975; Hickman, 1975a; Lipple, 1975). Angular unconformities are locally developed between rhyolite lava flows and the basaltic sequence. A suite of quartz-plagioclase porphyries, leucocratic tonalite and granodiorite porphyries and high-level granodiorites with associated stockwork Cu-Mo-W mineralization (see Marston & Groves, 1981) are penecontemporaneous with rhyolitic volcanism. Boulders of quartz-plagioclase porphyry occur in rhyolitic epiclastic conglomerates, and boulders and pebbles of quartz-plagioclase porphyry and granodiorite porphyry occur in fluvial sediments of the overlying Gorge Creek Group. . A thick (600 m) unit of massive rhyolite exposed near Budjan Creek in the Kelly Belt (Fig. 3) consists of a single large lava flow or a number of smaller flows. Individual flow lobes (cooling units) are sheet-like or bulbous and between 5 and 100m thick, with larger units more abundant in the central portion of the flow. The lava is locally flow banded or massive and columnar jointed. A brecciated flow top (upper 50 m of unit) contains abundant chalcedonic quartz veins and cavity fillings. Some coarse subaqueous rhyolite pyroclastic units near Emu Creek in the Kelly Belt (Fig. 3) contain devitrified glass and pumice fragments. Pyroclastic deposits of this type which contain essential pumice are probably the result of magmatic rather than phreatomagmatic eruptions (Heiken, 1972). An extensive volcanic sedimentary unit overlies rhyolites in the Kelly Belt. It is a 10 to 25 m thick, upward-fining sequence of sandstones and siltstones (now cherty sediments) which extends for
267 30km. Sandstone units are plane- and crosslaminated with ripple marks developed on some bedding planes. Individual units are 1 to 10cm thick and bimodal cross-laminations and ripple marks indicate deposition in a high-energy, shallow-water environment. The upper 5 to 10 m consists mainly of epiclastic volcanic siltstone layers. These are plane-laminated grey and white cherty sediments with minor low-angle (<10°) cross-laminations, intraformational conglomerates and interbedded sandstone units. The proportion of fine- to medium-grained units increases from south to north as sediments become more distant from rhyolite flows, coarse pyroclastics and quartz-plagioclase porphyries. PETROGRAPHY AND GEOCHEMISTRY OF VOLCANICS
Alteration
Recrystallization of Warrawoona Group volcanics during very low grade metamorphism has resulted in extensive redistribution of many elements and the formation of mesoscopic zones of alteration, or metadomains (Barley, 1980). Recrystallization and alteration have involved a fluid phase and mainly affect porous rocks (e.g. pyroclastic units and pillowed flows) along fractures, flow tops and other channelways. The fluid phase was apparently rich in C 0 and mineral assemblages developed reflect different activities of CO 2- Rocks with limited porosity contain pumpellyite and abundant epidote, whereas extremely porous pyroclastic and hyaloclastic units contain assemblages dominated by chlorite and carbonate. The centres of massive flows and some very large (>50 cm) fragments in pyroclastic units are relatively unaltered and contain relict igneous mineral phases in areas which have experienced metamorphism only of prehnitepumpellyite or lower greenschist facies. The style of heterogeneous alteration is also related to the composition of parent rocks. Mafic rocks developed domains rich in Ca-Al silicates, albite and quartz or chlorite and carbonate. On the other hand Ca-Al silicate metadomains are less commonly developed in rocks of dacitic and rhyolitic composition where more diffuse and pervasive sericite and sericite-carbonate alteration is prevalent. Hydrothermal alteration of this type probably occurred during burial of the volcanic units in a subaqueous volcanic environment with high heat flow. Comparison of analyses of samples from altered and least altered metadomains in single outcrops (Barley, 1980) indicates the degree of chemical mobility during recrystallization and alteration. Provided samples are carefully 2
M. E. BARLEY
268
selected Si0 , Ti0 , MnO, A1 0 , MgO, P 0 , Cr, Ni, Y, Zr and REE are close to original compositions in basic rocks. However, A1 0 , Ti0 , P 0 and Y have been mobile during sericitization of intermediate and acid rocks. The problem of alteration is accentuated by rugged hilly terrains where many outcrops are in the pallid zone of a Tertiary laterite profile and contain abundant clay minerals and secondary silica. Analyses of carefully selected samples from least altered metadomains probably approximate their original igneous compositions, as suggested by smooth and generally tightly clustered trends on variation diagrams (Barley, 1980). However, interpretations must be evaluated within the context of limitations imposed by this assumption. Reconnaissance geochemical studies do not provide an adequate basis for the recognition and interpretation of original chemical features. 2
2
2
3
2
2
2
3
5
2
5
The Dominantly Intermediate to Acid (Calc-alkaline) Sequences Members of the dominantly intermediate to acid volcanic sequences exhibit a continuum of textures, mineralogy and geochemistry, which is typical of many modern calc-alkaline volcanic series (e.g. Ringwood, 1974). The most mafic rocks are plagioclase-phyric basalts, which contain interstitial clinopyroxene (some relict grains are preserved in very low grade areas) and subhedral microphenocrysts of partly altered irontitanium oxides. These lavas can be distinguished from basalts from the dominantly basaltic sequences, with similar MgO contents, on the basis of: (a) lower FeO*, T i 0 a n d Y contents (Fig. 4), (b) slightly higher A1 0 contents, (c) fractionated REE patterns (Fig. 5; Barley, 1980). 2
2
Andesites contain more abundant piagioclase phenocrysts and quartz becomes a phenocryst phase in dacites. Prismatic chlorite, epidote and actinolite pseudomorphs after clinopyroxene microphenocrysts and subhedral partly altered iron-titanium oxide microphenocrysts are the major mafic phases in intermediate lavas. Chlorite and actinolite pseudomorphs after original amphibole (with typical rhombic cross sections) are developed in some andesites and dacites. Rhyolites contain quartz and piagioclase phenocrysts with rare mafic minerals (partly altered amphibole and biotite) in a recrystallized and commonly strongly altered quartzo-feldspathic groundmass. The calc-alkaline series exhibits fractionation trends towards S i 0 enrichment (52 to 70%), and FeO* (9 to 2%) and T i 0 (0.8 to 0.2%) depletion. Yttrium contents remain relatively constant (20 to 25 ppm) and light REE contents increase from 30 to 150 times chondritic values (Fig. 5) with increasing Si0 , while heavy REE contents remain relatively constant at approximately 10 times chondritic values (Barley, 1980). The major- and trace-element variations displayed by least altered and sparsely phyric calc-alkaline volcanics in the Kelly Belt and McPhee Dome can be quantitatively explained in terms of a spectrum of highlevel closed-system fractional crystallization trends (Barley, 1980) involving the observed phenocryst phases (i.e. cpx + pi + mt ± amphibole in basic and intermediate liquids ± biotite in acid liquids). Magmas of intermediate composition are being fractionated over a large crystallization interval (f = 0.7 to f = 0.4) for fractional crystallization schemes calculated. The 2
2
2
3
50 40 -
. ' Zr (ppm) ^ Fig. 4. Variation diagram Y us Zr for calc-alkaline (squares) and tholeiitic (circles) volcanics. N.B. : the scale is logarithmic. 20
56
1
* FeO = total iron expressed as FeO.
for calc-alkaline (stippled) and tholeiitic (plain) volcanics. The arrows indicate REE trends with decreasing MgO contents.
VOLCANIC ROCKS IN WARRAWOONA GROUP
crystallization of iron oxides or amphibole indicates fractionation at high fluid or total pressures and is consistent with the large proportion of pyroclastics in the calc-alkaline sequences. Rb-Sr whole-rock isotopic studies (Barle.y, 1980) indicate that the calc-alkaline magmas were derived from the mantle or from a short-lived crustal source with a low Rb to Sr ratio. Heavy REE contents (between 8 and 15 times chondritic values for the whole series) indicate that parent magmas were probably not derived by melting of unmodified amphibolite or eclogite (cf. Glikson, 1979). The model most successful in simulating the composition of the most mafic calc-alkaline lavas involves 5 to 7% partial melting of the mantle in which clinopyroxene is a stable residual phase, followed by the crystallization of olivine and clinopyroxene at depth. The Dominantly Basaltic (Tholeiitic) Sequences Members of the dominantly basaltic sequences display a continuum of textures, mineralogy and geochemistry which suggests that they were derived as a petrogenetically related suite. Sequences of interlayered lava flows contain rocks which range in composition from what are normally considered as pyroxenitic komatiite to iron-enriched tholeiite. The end members of this suite can be divided into komatiitic (8.5 to 12% MgO) or tholeiitic (4.5 to 7% MgO) lavas on the basis of chemistry, mineralogy, textures and field characteristics (using the criteria of Arndt et al., 1977). However, elinopyroxene-phyric basalts with between 7 and 8.5% MgO abound in the Warrawoona Group and subdivision of these lavas into komatiitic or tholeiitic groups using the criteria of Arndt et al. (1977) is completely arbitrary. Lava flows of different composition are interlayered in most sequences, and consistent trends towards less mafic compositions are seldom observed (Barley, 1980). No peridotitic lavas have been recognized in the Kelly Belt or McPhee Dome, and aphanitic or sparsely-phyric lavas with more than 12% MgO are rare. The most mafic rocks (12 to 20% MgO) are olivine and clinopyroxene cumulates from thick (10 to 50 m) differentiated flows (cf. Arndt, 1977) which have flow-top breccias with between 6 and 12% MgO. Phaneritic basalts, gabbros and pyroxenites from these flows typically have high Ca0/Al 0 ratios (>1) as the result of accumulation of olivine and clinopyroxene. Pillow lavas, margins of lava flows and thin (<5m) lava flows are generally aphanitic and contain abundant devitrified glass and skeletal crystals. The range of textures exhibited by 2
3
* FeO = total iron expressed as FeO.
269 aphanitic lavas parallels the textural variation observed in experimental studies relating crystal morphology to cooling rate (e.g. Lofgren et al., 1974) with the textures which reflect slowest cooling developed in the less mafic lavas. Lavas with between 9 and 12% MgO contain chlorite pseudomorphs after subhedral olivine microphenocrysts with skeletal overgrowths and skeletal and spherulitic relict clinopyroxene. Lavas with between 7 and 9% MgO contain prismatic clinopyroxene microphenocrysts and intergrowths of spherulitic clinopyroxene and plagioclase. The least-mafic lavas, with less than 7% MgO, contain prismatic clinopyroxene microphenocrysts, with plagioclase as prismatic microphenocrysts with skeletal overgrowths and as spherulites. Aphanitic and sparsely-phyric lavas in the tholeiitic sequence in the Kelly Belt and McPhee Dome exhibit trends towards FeO* (9.5 to 14.5%) and Ti0 (0.4 to 1.3%) enrichment. Contents of Y (13 to 37 ppm), Zr (30 to 120 ppm) and most other incompatible elements also increase steadily with decreasing MgO contents. Chondrite-nomalized REE patterns are flat to slightly enriched in light REE (Fig. 5) and vary between approximately 8 times chondritic values for lavas with 8 to 10% MgO to approximately 30 times chondritic values for lavas with 4.5% MgO. Chromium (500 to 100 ppm) and Ni (200 to 70 ppm) contents decrease as lavas become less mafic. The chemical variation of this series of lavas can be quantitatively explained in terms of a spectrum of low-pressure fractional crystallization trends involving the observed phenocryst phases (ol + cpx + pi). The more evolved lavas are related to parental (komatiitic) magmas with trace-element contents which closely reflect their mantle source (Barley, 1980). The low-pressure fractionation trends within the tholeiitic sequences are consistent with field and petrographic evidence including the abundance of clinopyroxene-phyric basalts with compositions transitional between typical komatiites and ironenriched tholeiites, -variations in phenocryst assemblages and the occurrence of numerous high-level differentiated sills in the Warrawoona Group (e.g. McCall, 1971; Hickman & Lipple, 1975). 2
DEVELOPMENT OF THE VOLCANIC SEQUENCE The type of volcanic landscape envisaged and the distribution of facies can be explained by the landforms produced by mafic, intermediate and
270
Fig. 6.
M. E. BARLEY
A. Top of coarse subaqueous pyroclastic unit overlain by medium-grained tuffaceous sediments. Kelly Belt (3 km north of Copper Hills): the compass is 8 cm in diameter. B. Subaqueous pyroclastic unit with reverse graded-bedding. Kelly Belt (6 km north of Copper Hills): the hammer is 35 mm long. C. Turbidite units in a sequence of medium- to fine-grained epiclastic volcanic sediments. McPhee Dome (1 km east of Gobbos): the note book is 5 cm long.
VOLCANIC ROCKS IN WARRAWOONA GROUP 271 acid volcanism. Mafic volcanism is generally ment of the large volume of acid magma (e.g. non-violent and mafic lava is much less viscous Rast, 1970). Rhyolite flows and pyroclastics are than either intermediate or acid lava. The bulk of preserved on the flanks of the volcanic centre the material erupted is lava which fills topo- (e.g. Lambert, 1978). graphic depressions eventually building extensive Tholeiitic volcanism and calc-alkaline volcanessentially flat lava plains (e.g. ocean floor or ism are probably at least partly penecontemporplateau basalts) or large gently sloping shield vol- aneous (Barley etal., 1979) as evidenced by intercanoes (e.g. Hawaii). layering of intermediate and acid pyroclastics and Intermediate lavas on the other hand are more epiclastic sediments which represent distal facies viscous and do not flow far from their site of of calc-alkaline volcanism, in the tholeiitic eruption. Violent and explosive eruptions are sequences. Tholeiitic volcanics pre-date interrelatively frequent, producing large elevated mediate pyroclastics in most sequences in the volcanoes or vent complexes, composed largely Warrawoona Group (Hickman & Lipple, 1975; of fragmented lava with extensive pyroclastic Hickman, 1975a). However, preliminary Bouger deposits, gravity flow units and associated epi- anomalies (reported in Hickman, 1975a) indicate clastic sediments mantling the surrounding topo- that it is unlikely that a thick sequence of basalts graphy. The frequency of eruption generally underlies proximal calc-alkaline volcanics in the decreases as the height of the volcano increases McPhee Dome. During periods of waning calc(Woodruff et al., 1979). Erosion of vent com- alkaline volcanism, and in areas remote from plexes and pyroclastic deposits also produces calc-alkaline volcanoes, sequences developed that abundant epiclastic sediments which are often consisted of basalt flows interlayered with sheets dispersed over considerable distances from vol- of silicic volcanic debris and a variety of cherty canic centres. Acid lavas are also viscous, com- sediments (Barley, 1978; Barleys a/., 1979). monly producing domes or very thick lava flows The type of volcanic landscape envisaged which do not move far from their site of eruption during the deposition of the Warrawoona Group (e.g. Guest & Sanchez, 1969). Infrequent violent in this part of the eastern Pilbara Block (Fig. 1) eruptions of acid lava produce extensive pyro- could have been produced by a series of tholeiitic clastic deposits which are subsequently eroded to basaltic fissure eruptions, or by a gently sloping shield volcano or volcanoes, with lavas interproduce silicic epiclastic sediments. The history of calc-alkaline volcanism indi- fingering with the products of more elevated calccated by the volcanic sequences in the Kelly Belt alkaline volcanic centres in a large shallow basin and McPhee Dome began with the eruption of or series of basins. Comparison of volcanic and sedimentary units plagioclase-phyric basalts and andesites into shallow water near McPhee Creek. The eruption observed in the Kelly Belt and McPhee Dome of viscous intermediate lava produced an elevated with those in a dominantly tholeiitic sequence at volcanic centre in an otherwise subdued topo- North Pole (Dunlop, 1978; Dunlop & Buick, graphy. Explosive eruption of voluminous inter- 1981; Barley et al., 1979) indicates that extensive mediate pyroclastics followed. Much volcanic shallow-water environments existed in this part debris was deposited from subaqueous gravity of the eastern Pilbara Block throughout the flows which formed an extensive sheet around the evolution of the Warrawoona Group. This is the elevated volcanic centre. Tuffaceous and epi- area in which Eriksson (1981) has recognized clastic sediments formed by erosion of volcanics platform (alluvial) sediments in the overlying in high-energy, shallow-water to subaerial Gorge Creek Group. Deep-water trough (turbienvironments, are locally important constituents dite) sediments occur to the southeast of the Kelly of these sequences. Deposits from turbidity Belt and McPhee Dome in the Mosquito Creek currents are developed in some distal sequences Synclinorium. The rapid transition from alluvial and tend to thin and become more fine-grained to marine sediments to the south-east has been with increasing distance from the volcanic centre. interpreted by Eriksson (1981) as representing the As the volcanic centre became larger and southern margin of an eastern Pilbara "protomagmas more fractionated, eruptions became craton" at the time of deposition of the Gorge more sporadic and sediments eroded from the Creek Group. elevated calc-alkaline volcanic pile were interlayered with tholeiitic sequences. Calc-alkaline DISCUSSION Many workers have emphasized the presence in igneous activity culminated with the intrusion of subvolcanic porphyries and high-level granitoids greenstone sequences of volcanic-sedimentary and the eruption of rhyolite flows and pyroclas- cycles consisting from the base upward of ultratics. Doming of the volcanic centre and adjacent mafic or mafic volcanics, intermediate or acid sequences probably accompanied the emplace- volcanics and a chert cap (e.g. Viljoen & Viljoen,
M. E. BARLEY
272
1969; Anhaeusser, 1971; Williams et al., 1976). and sites of eruption with time, and the erosion of However, in the eastern Pilbara Block calc- elevated volcanic centres are likely to result in alkaline volcanic sequences exhibit lateral facies numerous unconformities of both local and variations on the scale of several kilometres. regional significance. A detailed study of facies Sheets of largely dacitic volcanic debris inter- variations within both the volcanic sequences and finger with sequences of basaltic lavas and locally cherty sedimentary horizons is necessary before produce the impression of volcanic cycles. In the significance of any particular stratigraphic these sequences there is no genetic relationship break can be established (e.g. Glikson, 1979; between the tholeiitic and dacitic volcanics and division of Warrawoona Group into upper and no consistent trend towards more acidic composi- lower greenstone sequences) and before stratition which can be interpreted in terms of frac- graphic correlation at the level proposed by Hicktionation of a single parent magma. Clearly, care man (1977, 1981) or Glikson & Hickman (1981) is necessary in interpreting the bi-modal tholeiite- can be applied with confidence. dacite association (e.g. Barker & Peterman, 1974) observed in many Archaean terrains. This is particularly significant in isolated or small green- ACKNOWLEDGMENTS stone belts which may not exhibit full facies variThis work is the result of a Ph.D. project ations or in areas where high-strain or high-grade undertaken at the University of Western Ausmetamorphism have obliterated original textures. tralia, supervised by D. I. Groves, whose help is It is considered that the evolution of the acknowledged. The receipt of a University postWarrawoona Group was essentially a continuous graduate studentship and field support by I. D. process resulting in the formation of the observed Martin of Alcoa of Australia (W.A.) Ltd is also broadly conformable sequence. However, the acknowledged. The REE analyses reported were interaction of contrasting types of volcanism, obtained by G. D. Borley of Imperial College, variations in the importance of different styles London. REFERENCES R., 1971: The Barberton Mountain Land, South Africa—a guide to the understanding of the Archaean geology of Western Australia.
ANHAEUSSER, C .
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Komatiitic and iron-rich tholeiitic lavas of Munro Township, Northeast Ontario. J. Petrol., 18
319-369. BALLARD, R . D . , & MOORE, J . G . , 1977: Photographic
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BARKER, F . , & PETERMAN, Z . E . , 1974: Bimodal tholei-
itic-dacitic magmatism and the early Precambrian crust. Precamb. Res., 7, 1-12. BARLEY, M. E., 1978: Shallow water deposition during deposition of the Archaean Warrawoona Group, East Pilbara Block, Western Australia. Pubis geol.
Dep. and Extension Service, Univ. West. Aust., 2,
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BARLEY, M . E . , DUNLOP, J . S. R . , GLOVER, J . E . , & GROVES, D . I., 1979: Sedimentary evidence for'an
Archaean shallow-water volcanic-sedimentary facies, eastern Pilbara Block, Western Australia.
Earth planet. Sci. Lett., 43, 74-84.
BETTENAY, L . F . , BICKLE, M . J . , BOULTER, C . A., GROVES, D . I., MORANT, P . , BLAKE, T . S., &
JAMES, B.A., 1981: Evolution of the Shaw Batholith—an Archaean granitoid-gneiss dome in the eastern Pilbara, Western Australia. Spec.
Pubis, geol. Soc. Aust., 7, 3 6 1 - 3 7 2 .
DAVY, R., & LEWIS, J. D . , 1981: T h e geochemistry of
the Mount Edgar Batholith, Pilbara area, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 373-383.
DIMROTH, E . , COUSINEAU, P . , LEDUC, M . , & SAMSCHAGRIN, Y., 1978: Structure and organization of
Archean subaqueous basalt flows, Rouyn-Noranda area, Quebec, Canada. Can. J. Earth Sci., 15, 902-918. DUNLOP, J . S. R . , 1978: Shallow water sedimentation at North Pole, Pilbara, Western Australia. Pubis
geol. Dep. and Extension Service, Univ. West. Aust., 2, 3 0 - 3 8 .
DUNLOP, J. S. R., & BUICK, R., 1981: Archaean epi-
clastic sediments derived from mafic volcanics, North Pole, Pilbara Block, Western Australia.
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ERIKSSON, K. A., 1981: Platform-to-trough sedimenta-
tion, east Pilbara Block, Australia. Spec. Pubis
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A. Y., 1979: Early Precambrian tonalitetrondhjemite sialic nuclei. Earth-Sci. Rev., 15, 1-73. GLIKSON, A. Y . , & HICKMAN, A. H . , 1981: Geochemical stratigraphy and petrogenesis of Archaean basicultrabasic volcanic units, eastern Pilbara Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7,
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genesis of Archaean base-metal deposits in Western Australia. Spec. Pubis geol. Soc. Aust., 7, 409-420. MCCALL, G. J. H., 1971: Some ultrabasic and basic igneous rock occurrences in the Archaean of Western dating of the North Star Basalt, Warrawoona Australia. Spec. Pubis geol. Soc. Aust., 3, 429-442. Group, Pilbara Block, Western Australia. Spec. PIDGEON, R. T., 1978: 3450m.y. old volcanics in the Pubis geol. Soc. Aust., 7, 1 8 7 - 1 9 2 . Archaean layered greenstone succession of the PilHEIKEN, G., 1972: Morphology and petrography of volbara Block, Western Australia. Earth planet. Sci. canic ashes. Bull. geol. Soc. Am., 83, 1961-1988. Lett., 37, 421-428. HICKMAN, A. H . , 1975a: Explanatory notes on the RAST, N., 1970: The initiation, ascent and emplacement Nullagine 1:250000 Geological Sheet, Western of magmas; in Newall, G., & Rast, N. (Eds) Australia. Rec. geol. Surv. West. Aust., 1975/5. Mechanism of Igneous Intrusion, 339-362. Liverpool Letterpress, Liverpool. , 19756: Precambrian structural geology of part of the Pilbara Region. Ann. Rep. geol. Surv. West. RICHARDS, J. R . , 1978: Lead isotopes and ages of Aust. for 1974, 68-72. galenas from the Pilbara Region, Western Australia. J. geol. Soc. Aust., 24, 465-473. ' . ' 1977: New and revised definition of rock units in the Warrawoona Group, Pilbara Block. Ann. RINGWOOD, A. E., 1974: The petrological evolution of Rep. geol Surv. West. Aust. for 1976, 97. island arc systems. J. geol. Soc., 130, 183-204. 1981: Crustal evolution of the Pilbara Block. SANGSTER, D. F., & BROOK, W. A., 1977: Primitive lead Western Australia. Spec. Pubis geol. Soc. Aust., 7, in an Australian Zn-Pb-Ba deposit. Nature, Lond., 57-69. 270, 423. HICKMAN, A. H . , & LIPPLE, S. L . , 1975: Explanatory VILJOEN, M. J., & VILJOEN, R. P., 1969: A collection of notes on the Marble Bar 1 : 2 5 0 0 0 0 Geological 9 papers on many aspects of the Barberton graniteSheet; Western Australia. Rec. geol. Surv. West. greenstone terrain, South Africa. Spec. Pubis geol.
HAMILTON, P . J . , EVENSEN, N . M . , O ' N I O N S , R . K . , GLIKSON, A. Y . , & HICKMAN, A. H . , 1981: Sm-Nd
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JONES, J. G., 1969: Pillow lavas as depth indicators.
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7#-7/4,153-157. LIPPLE, S. L . , 1975: Definitions of new and revised
stratigraphic units of the Eastern Pilbara Region.
Ann. Rep. geol. Surv. West. Aust. for 1974, 5 8 - 6 3 . LOFGREN, G . , DONALDSON, C . H . , WILLIAMS, R . J . , MULLINS, O. JR, & USSELMAN, T. M., 1974: Experimentally reproduced textures and mineral chemistry of Apollo 15 quartz normative basalts.
Proc. Fifth Lunar Conf, 1, 549-567.
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WILLIAMS, I. R . , GOWER, C . F . , & THOM, R . , 1976:
Edjudina, Western Australia— 1:250 000 Geological Series. Explan. Notes geol. Surv. West. Aust.,
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WOODRUFF, L . G . , ROSE, W . I. JR, & RIGOT, W . , 1979:
Contrasting fractionation patterns for sequential magmas from two calc-alkaline volcanoes in Central America. J. Volcanol. Geotherm. Res., 6,
217-240.
ARCHAEAN CALC-ALKALINE VOLCANISM IN THE EASTERN GOLDFIELDS PROVINCE, WESTERN AUSTRALIA C. W . Giles
Department of Geology, University of Adelaide, G.P.O. Box 498, Adelaide, South Australia 5001 ABSTRACT Calc-alkaline volcanic rocks in the Yilgarn Block are confined to discrete centres and show no systematic relationship in time or space with other volcanic members of the greenstone successions. In many centres it is possible to recognise a marked fades change from dominantly subaerial volcanic rocks proximal to the vent zone to subaqueous tuffs and fine epiclastic sediments in distal regions. The latter rocks are commonly intercalated with tholeiitic basalts, indicating contemporaneity of volcanic activity rather than cyclicity. The geochemical characteristics of the basaltic and low-silica andesitic (<58% Si0 ) members of four calc-alkaline centres in the Eastern Goldfields Province are consistent with derivation by hydrous partial melting of a mantle source enriched in LIL elements over a vertical interval of 30-60km. Melting may have been promoted by introduction of water from an underlying zone of incipient melting, or by addition of heat from a mantle "hot-spot". The analogous geochemistry of low-silica andesites from two other calc-alkaline volcanic centres in the Yilgarn Block for which published data are available, suggests that such an origin may apply generally. Fractionation of the primary magmas has led to a range of acid differentiates including high-silica andesites, dacites and rhyolites. The variable contents of Y, HREE and Sr in the acid volcanic rocks in particular, indicates differing proportions of amphibole and plagioclase fractionation probably in response to variations of P TOTAL and/or P H 0 during crystallisation of the magmas before eruption. 2
2
INTRODUCTION Calc-alkaline volcanic rocks and associated epiclastic sediments form an important component of the Archaean greenstone succession in the northern portion of the Norseman-Wiluna greenstone belt (Fig. 1). Since these rocks are amenable to detailed petrological study, they offer the opportunity of advancing present knowledge of the processes of magma generation and hence of greenstone-belt development in this part of the Yilgarn Block. In this paper, previously unpublished geochemical data concerning four of the freshest and best-exposed calc-alkaline volcanic complexes in this region, are combined with other published data to investigate the possible sources of the primary magmas and their subsequent crystallisation histories. The physical processes by which the calc-alkaline volcanic piles evolved are also examined in the light of field evidence. An unavoidable aspect of the study of ancient igneous rock suites is the effect of alteration on rock composition. In the present case the influence of alteration has been minimized by analysing samples which show maximum retention of primary minerals, placing most emphasis on the
Spec. Pubis geol. Soc. Aust., 7 (1981)
reportedly relatively immobile elements and considering only consistent patterns of elemental behaviour. It is believed that this approach permits meaningful conclusions to be drawn from the geochemical data, with respect to the original magmatic processes. EVOLUTION OF THE VOLCANIC PILES A characteristic feature of the calc-alkaline volcanic rocks in the Yilgarn Block is their restriction to discrete centres. In general, calcalkaline volcanicity in a particular part of a greenstone belt can be related to the one centre or general vent zone, suggesting a brief and relatively simple episode of volcanism. The volcanic centres show no obvious association with recognised tectonic lineaments in the Yilgarn Block, although this does not exclude the possibility that crustal weaknesses may have contributed to their original development. In the Welcome Well complex (Gower, 1976), Hallberg (1980) has shown that epiclastic sediments, distal to the largely subaerial calc-alkaline centre, are intercalated with pillowed, tholeiitic basalts thus indicating approximate contemporaneity of the tholeiitic and calc-alkaline volcan-
276
C. W. GILES
Fig. 1. Location of the calc-alkaline volcanic centres referred to in the text. The complexes at Marda and Polelle have been previously described by Hallberg et al. (1976a, b). The remainder of the centres occur within the northern portion of the Norseman-Wiluna greenstone belt and have been examined and sampled by the writer.
ism. Ultrabasic and basic sills and dykes intrude the calc-alkaline volcanic pile, and at higher stratigraphic levels, both basic and ultrabasic lavas have been mapped (Hallberg, 1980). These observations do not support the concept of largescale volcanic cycles grading from ultramafic through mafic to felsic activity, as has been advocated for some greenstone belts in the Canadian Shield (e.g. Wilson & Morrice, 1977). Rocks in the calc-alkaline centres include basalt, andesite, dacite and rhyolite, although the relative proportions of each vary between centres. Acid and intermediate rocks are associated with all of the centres and it is the presence of the latter rock type which imparts the common, calc-alka-
line character to the centres considered in this paper. The acid volcanic rocks, including dacites and rhyolites, are usually pyroclastic. Typically, they are porphyritic and contain phenocrysts of zoned plagioclase and resorbed quartz in a devitrified glassy matrix which commonly shows relic vitroclastic textures. The andesites, on the other hand, are dominantly lava flows and commonly contain complexly-zoned phenocrysts of plagioclase, clinopyroxene and/or amphibole in a hyalopilitic matrix composed of altered glass and oriented plagioclase microlites. They grade, both petrographically and compositionally into basalts, which, although occurring in very minor proper-
VOLCANISM IN EASTERN GOLDFIELDS, W.A.
tion, hold the key to the origin of the primary magmas for the calc-alkaline suites. Reconstruction of the original form of most centres is limited by uncertainties arising from structural complications and lack of outcrop. However, for well-exposed centres such as the Welcome Well and Spring Well complexes, the rock types and their spatial relationships can be used to gain an insight into the volcanic environment. For example, field evidence suggests that the dominantly Intermediate Welcome Well complex represents the eroded remnant of a stratovolcano (Hallberg, 1980). This is indicated by an abundance of laharic deposits intercalated with massive and amygdaloidal lavas within the centre, and by thick deposits of epiclastic lithic wacke and greywacke adjacent to the centre. Both features indicate eruption on to a steep slope with subsequent mass-flow downslope, resulting in a skirt of epiclastic shed analogous in many ways to the ring-plain deposits that flank modern stratovolcanoes. The similarity of rock types in the other dominantly intermediate calc-alkaline volcanic centres (e.g. at Polelle, Hallberg et al., 19766; and at Ida Hill), suggests that they developed in analogous manner. In the complex at Ida Hill, the abundant intrusive plugs and dykes could be taken as evidence for exposure of a deeper level of the stratovolcano (Gower, 1976; Hallberg, pers. comm.). Contrasting with the previously described centre is the Spring Well complex (Bunting & Williams, 1976), which consists almost entirely of lensoid pyroclastic units that are inferred to have been dispersed from small vents due to nuee ardente eruptions of limited areal extent. Some of the coarse fragmental units are graded, and are thought to represent lag-fall deposits that settled out during collapse of vertical eruptive columns (Wright & Walker, 1977). Associated epiclastic sediments are relatively rare in the Spring Well complex, indicating that it may have formed a dome of less relief than the inferred stratovolcano form of the Welcome Well complex. The proximal volcanic units in both the Welcome Well and Spring Well volcanic complexes appear to have been subaerially erupted judging by the absence of hyaloclastite deposits and lack of intercalated fine sediments. However, this may not apply in all cases, since fine black shales intercalated with acid and intermediate volcanic rocks in the complex at Bore Well could indicate subaqueous volcanism, at least in part (Hallberg, 1980). A striking feature of the volcanic deposits in these centres is their marked similarity with rock
277
types found in modern volcanic environments. Thus it seems that the volcanic processes operative in the Archaean were little different to those observable today. Consequently, it may be inferred that the physico-chemical properties of the magmas generated in the Archaean and expressed in their movement, fractionation and subsequent eruption, were probably comparable to modern calc-alkaline magmas. However, the apparent absence of calc-alkaline volcanic arcs in the Yilgarn Block suggests that the tectonic processes which were ultimately responsible for the generation of the Archaean magmas differed from the subduction-related processes which are generally thought to have generated the primary magmas for modern calc-alkaline suites.
SOURCE OF THE PRIMARY MAGMA General Discussion The continuum of field and petrographic characteristics observed for volcanic rocks from the various centres could be taken as evidence for common origin..A quantitative evaluation of this proposal is facilitated by the data, plotted in Figure 2, for the most basic rocks from each of the centres (with < 6 0 % Si0 2 and > 4 % MgO). Selection of the most basic rocks, including basalts and low-silica, high-MgO andesites, minimizes the effect of elemental variations produced by differentiation. It is apparent from Figure 2 that rocks from each of the centres show analogous trace-element variations, but the absolute contents of individual trace elements range widely. The linearity of the Nb vs Zr, Y vs Ti, Sc vs Ti, V vs Ti and Cr vs Ni relationships in particular, strongly supports the view that the low-silica volcanic rocks had a common origin. If the volcanic rocks had been derived from various sources, the linearity of the plots would not be expected and indeed would be difficult to explain. Compositional limitations preclude a crustal origin for the basaltic samples (with <52% Si0 2 and > 6 % MgO, 150ppm Cr, 100ppm Ni). The comparable geochemical characteristics of the low-silica andesites plotted (with < 58% S i 0 2 a n d > 4 . 5 % MgO) indicates that they were unlikely to have been derived from an independent source. These conclusions are supported by modelling calculations which show that a basic crustal source with the composition of an average Archaean tholeiite (analysis 3, Table 1) yields levels of both compatible elements (Ni, Cr, V and Sr) and less-compatible elements (Zr, Nb, Ce and Sm) that are significantly less than those observed in the low-silica andesites (compare analysis 1,
278
Fig. 2.
C. W. GILES
Variation diagrams for minor elements in the most basic samples (i.e. basalts and low-silica andesites) from each of the centres and for selected tholeiitic basalts. The chondritic trend lines are after Nesbitt & Sun,1976. • Welcome Well, o Spring Well. • Bore Well. • Polelle (samples and some data after Hallberg et al., 19766). A Marda (samples and some data after Hallberg et al., 1976a). A Ida Hill, x Tholeiitic basalts (data from Hallberg & Williams, 1972; Nesbitt & Sun, 1976; Stolz, 1981 and unpublished data of the writer).
VOLCANISM
IN E A S T E R N G O L D F I E L D S ,
W.A.
279
TABLE I
Selected geochemical data and experimental data referred to in text Si02 A1203 fe2<5# MnO MgO CaO Na20 K20 m 2 M s
\
2
3
4 .
5
6
56.90 15-65 8.95 0.15 6.51 5.30 5.34 0.19 0.83 0.13
7
55.87 16.70 8.37 0.15 4.60 7.90 3-71 0.85 1.12 0.33
8
51.9 15.0 11.8
51.22 16.35 9-53 0.15 7.00 9.01 2.61 2.53 0.99 0.28
9 *
50.29 13-69 10.99 0.18 11.37 8.98 1.86 0.97 1.01 0.37
49.8 12.0 10.5 0.2 12.6 10.4 2.0 .0.5 •2.6
55.5 12.3 8.0 0.1 9-3 10. k 2.0 0.5 2.5
65.63 14.40 6.55 0.10 1.80 2.4 it 6.15 1.26 0.83 0.18
75-23 12.98 2.73 0.05 0.71 2-75 3.67 1.76 0.23 0.04
6.8 10.9 2.7 0.18 0.97 0.12
Total
99.9^
99.60
100.37
99.67
99-71
99.34
100.15
Zr Nb Y Ce Sm Yb Sc Ni Cr V Rb Sr Ba
106 5.0 20 *34 * 4.5 * 2.6 25 188 304 186 3 253 145 "
138 5.4 23 *36 * 3.8 * 2.0 18 20 51 154 18 278 251
61 + 3.2 +20 f 8 t 1.9 t 2.1 +40 170 367 320 9 105
188 9.1 26 50
111 6.1 23 "55 ^ 6.5 * 1.7 28 317 799 197 24 421 . 270
233 10.6 43 -57 * 6.5 - 3.3 15 13 6.1 126 23 129 613
199 13 45 *110 * 7.6 * 3.8 5 12 7 14 40 24 548
* contents determined by mass spectrometer.
20 93 150 134 20 394 346
Remainder of major and trace elements determined by XRF.
If Porphyritic andesite, containing phenocrysts of plagioclase and clinopyroxene in a matrix of plagioclase mi'crolites and altered glass, Welcome Well complex. 2: Description as for I, rock from Welcome Well complex. 3: Average of 337 tholeiitic basalts from the Eastern Goldfields region, after Hallberg & Williams, 1972. + signifies estimated values taken from Nesbitt & Sun (1976), Stolz ( I 9 8 I ) and unpublished data of the writer. 4: Porphyritic basalt containing abundant m1crophenocrysts of plagioclase and clinopyroxene in relatively mafic, glassy matrix, Welcome Well complex. 5: Basalt, containing abundant clinopyroxene, plagioclase and some possible olivine pseudomorphs, Welcome W^ll complex. 6: Preferred equilibrium liquid composition from water-saturated melting of pyrolite at 20 kb and 1100°G, after Table 3 of Green, 1976. 71 As for 6, but at 10-kb and T100°C. §{ Porphyritic daci te, containing phenocrysts of euhedral plagioclase in a finely-crystalline, quartz-feldspar matrix, Spring Well gomp1ex §? Porphyr.itic rhyolite, containing phenocrysts of euhedral plagioclase and resorbed quartz in a finely-crystalline, quartzfeldspar matrix, Spring Well complex.
Table I), Contribution of large-ion lithophile (LIL) elements from an acid crustal source (e.g. tonalitic gneiss) could be invoked to explain the deficiency in Zr, Nb, Ce and Sm, but the discrepancy in the more compatible elements (i.e. Ni, Cr, V and Sr) would still remain. An ultramafic crustal source may eliminate the calculated deficiency in Ni, Cr and V, but the possibility of greater enrichment in LIL elements (e.g. Zr, Nb, Ce fnd gm) due to the lower degrees of melting in this case is negated by the lower LIL contents of Archaean ultramafic rocks compared with mafic rocks (Nesbitt & Sun, 1976; Sun & Nesbitt, 1978). A crustal source relatively enriched in Ni, Cr, V, Sr and LIL elements compared with typical Archaean tholeiites might be suitable in theory, but such rocks are notably absent from the Archaean rock record. It therefore appears that the low-silica intermediate volcanics have not originated in the crust. Rather, their common geochemical characteristics with the basaltic members of the calcalkaline volcanic complexes (Fig. 2), supports a genetic relationship via differentiation. Thus an ultimate mantle origin is indicated, and, in view of this, it is pertinent to examine the conditions under which the primary magmas were formed in the mantle.
Conditions of Magma Generation Among the relatively fresh rocks analysed from the calc-alkaline centres, only one sample has Ni and Cr contents and an Mg number compatible with equilibration with mantle peridotite (analysis 5, Table I). Least squares approximation of the major elements of this sample in terms of a hypothetical mantle composition (calculated Archaean pyrolite: Green, 1975) and possible residual minerals (compositions after Green, 1976) shows that a good match can be obtained assuming 20-25% partial melting leaving a residue of olivine, orthopyroxene and clinopyroxene. This result accords with trace-element modelling which demonstrates that in order to duplicate the absolute contents of V (197 ppm) and Sc (28ppm) in the primitive basalt, approximately 20-25% melting of the mantle leaving a residue of 70% olivine, 20-22l/2°/o orthopyroxene and 7!/?-10% clinopyroxene is necessary (mantle abundances of Sc and V and Kds after compilation ofFrey etal.9 1978). Significantly, at the relatively high degrees of melting indicated (i.e. 20-25%), clinopyroxene is stable only under relatively hydrous conditions and at low pressures (<20kb: Green, 1973, 1976). The important implication of this is that
C. W. GILES
280
the primitive basalt was probably generated by relatively high degrees of melting of hydrous mantle at shallow depths. As a group, the more evolved basalts and lowsilica andesites (<58% Si0 ) show a significant range in MgO contents (4-8%), which appears to be quite independent of Si0 (Fig. 3C). It is also notable that within the 6-8% MgO range there is a large variation in the Ni and Cr contents (Figs 3H and I). Neither feature can be reconciled with a simple differentiation series in which the evolved basalts and andesites have fractionated from parental magmas of similar type. Rather, the degree of fractionation of individual samples, as judged by their Ni and Cr contents, ranges widely although the major-element geochemistry may be quite similar (compare analyses 1 and 2, Table I). Alteration or preferential partitioning of Ni into sulphide phases are considered to be unlikely explanations for this behaviour, owing to the moderately good positive correlation of Ni and Cr which indicates that some common factor, other than differentiation, is controlling the behaviour of both elements simultaneously (Fig. 2A). A possible explanation is provided by the experimental work of Green (1973, 1976), which demonstrates that under conditions of hydrous melting in the upper mantle, a variety of primary basic magmas, differing chiefly in their SiC>2 and MgO contents, can be produced in response to varying pressures. At 20 kb, basic magmas comparable in composition to the primitive basalt are produced (compare analyses 5 and 6, Table I), while at lOkb the primary basic magmas are significantly higher in Si0 (55-56% Si0 ) and lower in MgO (approx. 9%; see analysis 7, Table I, after Green, 1976). Thus it is possible that the primary magmas which gave rise to the low-silica, high-MgO andesites (e.g. analysis 1, Table I), were initially relatively high in Si0 and low in MgO, comparable with melts produced at 1100 °C, 10kb. Consequently many of these andesites may be little removed from their parental magmas (cf. analysis 7, Table I), hence explaining why they are relatively enriched in MgO, Ni and Cr for their particular Si0 contents. On the other hand the more fractionated andesites with lower Ni and Cr contents (e.g. analysis 2, Table I), may have been derived by extended differentiation from a parent lower in Si0 and higher in MgO, more like the primitive basalt (analysis 5, Table I.). It therefore appears that the variability of MgO, Ni and Cr observed in the low-silica andesites, can be satisfactorily explained in terms of differentiation from a range of primary magmas that have been produced as a result of partial 2
2
2
2
2
2
2
melting of the uppermost mantle over a pressure interval of roughly 10-20 kb. Nature of the Mantle Source An examination of the variation diagrams in Figure 2 shows that the elemental ratios of the basalts and low-silica andesites can be divided into three main groups including those that are: (1) less than chondritic—Sc/Ti, V/Ti; (2) approximately chondritic—Zr/Nb, Y/Ti; (3) greater than chondritic—Zr/Ti, Zr/Y, Ce/Zr. Nesbitt & Sun (1976) have demonstrated that Archaean ultramafic rocks are characterized by approximately chondritic values for these ratios, suggesting that the mantle source for the highmagnesium lavas had chondritic relative elemental abundances. Thus, the non-chondritic elemental ratios of the calc-alkaline rocks suggests a non-chondritic mantle source and/or control on some elements by residual minerals during magma segregation. The alternative of a nonchondritic mantle source necessarily implies a two-stage process in which the mantle is selectively enriched in certain elements before magma segregation. The elemental ratios indicate an order of enrichment in the primary basic magmas, compared with chondrites, of: Ce (~P) >Zr, Nb >Ti, Y >Sc, V. It has been previously shown that control by residual clinopyroxene during magma segregation can account for the levels of Sc and V in the primitive basalt. It is possible that Ti and Y behaved essentially incompatibly during the magma segregation event, thus preserving the chondritic Ti/Y mantle value in the melt. The observation that Zr/Ti and Zr/Y ratios are greater than chondritic values implies that Zr was probably enriched in the mantle source before magma segregation. If so, Zr and Nb must have been added to the mantle source in equal proportions to maintain the chondritic Zr/Nb value (assuming the source from which they were derived had chondritic Zr/Nb). The alternative interpretation that Ti and Y were controlled during the magma-segregation event (in addition to Sc and V) and that Zr and Nb behaved incompatibly, is not favoured, since it relies on the improbable assumption that Ti and Y were controlled in the correct proportions to yield a chondritic Ti/Y value. The combination of Ce/Zr ratio greater than the chondritic value and a chondritic Zr/Nb ratio for the low-silica volcanic rocks suggests, that LREE (and probably P) were added to the mantle source in an enriching event quite independent of that for Zr and Nb. Had all these elements been added at the same time (from a primitive chondritic mantle), it would be necessary to propose
VOLCANISM IN EASTERN GOLDFIELDS, W.A.
that Zr and Nb were controlled relative to Ce in the correct proportions to maintain a chondritic ratio, a situation which, as previously argued for Ti and Y, is considered unlikely. The geochemical data can offer no evidence concerning the order in which Zr and Nb, and LREE were enriched in the mantle source, nor the timing of the enrichment events with respect to magma segregation. However, the variability in the relative abundances of LREE in the low-silica volcanics and the relative variability in Ce/Zr (Fig. 2D), suggests that the processes leading to LREE enrichment in the mantle source were rather more erratic than those for Zr and Nb. The important conclusion arising from this discussion is that the elemental ratios of the lowsilica calc-alkaline volcanic rocks record complex processes of LIL-element enrichment in the mantle source before magma segregation. Some of the elemental ratios (e.g. Ti/Y, Sc/Ti and V/Ti) have, however, probably been inherited during the melting event, due to the control of V and Sc by residual clinopyroxene and the incompatible behaviour of Ti and Y. CRYSTALLISATION HISTORY In view of the common mantle origin proposed for the primary magmas of the calc-alkaline rock suites, it is relevant to examine the subsequent crystallisation of the primary magmas with the object of establishing whether the acid members of the suites are related to the andesites and basalts via differentiation. The linear inverse correlation of A1 2 0 3 , Fe 2 0 3 t , MgO, Ti0 2 and Sc with Si0 2 for rocks with >58% Si0 2 , suggests a differentiation series (Fig. 3). Notably, the Archaean rocks plot within fields defined by Cainozoic calc-alkaline volcanic rocks from a variety of localities, thus indicating marked analogies in composition. Magnetite fractionation was probably largely responsible for the control on Fe 2 0 3 t , T i 0 2 and Sc at high Si0 2 levels, although mafic minerals (e.g. clinopyroxene and amphibole) no doubt contributed during the earlier stages of differentiation. The Sr vs Si0 2 plot demonstrates two distinct trends, which could reflect fractionation of differing proportions of plagioclase (Fig. 3G). The acid rocks are also grouped into two fields according to their Y contents and this behaviour parallels that of the HREE. It is notable that the samples showing a relative depletion in Y are also relatively enriched in Sr and vice versa. This behaviour suggests that a mineral capable of controlling Y is fractionating to the exclusion of plagioclase, assuming that plagioclase is exerting the dominant control on Sr. The most probable
281
mineral is amphibole, since its crystallisation field is expanded by the same conditions that tend to contract that of plagioclase (Yoder & Tilley, 1962; Cawthqrn & O'Hara, 1976). Moreover, of the common modal minerals, amphibole is expected to exert the major control on Y and HREE, since it has the highest distribution coefficients for these elements in equilibrium with acid and intermediate melts (Arth & Barker, 1976; Gill, 1978). Lambert et al. (1974) have recognised similar chemical trends, particularly for Y (Fig. 3F), in a suite of calc-alkaline rocks from Mt Ararat. They attributed the high- and low-Y trends in their rocks to variable amounts of amphibole fractionation produced in response to fluctuating P H 2 0 during crystallisation. This explanation relied on the experimental observation that the crystallisation field of amphibole was expanded (while that of plagioclase was contracted) by increased P H 2 0 (Green, 1972). Under hydrous conditions, the load pressure can also have a large influence on the relative proportions of minerals that fractionate. For example, decreasing load pressure will cause a hydrous liquid of the appropriate composition to move from the amphibole field through the amphibole + plagioclase crystallisation field and eventually, at shallow depths (probably <10 km), into the clinopyroxene + plagioclase crystallisation field (Yoder & Tilley, 1962). The precise temperatures at these transitions will depend largely on the composition of the liquid and its water content. It is therefore evident that the depth at which the bulk of the crystal fractionation occurred, in addition to the water-vapour pressure, could have exerted a major control on the crystallisation history. If the magma was hindered during its ascent, at depths possibly corresponding to the base of the Archaean crust (e.g. 20-30 km), and if it resided at this level for some time, then prolonged amphibole-dominated fractionation would have occurred. This may have resulted in the differentiation series low in Y and HREE and high in Sr, typified by the rocks comprising the greater part of the complexes at Welcome Well, Ida Hill and Polelle. If, on the other hand, the magma ascended rapidly to shallow depths, then limited early amphibole fractionation would have been followed by prolonged clinopyroxene and plagioclase fractionation to yield the rocks high in Y and HREE and low in Sr found in the complexes at Spring Well and Bore Well. It is notable that an alternative origin, involving crustal anatexis, has been proposed by Taylor & Hallberg (1977) for the high-Y, low-Sr acid volcanic rocks of the Marda complex. This possi-
282
C. W. GILES
bility has been tested for acid rocks of similar probably only yield acid melts owing to the high composition in the Spring Well complex. Major- temperatures required to generate melts of interand trace-element modelling demonstrated that a mediate composition (Wyllie et al1976). Thus, source with the composition of an average the presence or absence of associated rocks of Archaean tholeiite (analysis 3, Table I) was too intermediate composition may be a reliable field high in Ni and Cr and too low in Zr and all REE indicator of mantle vs crustal-derived felsic to account for the levels of these trace elements in plutonic or volcanic rocks. This distinction apthe acid rocks (e.g. analyses 8 and 9, Table I). A pears to be valid in the Norseman-Wiluna greensource with the composition of an Archaean stone belt, where acid volcanic rocks lacking tonalitic gneiss also proved unsatisfactory, since associated andesites (not considered in this the characteristically low Y and HREE contents paper), have field and geochemical characteristics of these rocks (Glikson, 1979) combined with the much more consistent with an anatectic origin high proportion of residual amphibole and the (Hallberg, 1980). high degree of melting (>50%) contributed to a A further important finding of this work is that significant deficiency of Y and HREE in the cal- the primary magmas which gave rise to the calcculated melt compared with that observed (e.g. alkaline suites, appear to have been derived analyses 8 and 9, Table I). Moreover, the rela- independently of the magmas for the tholeiitic tively high levels of V and Sc in the dacites could and komatiitic series. This is supported not only not be duplicated from such a source (analysis 8, by the lack of any systematic field relationship Table I). The latter observation also argues between the three rock series, but also by the against a more acid, granodioritic source for the respective geochemical characteristics which are acid magmas. incompatible with a relationship via crystal fracIn summary, the difficulty of matching the tionation or partial melting of a common source. observed trace-element levels in the acid volcanic For example, at comparable MgO contents, the rocks by partial melting of likely crustal sources tholeiitic basalts are significantly higher in argues against an anatectic model. Moreover, the F e 0 , CaO, Sc and V, but lower in A1 0 , genetic relationship implied by the association of P 0 , Zr, Nb, LREE and MREE than the LIL element-enriched acid, intermediate and basaltic members of the calc-alkaline suites, thus often basic volcanic rocks within the one localised precluding a crystal fractionation relationship volcanic centre, cannot be reconciled with a (compare analyses 3 and 4, Table I). The approximulti-source hypothesis. This evidence, com- mately chondritis Zr/Ti, Zr/Y and Ce/Zr. values, bined with the demonstrated consistency of the characteristic of the tholeiitic basalts (Fig. 2; see trace elements, favours the previously advanced also Nesbitt & Sun, 1976), imply a mantle source hypothesis involving derivation of the acid quite distinct from that which gave rise to the magmas by crystal fractionation from inter- primary magmas for the calc-alkaline series. Premediate parents. sumably, it escaped the "enriching" process which affected the^ calc-alkaline magma source. IMPLICATIONS FOR ARCHAEAN This could indicate that the primary magmas GREENSTONE-BELT DEVELOPMENT which gave rise to the tholeiitic series were deAn important conclusion arising from this rived from primitive mantle material at greater work is that the calc-alkaline volcanic rocks in the depths, although magma segregation could have Norseman-Wiluna greenstone belt represent new occurred at considerably shallower levels if the additions to the crust from the mantle. The volu- primitive mantle material rose as a diapir (Fig. 4). minous granitoid rocks in the region, by contrast, Evidently, the conditions of magma segregation appear to be the products of major episodes of were such that clinopyroxene was rare or absent crustal reworking (Glikson & Lambert, 1976; as a residual mineral, judging by the higher Sc Archibald et a', 1978; Cooper et al., 1978). In and V contents of the tholeiitic basalts (cf basalts spite of their different origin, many granitoid of calc-alkaline series), and this may indicate a rocks have compositions comparable to the acid less hydrous mantle source (cf Nesbitt & Sun, members of the calc-alkaline suites. This arises 1976). because partial melting of a given source or Since the calc-alkaline rocks are genetically crystal fractionation from will yield acid tq the.oth^r majqr igneous rock series liquids of similar composition, provided the inunrelated Archaean granite-greenstone terrains, it respective residual and fractionating minerals follows that primary magmas were generated have analogous compositions. However, while in response tothea unique set of circumstances. Any crystal fractionation of a hydrous basic liquid petrogenetic model must account for will yield the calc-alkaline, andesite-bearing the rather sporadic, proposed occurrence of the association, partial melting of a basic source will calc-alkaline rocks localized and the relatively small 2
2
3
5
t
2
3
ppm Sr
o o
80
<
O r n > z (0) o \ • .
60
0/oMgO
&
*
• P.
0A
.
•• •
O ^A o Ogo^o
•
m >
O.* A • ^ 8 • /
2 1
f°°
GO
H ffl j
•
1
ppm Cr
ppm V
0
AO
•
400
m t-
V »i ^i
50
60
20
70 7o5i09
Fig. 3.
80
50
..
tl
^ 60
Mt Ararat-low Y
70 %Si09
J
80
200 1
Ion 0
A
A
A O
L_lj
ft
Z o o r
a 5 r
a G/3
••
6
1
>
I
%MgO
Selected variation diagrams in which the complete compositional range of samples has been plotted. The dotted field indicates the range of values observed in tholeiitic basalts and the fine stipple outlines the field of A1 2 0 3 , MgO, Fe 2 03 l and T i 0 2 in Cainozoic calc-alkaline volcanic rocks, based on a compilation by the writer. The Mt Ararat low-Y trend is after Lambert et al. (1974). Symbols as for Figure 2.
.
284
Fig. 4.
C. W . G I L E S
Schematic diagram summarizing possible hypothetical origins of the primary magmas for the calc-alkaline series, tholeiitic series and sodic granitoids. The critical feature is the L I L element-enriched mantle soiree for the primary magmas of the calc-alkaline series, in contrast with the deeper, non-enriched mantle source for the primary magmas of the tholeiitic series. In the diagram, the case of melting in response to the introduction of water and a lowered mantle solidus is illustrated, although as discussed in the text, melting above a mantle "hot-spot" is an equally plausible explanation for the origin of the magmas for the calc-alkaline series.
volumes of magma generated, and it must also provide a means whereby hydrous melting of the upper mantle can occur over a significant vertical interval. Subduction-related models, universally proposed to explain the origin of modern calcalkaline magmas, satisfy the latter requirement by invoking dehydration of a descending oceanic crustal slab and flooding of the overlying mantle wedge with the released LIL element-enriched fluids or silicate melts (Best, 1975). Melting of the mantle wedge at relatively shallow levels presumably occurs in response to the depressed mantle solidus in the presence of water. It has been noted that the calc-alkaline volcanic rocks in the Yilgarn Block do not occur in arcs, but rather in discrete, isolated centres. This is difficult to reconcile with subduction, even on a small scale, and unless field evidence in support of subduction-related processes can be demonstrated in the Yilgarn Block, this hypothesis is unlikely to be directly applicable. It is possible that melting of the upper mantle in the Archaean could have been triggered by the localized introduction of water from a zone of incipient melting at greater depths (Green, 1973;
see Fig. 4). Significant melting would have only occurred if the input of water and resultant depression of the mantle solidus was sufficient, and thus the sites of melting may have been largely controlled by zones of weakness such at faults which would have facilitated tapping and transport of the water-rich component from the zone of incipient melting to higher levels. This model requires the geothermal gradient to be of the correct order in the Archaean (comparable with that in modern ocean basins) to permit the development of a zone of incipient melting, possibly similar to the modern low-velocity zone (Green & Liebermann, 1976). An alternative mechanism, involving localized melting above a mantle "hot spot" (c/. Goodwin, 1973) has been proposed by Hallberg et at (1976a) as a possible explanation for the origin of the parental magmas which gave rise to the Marda calc-alkaline volcanic complex. In this case, the increased mantle heat flow, perhaps produced by diapirism or a radioactive decay process, is believed to have been sufficient to melt the upper mantle as it moved with the lithosphere over the "hot spot". Both models are similar in their end result in
VOLCANISM
IN E A S T E R N G O L D F I E L D S ,
terms of the m a g m a s g e n e r a t e d , a n d distinction between t h e m o n g e o c h e m i c a l g r o u n d s w o u l d be difficult, if n o t i m p o s s i b l e . In either case the conditions of m a g m a g e n e r a t i o n c r e a t e d in t h e u p p e r mantle w o u l d essentially d u p l i c a t e t h o s e p r o duced in the m a n t l e w e d g e a b o v e a descending, dehydrating o c e a n i c c r u s t a l slab in a s u b d u c t i o n zone. This m a y o f f e r a n e x p l a n a t i o n f o r the a p parent p a r a d o x in t h e E a s t e r n .Gold-fields P r o vince, where the field r e l a t i o n s h i p s of the Archaean calc-alkaline volcanic r o c k s o f f e r n o direct s u p p o r t t o s u b d u c t i o n m o d e l s , b u t the geochemistry indicates m a r k e d analogies with
ARCHIBALD, N . J . , BETTENAY, L . F . , BINNS, R . A . , GROVES, D . I . , & GUNTHORPE, R . J . , 1 9 7 8 : T h e
evolution of Archaean greenstone terrains, Eastern Goldfields Province, Western Australia. Precamb. Res., 6, 103-131. ARTH, J. G., & BARKER, F., 1976: Rare-earth partitioning between hornblende and dacitic liquid and implications for the genesis of trondhjemitic-tonalitic magmas. Geology, 4, 534-536. BEST, M. G., 1975: Migration of hydrous fluids in the upper mantle and potassium variation in calcalkalic rocks. Geology, 3, 4 2 9 - 4 3 2 . BUNTING, J . A . , & WILLIAMS, S . J . , 1 9 7 6 : E x p l a n a t o r y
notes on the Sir Samuel 1:250 000 Geological Sheet, Western Australia. Rec. geof. Surv. West. Aust., 1976/8. CAWTHORN, R . G . , & O ' H A R A , M . J . , 1 9 7 6 : A m p h i b o l e
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COOPER, J . A . , NESBITT, R . W . , PLATT, J . P . , & MORTI-
MER, G. E., 1978: Crustal development in the Agnew region, Western Australia, as shown by Rb/Sr isotopic and geochemical studies. Precamb. Res., 7, 31-59. FREY, F . A . , G R E E N , D . H . , & ROY, S . D . , 1 9 7 8 : I n t e -
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W.A.
285
m o d e r n calc-alkaline volcanic rocks which h a v e f o r m e d in s u b d u c t i o n e n v i r o n m e n t s . ACKNOWLEDGMENTS This w o r k was supervised at the University of A d e l a i d e by Dr R. W . Nesbitt while the writer was in receipt of a C o m m o n w e a l t h Research A w a r d . T h e study has b e n e f i t t e d greatly f r o m the assistance a n d i n s p i r a t i o n of Dr J . A . H a l l b e r g w h o has c o n t r i b u t e d m a n y of the ideas expressed in this p a p e r . T h e h e l p f u l c o m m e n t s of Dr Hallberg a n d an u n k n o w n reviewer substantially imp r o v e d the m a n u s c r i p t .
GREEN, D. H., 1973: Experimental melting studies on a model upper mantle composition at high pressure under water-saturated and water-undersaturated conditions. Earth planet. Sci. Lett., 19, 37-53. _, 1975: Genesis of Archaean peridotitic magmas and constraints on Archaean geothermal gradients and tectonics. Geology, 3, 15-18. , 1976: Experimental testing of "equilibrium" partial melting of peridotite under watersaturated, high-pressure conditions. Can. Mineral., 14, 255-268. GREEN, D . H . , & LIEBERMANN, R . C . , 1 9 7 6 : P h a s e e q u i -
libria and elastic properties of a pyrolite model for the oceanic upper mantle. Tectonophys., 32, 6 1 - 9 2 . GREEN, T. H., 1972: Crystallization of calc-alkaline andesite under controlled, high-pressure hydrous conditions. Contr. Mineral. Petrol., 34, 150-166. HALLBERG, J. A., 1980: Archaean geology of the Leonora-Laverton area. International Archaean Symposium, Perth 1980. Northeast Yilgarn Block Excursion Guide. HALLBERG, J . A . , JOHNSTON, C . , & BYE, S . M . ,
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The Archaean Marda igneous complex, Western Australia. Precamb. Res., 3, 111-136. HALLBERG, J . A . , CARTER, D . N . , & WEST, K .
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J.
A.,
&
WILLIAMS,
D.
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C.,
1972:
Archaean mafic and ultramafic rock associations in the eastern goldfields region, Western Australia. Earth planet. Sci. Lett,, 15, 191-200. LAMBERT, R . ST. J . , HOLLAND, J . G . , & O W E N , P . F . ,
1974: Chemical petrology of a suite of calc-alkaline lavas from Mt Ararat, Turkey. J. Geol., 82, 419-438. NESBITT, R . W . , & SUN, S . - S . , 1 9 7 6 : G e o c h e m i s t r y o f
Archaean spinifex-textured peridotites and magnesian and low magnesian tholeiites. Earth planet. Sci. Lett., 31, 433-453. STOLZ, G. W., 1981: A Petrographic and Geochemical Investigation of the Archaean Volcanic Succession in the Vicinity of the Scotia Nickel Deposit. Ph.D. Thesis, Univ. Adelaide [unpublished].
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SUN, S . - S . , & NESBITT, R. W . , 1978: Petrogenesis of
WRIGHT, J. V., & WALKER, G. P. L., 1977: The ignim-
65, 301-325. TAYLOR, S. R., & HALLBERG, J . A., 1977: Rare earth
WYLLIE, P . J . , HUANG, W . L . , STERN, C . R., MAALOE,
Archaean.ultrabasic and basic volcanics: evidence from rare earth elements. Contr. Mineral. Petrol.,
elements in the Marda calc-alkaline suite; an Archaean geochemical analogue of Andean-type volcanism. Geochim. costnochim. Acta, 41,
1125-1129. WILSON, H . D . B., & MORRICE, M . G . , 1977: T h e vol-
canic sequence in Archaean Shields; in Baragar, W. R. A., Coleman, L. C., & Hall, J. M. (Eds) Volcanic Regimes in Canada, 355-376.Spec. Pap. geol. Ass. Can., 16.
brite source problem: significance of a co-ignimbrite lag-fall deposit. Geology, 5, 729-732.
S., 1976: Granitic magmas: possible and impossible sources, water contents, and crystallization sequences. Can. J. Earth Sci., 13, 1007-1019. YODER, H . S . , & TILLEY, C . E . , 1962: Origin of basalt magmas: an experimental study of natural and synthetic rock systems. J. Petrol3, 342-532.
GEOCHEMICAL STRATIGRAPHY AND PETROGENESIS OF ARCHAEAN BASIC-ULTRABASIC VOLCANIC UNITS, EASTERN PILBARA BLOCK, WESTERN AUSTRALIA A. Y. Glikson & A. H. Hickman 1
1
2
Bureau of Mineral Resources, Geology and Geophysics, PO Box 378, Canberra City, ACT 2601 Geological Survey of Western Australia, Adelaide TerracePerth, Western Australia 6000 2
ABSTRACT A major and trace-element study of 442 eastern Pilbara volcanic rocks allows construction of a regional geochemical stratigraphy and insight into the petrogenesis of the Archaean crust. The 3.556 ± 0.032Ga old Taiga Taiga Subgroup (Hamilton et al., 1981) is dominated by tholeiitic basalt and dolerite of high Ti and low K, Al, Ni and Cr. High-Mg basalt, peridotitic komatiite, and lenses of high-Al dacite-andesite are interspersed with the tholeiites. The overlying Salgash Subgroup consists of an ultramafic unit which includes peridotitic komatiites (Apex Basalt), an upper unit of tholeiitic to high-Mg basalts (Euro Basalt), and intervening dacite-rhyolite lenses (Kelly and Panorama Formations). Tholeiitic basalt units in the overlying sedimentary Gorge Creek Group (Charteris Basalt, Honeyeater Basalt) have high K and Al and low Ti. A progressive depletion in Fe, Ti, P, Zr and Y is shown by tholeiitic basalts, constituting both an overall trend and repeated smaller-scale cycles. These trends cannot be explained by magmatic fractionation, and reflect secular mantle depletion owing to repeated partial melting events and/or progressively deeper melting in a geochemically zoned low-velocity zone. An alteration screening method using log molecular proportion plots (LMPR) indicates considerable mobility of alkali and alkaline-earth elements and relative stability of siderophile (Fe, Ti, Mn, V), some LIL (P, Zr, Nb, Y), and magnesia-related (Ni, Cr, Co) elements. A concentration toward the top of sequences of volatiles, alkalis, and Cu is attributed to synvolcanic carbonation and/or secondary leaching. Petrogenetic calculations suggest iron-rich mantle source compositions (Mg numbers in the range 80-90). Peridotitic komatiites formed by over 50% melting of mantle peridotite, possibly represent high-temperature diapiric events. Compositional gaps between peridotitic komatiites and high-Mg basalts suggest they are unlikely to be related by crystal fractionation. A continuous chemical spectrum between high-Mg basalts and tholeiitic basalts is consistent with the latter having formed from the former by fractionation of clinopyroxene, olivine, orthopyroxene and minor plagioclase. Geochemical comparisons between these Archaean basalts and modern basalts provide no conclusive evidence on Archaean tectonic environments. INTRODUCTION The Pilbara Block contains thick metamorphosed greenschist-facies volcanic sequences of a wide compositional spectrum. The rocks -studied include peridotitic komatiites (defined here as volcanic rocks with over 20% MgO), high-Mg basalts (below 20% M g O and of MgO/Al2O3 higher than 0.6), tholeiitic basalts ( M g O / A l 0 less than 0.6), dolerites, high-Al andesites, sodic to potassic dacites, sodic to potassic rhyolites and ultrapotassic rhyolites, of an age-range from 3.560 ± 0.032Ga (Hamilton et al., 1981) to ca 2 . 6 G a (J. R. de Laeter, pers. comm., 1980). Excellent outcrop, stratigraphic and structural definition, and the establishment 2
3
Spec. Pubis geol. Soc. Aust., 7 (1981)
of regional geological controls (Hickman & Lipple, 1975; Lipple, 1975; Hickman, in press) (Fig. 1), render the eastern part of this terrain particularly suited to systematic geochemical investigations. For details of the geology refer to Hickman (1981). Previous geochemical investigations have been limited to analyses of 28 volcanic rocks by Hallberg (1974) and 34 analyses of dacitic arid rhyolitic rocks reported by Hickman & Lipple (1975). A joint BMR-GSWA geochemical study commenced in 1975, including systematic collections from units indicated in Table I, and 442 volcanic samples were analyzed for all major elements, plus Ba, Rb, Sr, U, Th, Zr, Nb, La, Ce, Y, Li,
288
A. Y. GLIKSON & A. H. HICKMAN 120c
F50/A8/I
Other ° o ° O o
Honeyeater Basalt Charteris
Vv 03
1
formations
PRO TEROZOIC rocks | t 'J ARCHAEAN granitic
Euro Basalt
Stratigraphic section on which geochemicai samples were collected (arrows indicate top}
| :;>!•:';'[ Apex Basalt |
rocks
Basalt
j Towers Formation Mount Ada Basalt North Star Basalt // \ \
Fig. 1.
0 L_
20 i -
40 km I
Stratigraphic sketch map of the eastern part of the Pilbara Block, showing basalt formations and sampling traverses. Traverses: 1, Taiga River to Coongan River; 2, McPhee Reward; 3, Taiga River; 4„Shark; 5, Camel Creek; 6, Sandy Creek; 7, Spinaway Creek; 8, Charteris Creek; 9, Soanesville; 10, Bowls Gorge; 11, WymanWell; 12, Emu Creek; 13, Budjan Creek.
BASIC-ULTRABASIC VOLCANIC UNITS, PILBARA
289
TABLE I
Stratigraphic succession of the Pilbara Supergroup, indicating formations sampled Group and
Format ion"
Traverse + Numbers
Samples Ana lysed HMB
TB £ HALB
DG
9
7
21
12
8
6
15
5
PK
And
Dac
Rhy
Other
Total
Negri Volcanics
!
Rushall Slate
GROUP
WHIM
CREEK
Louden Volcanics
Hons Cupri Volcanics Warambie Basa11 Mosquito Creek Formation Lai la Rookh Sandstone SOANESV1LLE SUBGROUP
GROUP
GORGE CRE
I
Honeyeater Basalt Cleaverville Charteris
Formation
Euro Basalt Panorama
11,12,13
Apex Basalt
1,5,6,7
Towers F o r m a t i o n
1
Duffer Formation
1,6,10
Mount Ada Basalt
1,2,3,4
McPhee
U
5,7
Formation
< i co
TALGA i TALGA SUBGROUP
'1
i C9 <Z o 0 1 • OH < ZS
27
5
23
5
9
1
23
15
1 1
5
1
3
1
62 1
4
12
42
45
16
3
2
22
Formation
North Star Basalt
23
2
9
)k 5
5
91
10
90
1,2,3,4
5
7
39
24
5
8
1
k
93
Total:
15
57
148
74
27
53
46
22
442
» Li thological c o m p o s i t i o n , thickness and geochronological data presented by Hickman +
1
Corboy Formation Wyman
Q. Z3
40
Formation
Basalt
C1981)-
Traverse, locations shown o n Figure I. PK;, peridot.it ic komat i i te; H M B , high-Mg basalt; T B , tholeiitic basalt; H A L B , high-Al basalt; D G , dolerite and gabbro; And, andesite (on silica content); D a e , dacites; R h y , rhyolite; O t h e r , intrusive p o r p h y r y , pyroxenite, s c h i s t , etc.
Ni, Co, Cr, V, Cu, Zn, Ga and S. This paper discusses stratigraphic geochemical trends observed in the basic and ultrabasic volcanic units, discrimination between primary and secondary chemical characteristics, magmatic pedogenesis, and implication to the tectonic environment of volcanic activity. The geochemistry of the silicic volcanic units is discussed elsewhere (Glikson & Hickman, in press) and sampling, analytical and computing methods are described in that paper. ALTERATION SCREENS Average compositions are presented in Table II. The significance of geochemical data to magmatic processes may be obscured by interaction of water with the lava surface, burial metamorphism, thermal metamorphism, attendant metasomatism, and/or dehydration (Smith, 1968, 1969; Hart, 1971; Miyashiro, 1972; Jolly & Smith, 1972; Smith & Smith, 1976; Condie et al 1977; Glikson & Derrick, 1979; Glikson, 1979). The compositions of Pilbara basalts, dolerites and gabbros are similar (Glikson & Hickman, in press) and because the hypabyssal rocks could not have originally reacted with surface water, it follows that any alteration must be primarily of metamorphic origin. The distinct behaviour of
the elements, namely mobility of alkali and alkaline-earth elements as compared to relative stability of rare-earth and ferromagnesian elements (Smith & Smith, 1976; Condie et al., 1977; Sun & Nesbitt, 1978tf), indicates that the rocks are differentially altered. Beswick & Soucie (1978) have shown that least-altered modern oceanic, arc-trench, and calc-alkaline volcanic rocks plot in well-defined fields on log (major element A/alkali element) vs log (major element B/alkali element) plots. Because ratioing oxides are K^O or N a 0 , changes in the abundance of these mobile components would result in dislocation of points mainly within the least-altered field limits. Thus only alteration which affects other major elements is represented by deviations from these fields. A limitation of the method as applied to the present study is the lack of data for leastaltered high-Mg basalts and peridotitic komatiites. In the present study, the deviation of samples from the main Phanerozoic field has been measured using a specially designed computer program. Warrawoona Group and Gorge Creek Group data were plotted on five types of LMPR plots (Fig. 2). Data for each volcanic formation were plotted and compared separately (Glikson & Hickman, in press). S i 0 - A l 0 plots yield the 2
2
2
3
290 A. Y. GLIKSON & A. H. HICKMAN least scatter. Basalts and andesites correspond in subsequent treatment of the data. Many deviclosely to the main Phanerozoic field, whereas ating samples have anomalous trace-element komatiites, pyroxenites, and high-K rhyolites characteristics, examples being Ti, Ni, Cr, Zn often plot above this field, reflecting high Si/ A1 and Li in basic rocks. With the exception of ratios. Basalts correspond closely to the least- chloritized basic rocks there is little correlation altered Ca0-Si0 field; however, many samples between the major-elemerit anomalies and volain the andesite-dacite range tend to have low tile contents of these rocks, i.e. disturbance of Ca/Si ratios, probably reflecting loss of CaO primary element patterns was not necessarily upon alteration. Ca0-Al 0 plots also indicate related to hydration. Alternatively, later deCaO loss of the andesites and dacites, and high hydration related to prograde metamorphism Ca/Al ratios of komatiites and pyroxenites—a may have masked such a relationship. common feature of Archaean ultramafic volcanic Frequency-distribution curves for basic volrocks (Viljoen & Viljoen, 1969; Brooks & Hart, canic units (Glikson & Hickman, in press) sup1974; Arndt et al., 1977; Nisbet et al., 1977). port the assumption that Si and Al are relatively Plots on the Si-Fm diagram indicate close agree- stable (Beswick & Soucie, 1978), and are consisment of basalts with the least-altered field, tent with the interpretation that certain metallic generally high Si/Fm ratios of andesites and elements (Ti, Mn, Mg, V, Ni, Co, Zn, Ga) are dacites, and low Si/Fm ratios of komatiites. relatively immobile during alteration (Pearce & Good fits for the basalt-andesite-dacite range are Cann, 1973; Pearce et al., 1975; Floyd & Winobserved on the Ca-Fm diagram; ultramafic chester, 1975, 1978; Winchester & Floyd, 1976; rocks show low Ca/Fm ratios. Nesbitt & Sun, 1976; Condie et al., 1977; Floyd, All samples which deviate markedly from the 1977). Ce shows a wider scatter than Y, in accord least-altered fields were allotted special symbols with the relative mobility of the rare-earth WARRAWOONA GROUP & GORGE CREEK GROUP 2
2
3
log [Si02/K2Q]: log[A12p3/K20]
log [Ca0/K20]:log[Si02/K201
KOMATIITES & PYROXENITES BASALTS & DOLERITES" DACITES & ANDESITES RHYOLITES
-S 0 log[A1203/K20]
iog[SiqicK20j
log [Ca0/K20]: log [A1203/K20]
cm iogCA1203/K203
log[Si02/K20}? log[FM/K2Q]^
log [Ca0/K20]: log[FM/K2Cl
S
^ 0
no
log[FM/K20]
log CFM/K20D
. V
S i T CaO A m S p T r ^ , I S Z v S t S S : ^07 ; o • ° " dac te and rhyol t' n ,
2
a
C a
P l 0 t S
F m
2 ° >°g 20>. for A-B pairs Si0 Al 0 , CaOareas represent least-altered modern volcanic fields komatiites, pyroxenites, tholeiitic basalts, dolerites, andesites,
( l 0 g
A / K
D e l i n e a t e d
r e p r e s e n t
W
B/K
r
2
3
TABLE
A verage compositions 1
2
N S B
Si02
3
4
MAB
AB
EB
a ov fe r a g e 37 tholeiites
average of 36 tholeiites
average of 13 tholeiites
average of 5 tholeiites
51.19
48.22
49.7
4 8 , 12
II
of least-altered basic and ultrabasic components of the Warrawoona and Gorge Creek groups* 5.
6
CB
HB
7
average average of 12 of 10 Soanesville tholeiites tholeiites 48.92
50.47
8
9
10
11
12
13
14
15
NSB
MAB
AB
EB
CB
HB
NSB
average of 6 high-Mg basalts
average of 9 high-Mg basalts
average of 20 high-Mg basalts
average of 3 high-Mg basalts
average of 6 high-Mg basalts
average of 5 high-Mg basalts
average of 5 peridotite komatiites
average of 4 peridotite komatiites
MAB
average of 5 peridotit6 komatiites
AB
50.40
49-99
50.09
46.90
47.83
49.29
41.47
41.6*1
42.36
T i 02
1.24
1.61
0.87
1.22
0.53
0.81
0.61
0.9b
0.58
1.03
0.55
0.60
0.35
0.66
0.18
A1203
13.88
12.86
13.46
13.75
14.90
14.38
11.44
10.15
11.50
9.41
14.24
9.65
5-51
*».66
4.10
FeO
1 1.88
13.52
10.86
13.32
9.16
10.83
9.82
11.50
10.19
10.90
9.52
9.84
11.07
13.83
8.73
MnO
0.23
0.21
0.20
0.21
0.17
0.21
0.20
0.19
0.18
0.18
0.25
0.14
MgO
6.28
5-39
6.62
6.31
7.71
6.31
11.27
10.15
11.01
10.08
9.52
11.06
28.70
25.64
31.56
CaO
9-97
8.69
10.81
8.48
10.48
8.15
10.67
10.90
10.11
10.60
10.14
12.77
4.02
5-63
4.06
Na20
2.20
2.14
1.96
2.30
1.94
1.78
0.77
1.65
2.12
0.99
2.22
1.19
0.05
K20
0.38
0.22
0.31
0.17
0.24
0.73
0.37
0.24
0.17
0.25
0.44
0.36
0.03
0.02
0.02
P2O5
0.12
0.15
0.08
0.13
0.06
0.10
0.10
0.08
0.06
0.09
0.06
0.08
0.03
0.06
0.03
Ba
1 10
52
106
346
Fe203
93
0.22
0.19
0.19
401
77
Rb
13
7
14
6
15
21
20
9
10
23
14
20
Sr
140
142
129
112
175
119
71
110
109
98
164
41
Pb
5
6
3
2
Th
2
3
2
3
55
19
6 2
58
45
11
1
1
2
2
3
2
4
3
103
122
51
74
29
66
59
74
32
58
Nb
5
6
2
3
3
4
3
4
0
Y
26
29
18
23
24
,16
21
12
16
La
11
19
27
17
17
0.27
16
3
Zr
U
0.05
2
1
28
32
24
35
6
11
14
6
9
5
7
9
5
6
3
5
4
4
3
4
Ce
17
21
19
22
16
18
8
11
16
21
13
16
5
Li
28
20
17
8
16
9
31
14
18
12
20
6
3
V
327
350
280
324
201
317
264
258
219
264
203
268
122
173
92
Cr
169
128
236
109
414
176
875
784
749
783
697
926
4350
3721
1322
Co
45
44
32
51
47
49
94
112
58
104
216
1435
1771
Ni
66
98
156
53
5
5
3
2
13 3
4
137
198
198
177
158
Cu
47
72
90
162
77
74
80
92
85
151
84
63
30
50
14
Zn
97
108
80
102
54
84
69
75
67
76
60
55
58
68
34
Ga
17
18
14
11
12
11
.7
10
278
672
417
200
163
205
160
200
S
515
* Altered r o c k s , as defined by LMPR p l o t s , omitted NSB MAB AB EB CB HB
- North Star Basalt - Mount Ada Basalt - Apex Basalt - Euro Basalt - Charteris Basalt - Honeyeater Basalt
268
168
127
317
115
1435
120
TALGA-TALGA SUBGROUP, WARRAWOONA GROUP; McPHEE REWARD METRES
02 91 w -
01
mr
•
"
» #
-
I I I I I I I I I I I f i l l I I I I I I I t l I I I I I I* I | | I
»
9
.I I
0092 0002 0091 082 091 001
' ' I I I I I I I I I I 1*1 I I I I I I I I ! . * | ! | t
| | f*,
01
•• •
tl I I I I I I
II II l i l t
009
it
09 Ti 0
II I I I I I I I I M 1*1 I I I I 1 I I I I I I I I I I ! I . 1 1 1 1 , 1 1 !
1 111
• — • •• • ' I' IIIIIIII iIiii» i' ' '
• • •• +] i iiiiiiiii •
0L. 002 091
•
001
" •••
09
' i i i i i ' i • • i r»i i i » i i ii i i i i i i i i i i i i i , i i i ,
' '' I I I I I I I I I I l l I I i i i i i i i i i i i i i i i i i i i^ . f ,
002 009
091 •
• i n mil
i I I i rii 1 1 f •
• ?• ^
•
•
001
I I I I I I I I I I I I I• M M I I I I I I I I I l i i i i i i i i i i •. i l
09
•
•••
•
- • •
• w.
• • •
I K
j 0't J."•
v
.
-
.
"I 0.-T I I I I I I I I I I I I I I I l I L-L_l_i I I I I I•I I •l l• ;•l •I •l I•I l I .I l (1
1 1 1
' I• I I I I I |: I Ul I I I I I I •I I * I I-
Fig. 3.
I A l+l
1
fc 4 I
Stratigraphic-geochemical columnar plots of the McPhee Reward section through the North Star Basalt and the Mount A d a Basalt. The vertical scale relates to stratigraphic level above the base of the North Star Basalt. The McPhee Formation is positioned between 2100 and 2200 metres. Only values above detection limits are plotted.
A. Y. G L I K S O N & A. H. H I C K M A N 294 elements (Floyd, 1977; Sun & Nesbitt, 1978a; (best-defined trends are italicized), and a general increase in H 0 (total). The lowest units in the Condieetal., 1977; Collerson & Fryer, 1978). Mount Ada Basalt have higher Ti, Fe, Mn, Na P, Sr, Zr, Nb, Y, La, Li, V, Cu, Zn, Ga, S, CO* GEOCHEMICAL STRATIGRAPHY and H 0 , and lower Si, K, Rb and Ba than the The geochemical data (Table II) show that the highest units of the North Star Basalt. Upward, North Star Basalt and Mount Ada Basalt of the the Mount Ada Basalt shows a decrease in 77, Fe Taiga Taiga Subgroup are dominated by high-Ti, Mn, Na, P, Sr, Zr, Nb, Y, La, Ce, Li, andZn! low-K, low-Ni and low-Cr tholeiitic basalts. Only The trends in Figure 3 become even more proabout 15 percent of these samples consist of ultra- nounced when the minority of ultramafic, felsie mafic rocks, i.e. high-Mg basalts, pyroxenitic and intrusive components of the section are exand peridotitic komatiites. The overlying Salgash cluded. The most important differences between Subgroup has a significantly higher ultramafic the two formations in this area are the high component—about 40 percent of the samples H 0 + , C 0 , and S levels in the Mount Ada consisting of high-Mg basalt and peridotitic Basalt. Thus, two cycles, each showing upward komatiite, with subordinate pyroxenite (Table I). decrease in siderophile, large-ion-lithophile The remaining parts of this subgroup are mainly (LIL), rare-earth (REE), and some chalcophile composed of tholeiitic basalt and dolerite. Basic elements, occur in the McPhee Reward section. volcanic units intercalated with clastic sediments Superimposed on these are associated effects of of the Gorge Creek Group consist of Ti-poor, K- carbonation and hydration, i.e. alkali depletion rich, high-Al tholeiitic basalt and dolerite, and and possibly the introduction of sulphur. The significance of the above trends is enhanced by minor high-Mg basalt. Distribution patterns of iron and related similar patterns observed in the Taiga River siderophile elements show a general although section (Fig. 1). It is significant that, whereas Si somewhat irregular decrease in total Fe expressed and Mg levels decrease upward (partly because of as FeO from the Taiga Taiga Subgroup (11-15%) the diluting effect of higher volatiles), Ca shows to the overlying Salgash Subgroup and Gorge little stratigraphic variation in these sections, sugCreek Group (8-11%). This trend results in part gesting little correlation between Ca and C0 . It from the increase in abundance of komatiites in is not clear whether the sharp increase in LIL and the Salgash Subgroup, but also reflects a decrease REE elements upward across the chemical break in the FeO /MgO ratio and siderophile trace between the North Star Basalt and the Mount elements in tholeiitic basalts (Glikson & Hick- Ada E^salt reflects primary igneous cyclicity or is man, in press). Similar upward trends are dis- due to parbonation. The second interpretation played by Ti, Mn, V, Zr, Nb, Ce, Y and Zn would imply carbonatite-like affinities of the introduced volatiles, i.e. concomitant enrichment (Table III). The detection of geochemical trends depends of C 0 , Zr, Nb, Y, La and Ce. However, varicritically on sampling densities. Single sample- ation diagrams show little positive correlation beper-point collections along regular intervals often tween C 0 and these elements. Furthermore, reveal systematic stratigraphic trends. By con- because systematic decrease in siderophile and trast, multiple sample-per-point collections, e.g. lithophile elements occurs in both carbonated and in the 16000-m thick section between Taiga River uncarbonated successions, a primary (igneous) and Coongan River north of Marble Bar (Fig. 1), reveal wide compositional differences in indiTABLE III vidual localities. In this area local scatter is sufof some siderophile and lithophile elements in ficient to obscure possible vertical variations, Modesbasic volcanic units of the Warrawoona and except of a general upward increase in the scatter Gorge Creek groups * of H 0 + and C 0 . However, because care was taken to ensure that single samples were representative of their localities, and because many of the geochemical trends which emerge are systematic, we consider that trends along single sample-per-point traverses are meaningful. Such trends are revealed in the McPhee Reward section data, including a geochemical discontinuity between the North Star Basalt and the Mount Ada Basalt (Fig. 3). The former shows a general upward decrease in 77 Fe, Na, K, />, Rb, Zr, Nb, Y, La, Ce, Zn, and Ga 2
2
2
2
2
t
2
2
2
2
TTS
MAB
AB
11-12
12-15
9- 10
Ti02/
1.0-1.25
ppm 300-350
Zr ppm
75-150
ppm
<30
20-25
15-30
HB
9-11
0.5-0.75
0.15-0.2
200-350 150-200
250-300
25-50
25-50
2-3
2-3
Ce ppm
Zn ppm
CB
8-9
0.15-0.2
250-300 200- 250
75-100
Nb ppm
Y
EB
0.75-1.0 0.5- 0.75
0.2 -0.25
V
GCG
SS
NSB
FeO t /
15- 20
20-25
10- 15
25
75 -100
* Abbreviations as in Table II.
50-75
10-15
1-15
50-75
BASIC-ULTRABASIC VOLCANIC UNITS, PILBARA
295
nature for these trends is favoured. These intra- Ti0 , V, Zr, Y, and related elements, reflecting a formational trends are similar to the previously tholeiitic liquid line of descent, were recorded in discussed overall siderophile and lithophile ele- the Abitibi belt (Gelinas et al., 1977) and the ment depletion, indicating that the long-term Snake Lake-Kakagi Lake area of southwestern depletion involved repeated local depletion Ontario (Wilson & Morrice, 1978). cycles. A comparison between the averages of stratiSystematic trends are less obvious in the 4300 graphically low and high greenstone successions m thick Shark Gully section through the Taiga in the Yilgarn Block (Hallberg & Glikson, in Taiga Subgroup, mainly because of a greater con- press) also indicates an increase in Fe, Ti, Mn, P, tent of dolerite sills (here 45 percent of the sample Zn, Nb, Y, V, and Zr. The latter trends are akin population). The overall trend from the North to iron-enrichment shown by some Proterozoic Star Basalt to the Mount Ada Basalt, and within sequences, i.e. basic volcanics in the Mackenzie each of these units, includes a decrease in Si and District, Canada (Baragar, 1969), and Eastern Ti and an increase in K, C 0 , H 0 +, Ba, Rb, Cu Creek Volcanics, northwestern Queensland (Glikand S. When intrusive rocks are excluded, the son & Derrick, 1979). However, whereas ironNorth Star Basalt shows a well-defined upward enrichment trends arise by crystal-fractionation decrease in Ti, Fe, P, Y, Zn and Zr, as at McPhee of olivine, because of the low K values of sideroReward. However, the general upward increase phile and lithophile elements relative to mantle in alkalis is the reverse of the trend observed at residues, no petrogenetic mechanism for a deMcPhee Reward, indicating variable secondary crease in the abundance of these components by redistribution patterns in carbonated and crystal fractionation is evident. The similarity hydrated zones, involving either addition or between the MgO levels of tholeiitic basalts of the depletion of alkalis. Taiga Taiga Subgroup, Salgash Subgroup, and A detection of trends in the Salgash Subgroup Gorge Creek Group, indicates that the decrease in is hampered by its wider compositional variation FeO and Ti0 in these rocks is unlikely to reflect i.e. from tholeiitic basalt to peridotitic komatiite. crystal fractionation effects, Likely explanations In the 3000m-thick Camel Creek section (Fig. 1) of these variations, therefore, are either (1) a the Apex Basalt shows an upward decrease in Mg, vertically zoned upper mantle, in which a lowNi, and Cr, and an increase in Ca, Na, K, Sr, Rb, velocity zone is enriched upward in siderophile Cu, and S. The Euro Basalt has higher Ti, Fe, P, and lithophile components, and with partial C0 , Sr, Zr, Y, La, Ce, V, Cu, and Zn, and lower melting proceeding downward with time (S.-S. Mg than the underlying Apex Basalt. This trend is Sun, pers. comm., 1979), or, alternatively (2) the reverse of that in the Taiga Taiga Subgroup, progressive removal of siderophile and lithophile and is produced by the upward decrease in the elements from mantle source regions through suc.role of komatiites, i.e. peridotitic komatiites cessive melting events. Phanerozoic Ti depletion which occur up to the 600 metres level and high- trends have been reported by Smewing et al. Mg basalts which occur up to the 1100 metres (1975), Church & Riccio (1977), and Sun & Nesbitt (1978&). The commonly low T i 0 of arclevel. Columnar plots for the two basaltic formations trench basalts may be ascribed to a previous of the Gorge Creek Group show that the Char- depletion of the source lithosphere beneath midteris Basalt exhibits an upward increase in Na, K, ocean ridge or back-arc spreading centres. It Ba, Rb, V and H 0—indicating alkali enrich- would be of interest to examine whether early ment with hydration at high stratigraphic levels. Proterozoic basic volcanic rocks in the eastern The Honeyeater Basalt shows a weak upward Pilbara show further depletion, or whether, like decrease in Mg, Ca, K, Rb and Zr and an increase other Proterozoic basalts (Baragar, 1969; Glikin H 0 , CU, and possibly S. These trends, son & Derrick, 1979), they are rich in siderophile though not clearly pronounced, suggest sulphide elements, implying early Proterozoic mantle reenrichment associated with addition of volatiles. plenishments. An important clue to the nature of secondary The secular depletion in FeO , Ti0 , P 0 , Zr and Y abundances is analogous to trends in the metamorphic processes is furnished by the inTjakastad Subgroup, lower Onverwacht Group, crease in H 0 and C 0 with higher stratigraphic Transvaal (Glikson, 1979), the Duparquet and levels in the North Star Basalt, Mount Ada Aigubelle sections in the Abitibi belt, Ontario- Basalt, Charteris Basalt, and Honeyeater Basalt. Quebec (Baragar, 1968), and the Kenogami- These increases are related to the vertical Matheson, Birch-Uchi, and Lake of the Woods redistribution of mobile elements, i.e. alkali greenstone belts of the Superior Province elements, alkaline-earth elements, sulphur, cop(Baragar & Goodwin, 1969). However, geo- per and zinc. This can be interpreted in two difchemical trends showing increases in FeO , MnO, ferent ways: 2
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(1) Uppermost volcanic flows became more so at McPhee Reward, Shark, Bowls Gorge, hydrated and carbonated because of pro- Talga-Coongan, Camel Creek and Soanesville/ gressively lower rates of extrusion and thus The mechanism may be compared to the Cugreater opportunity for sea water-lava inter- leaching and consequent copper-carbonate ore formation at Mount Isa (Smith & Walker, 1971), action. (2) Upward migration of volatiles, alkalis and and to transport of Cu by complexing with CO2 base metals leached from the volcanic rocks molecules (Collerson & Fryer, 1978), rendering upon metamorphically induced dehydration the tops of volcanic sequences favourable loci of resulted in their concentration toward strati- base-metal mineralization. graphically high levels or beneath relatively impermeable units. Both processes may have operated. However, IGNEOUS PETROGENESIS Olivine tholeiite can form by 20-30% melting the existence of sharp breaks and the absence of gradational relations between little-carbonated of mantle peridotite, leaving a residue of olivine, and heavily carbonated volcanic units, i.e. in the orthopyroxene ± spinel (Green & Ringwood, McPhee Reward section (Fig. 3), is difficult to 1967; Green, 1971), or by fractional crystallizainterpret in terms of secondary redistribution of tion of picritic or komatiitic liquids (Arndt et al. volatiles. More likely, syn- or late-volcanic addi- 1977). Allegre et al. (1977) modelled the derivation of CO to the Mount Ada Basalt has occur- tion of bulk partition coefficients K^ (e—any red, whereas the North Star Basalt was originally trace element) from a log e vs log i plot (i—inrelatively carbonate-free. Hallberg (1974) sug- compatible "hygromagmatophile" element, gested that the carbonated volcanics formed by whose K relative to mantle refractory residues is permeating C0 -rich fluids which affected the nil). volcanic piles during deposition. Gradational inIn the present study Zr was selected as the creases in C 0 are observed in the Salgash Sub- reference incompatible element because of the group section of Chinamans Creek west of low relative to residual mantle (Allegre et ak Marble Bar (Glikson & Hickman, in press), sup- 1977; Frey et al1978) and the analytical accuporting secondary upward migration. The occur- racy of this element. Log (e) vs log (Zr) plots rence of carbonated hypabyssal dolerites sup- (Glikson & Hickman, in press) allow the derivaports secondary redistribution of C 0 , as these tion of coefficients. Only least-altered basalts, rocks would not have been exposed to sea-water. as defined on LMPR plots, are considered. From Whether carbonation occurred by syn-volcanic the K^ values the C (concentration of trace elemagmatically-related processes and/or by post- ment in the primary magma), F (wt% residual burial diagenesis and metamorphism, such car- liquid) values and residue norms related to the bonation is invariably associated with strong derivation of tholeiite basalt from primary highhydration, and in several instances with addition Mg basalt magma were computed (Table IV). of sulphur. Other concomitant element migration Only K^ values read from regressions whose patterns are very variable. The carbonation of the correlation coefficients (r) are better than 0.5 Mount Ada Basalt in the Marble Bar area has in- were considered. Where the slope exceeds 1.0, volved extreme leaching and removal of K and factors other than crystal fractionation probably Rb, whereas in the Shark Gully section of the operated (Allegre et al., 1977). Good regressions, Talga-Talga Subgroup concomitant increase in and thus possibly valid K^j values, are derived for CO and K is recorded. The widening composi- Ti, Y, P, Ga and Nb, and in some volcanic units tional ranges observed for most elements toward for Zn. For the calculation of F, the C^ Of the carbonated tops of volcanic sequences suggest primary magma is derived from average values geochemical mobility imparted by volatile migra- for the high-Mg basalts of the corresponding tion. However, the relative stability of sidero- units as shown in Table II. For the calculation of phile, magnesium-related, and some LIL ele- normative residual mineralogy, the parental ments in basic volcanic rocks is illustrated by magma was assumed to be a high-Mg basalt of well-defined frequency distribution curves for Al, Si0 = 51%; Ti0 = 0.65%; Ti, Mn, Mg, P, Zr, Nb, Y, V, Ni, Co, Zn and Ga! the 0composition: = 11%; FeO =. 12%; MnO = 0.2%; and by high correlation coefficients between Fe A1 = 13%; CaO = 11%; Na 0 = 1.0%; K 0 Ti, Mn, V; P, Zr, Nb; Ce, La, Y, and Cu, Zn Ga,' MgO = 0.2%. The results of corresponding model calS. Altered basic rocks—as defined on LMPR dia- culations are summarized as averages for the grams— tend to have low Cu abundances. How- various formations in Table IV. ever, Cu is commonly enriched toward the top of Plots of liquid fraction (F) versus normative both basic and silicic volcanic sequences along residual plagioclase yield a negative linear correwith H 0 , CO , S and/or alkali elements. This is lation, reflecting the progressively ultramafic . y
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tionation is best explained by super-heating i.e. Averages of calculated degrees of fractional crystallisa- peridotitic liquids may represent high-rate tion (F = % of residual liquid), norms of fractionated shallow-level adiabatic fusion of mantle diapirs residues (01, Op, Cp, Plag), and Ti, Y, V and Zn levels derived from depths greater than 200km. The of the primary high-Mg magma, based on data for steeper the Archaean geothermal gradient, the tholeiitic basalts of the North Star Basalt (NSB), Mount Ada Basalt (MAB), Apex Basalt (AB), Euro Basalt less the required depth of derivation of such (EB), Charteris Basalt (CB) and Honeyeater Basalt diapirs. For a modern ocean-ridge geotherm of (HB). Assumed Zr levels of the primary magma used in 15 °C/km dry liquidus temperatures (1650°C) the calculations are average values for high-Mg basalts would be attained at a depth greater than 100 km. from Table II. Only Kd values for which correlation coHowever, very high geothermal gradients of ca efficients (r) are greater than 0.5 were used in the calcu- 40°C/km may have pertained in the vicinity of lations. mantle diapirs. With a gradient of about 40°C/ km (Glikson & Lambert, 1976; Lambert, 1976) 01 Op Cp Plag Ti FormaY V dry liquidus conditions would be achieved at a tion : • n (F) cumu lates norm Primary inagma (ppm) depth of ca 40 km. From the PT diagram of 0.62 NSB 37 35 20 31 13 4227 18 281 69 Green et al. (1975), 50 percent melting of dry 0.68 MAB 36 24 19 35 20 6227 21 291 91 mantle peridotite (0.1 °7o H 0 ) would occur at 26 28 31 16 AB 3428 13 2^7 13 ' 0.70 65 about 1400°C at atmospheric pressure and about 0.84 EB " ' 5 37. 7 52 5 6227 22 325 104 1850°C at 50 kb. With mantle H 0 contents of GB 10 ' • 25 48 21 2990 1 1 191 53 HB 14 0.66 0.1-0.2%, 50 percent melting would occur at 16 19 39 25 3727 18 316 56 somewhat lower temperatures. If Archaean n = number of samples. mantle peridotite was more Fe-rich than modern pyrolite (Glikson, 1979), lower liquidus composition of the residue with increasing temperatures would apply, which may help to melting. Primary-magma-model abundances of explain the apparent extreme melting. Ti, V, Y and Zn (Table IV) commonly show Compositional gaps are observed between excellent correlations with the average of Warra- peridotitic komatiites and high-Mg basalts for woona Group and Gorge Creek Group high-Mg A 1 0 (7-10%), MgO (15-25%), Mg number basalts (Table II) which supports the validity of (60-70), Ni (above 1000 ppm in peridotitic the above calculations. Thus, it is likely that the komatiite), V (about 100 ppm or less for peridotholeiitic basalts are products of fractional titic komatiite) and on the Ol-Op-Cp-Qz diacrystallization of olivine, clinopyroxene and gram. These gaps are inconsistent with an interplagioclase from primary high-Mg magmas. This pretation of high-Mg basalt as a product of is consistent with the continuous transition, i.e. crystal fractionation of peridotitic komatiite lack'of a compositional gap, between the fields of magma, which would have resulted in a continutholeiitic and high-Mg basalt. The model high- ous compositional spectrum. By contrast, the Mg magma considered here may form by about geochemical overlap between high-Mg basalt and 33% melting of pyrolite composition, as calcu- tholeiitic basalt (and dolerite) suggests their colated by MgO + FeO mass balance (Glikson & magmatic relation. Although high-Mg basalt can be modelled from major-element composition as Hickmani in press). Peridotitic komatiites (Mg number above 75) the product of olivine fractionation from peridoof the Warrawoona Group formed by over 50 titic komatiite liquids, the compositional break percent melting of the mantle peridotite (Glikson between peridotitic komatiite and high-Mg basalt & Hickman, in press). The ascent of such liquids favours an origin by 30-50 percent melting of to the surface without crystal fractionation must mantle peridotite. Liquidus temperatures of highhave been very rapid (Green, 1972). It is possible Mg basalt liquids range between 1200-1500 °C for that, in part, peridotitic komatiites resulted from dry melting at 1 atmosphere, and lower temperaremelting of olivine-rich cumulates in shallow- tures at higher partial-H 0 pressures. In addition level magma chambers, though this requires to their compositional continuity, the intimate extreme geothermal rises. The generally fine- spatial association between high-Mg basalt and grained texture of these.rocks and the occurrence tholeiitic basalts in the field favours a view of of skeletal "spinifex-textured" features cor- high-Mg basalt as a primary magma composition roborates the view of peridotitic komatiites as which may have undergone small to moderate quench products of liquidus-temperature or degrees of olivine fractionation. The quartzsuperheated ultramafic liquids (Naldrett & normative chemistry of many of the high-Mg Mason, 1968; Lewis, 1971; Nesbitt, 1971; Arndt basalts militates against equilibration with et crl., 1977). Conceivably the ascent to the sur- (siliceous) orthopyroxene-bearing refractory face of ultramafic magmas without olivine frac- residues, which suggests that partial melting TABLE IV
Zn
U)
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A. Y. GLIKSON & A. H. HICKMAN 298 To test possible geochemical-tectonic correlaoccurred under pressures of less than 20 kb. Model calculations based on Rayleigh's. non- tions further, the data were plotted on the O.OlTiequilibrium fractionation law (Neumann et al., Zr-3Y ternary, the 0.01Ti-Zr-0.5Sr ternary 1954) suggest that the tholeiitic basic magmas (Pearce & Cann, 1973), the Ti0 -K 0-P 0 terhave formed by !4 to fractionation of clino- nary (Pearce et al., 1975) and the FeO MgOpyroxene, olivine, orthopyroxene, and plagio- A1 0 ternary. These authors showed that many clase in order of abundance, from a parental modern basalts concentrate in separate fields high-Mg magma. The evolution of tholeiitic basic corresponding to their tectonic setting, i.e. midmagmas by fractional crystallization of primary ocean ridge, ocean island, spreading-centrehigh-Mg melts has probably taken place at island, arc-trench, orogenic, and continental shallow depths, because the scarcity of high- environments. The Ti-Zr-Y diagram shows a conalumina basalts in the Warrawoona Group indi- centration of eastern Pilbara basic rocks in the cates equilibration of parental tholeiitic picrite MORB field, although boundaries with other with harzburgitic residues under pressures of less fields (especially ocean-island to continental) are than 20kb (Green et al., 1979). Extrusion of commonly transgressed. The Ti-Zr-Sr diagram high-Mg basalt and tholeiitic lavas was probably (which does not recognize ocean-island and cona more continuous and common process than the tinental environments) indicates a concentration production of peridotitic komatiite melts, whose of North Star Basalt and Mount Ada Basalt generation probably depended on transient samples in the oceanic field, Apex Basalt samples mantle diapirism accompanied by catastrophic in the fields of oceanic and arc-trench low-K melting. tholeiites, Euro Basalt samples in the oceanic field, Charteris Basalt samples in the arc-trench low-K tholeiite field, and Honeyeater Basalt in TECTONIC SETTING Determination of Archaean tectonic environ- the oceanic and arc-trench low-K tholeiite fields. ments by geochemical comparison with modern The Ti-K-P diagram shows a marked coincidence volcanics is impeded by unknown factors such as of Warrawoona Group samples with the oceanic possible changes in mantle composition since the field and a scatter of Gorge Creek Group basalts Archaean, changes in tectonic environments, across the boundary of oceanic basalts and nonsecondary chemical changes such as migration of oceanic basalts (there is a slight concentration in alkalis, and also the uncertainty as to what degree the non-oceanic field). The Fe-Mg-Al diagram existing data on modern volcanics are represen- reveals a concentration of Taiga Taiga Subgroup tative. That comparisons between the basalts of analyses about the boundary between continental the Pilbara Supergroup and modern basalts pro- and ocean island basalts, i.e. in the combined vide no conclusive evidence on Archaean en- field of within-plate basalts. Apex Basalt samples vironments seems to reflecft those problems. For plot in the MORB-ocean island-continental example, data provided by Glikson & Hickman fields, Charteris Basalt data in the MORB field, (in press) show that the North Star Basalt and the and Honeyeater Basalt data across the MORB Mount Ada Basalt resemble mid-ocean ridge and continental fields. In conclusion, these plots basalts (MORB) with respect to Ti, P and Sr, but generally indicate similarities between oceanicocean-island basalts (OIB) with respect to Ti, Na, type basalts and the Warrawoona Group basalts K, Rb, Sr, Zr and Y, arc-trench basalts (ATB) and arc-trench or non-oceanic type basalts and with respect to Mg, Ca, P, Ba and Cr, and conti- the Gorge Creek Group basalts, with the Fe-Mgnental plateau basalts (CPB) with respect to Al, Al diagram constituting an exception. However, Fe, Mg, Ca, V, Cr and Ni. Tholeiites of the Apex plots including mobile alkali and alkaline-earth Basalt and Euro Basalt exhibit the following simi- elements (K, Sr) cannot be considered as reflectlarities: MORB—Ti, Na, K, P, Ba, Sr, Zr and Ni; ing primary relations, due to secondary redistriOIB—Fe, Na, K, Rb, Y and Ni; ATB—Ca, P, bution of these elements (Smith & Smith, 1976). Ba, Rb, Zr, Y, V and Cr; CPB—Al, Fe, Mg, Ca, Evidently, while geochemical stratigraphic variV, Cr and Ni. Most geochemical comparisons ations described above are significant to mantle indicate that the Charteris Basalt resembles source composition and petrographic fractionaocean-island basalts (e.g. on the basis of Al, Mg,tion processes, they are not necessarily diagnostic Na, K, Sr, Y, V, Cr and Ni concentrations) of specific modern tectonic domains. whereas the Honeyeater Basalt is similar to continental basalts (Al, Mg, Ca, K, P, Ba, Rb, V, Cr ACKNOWLEDGMENTS and Ni). Many of the elements occur in similar paper is published with the permission of concentrations in different tectonic environments theThis Bureau of Mineral Resources, Geo(e.g. basalts containing 1.00% T i 0 a r e recordedlogyDirector, and Geophysics and the Director, Geological from MORB, OIB and ATB environments). Survey of Western Australia. 2
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Metabasalts from the Troodos Massif, Cyprus: genetic implication deduced from petrography and trace element geochemistry. Contr. Miner. Petrol 5,49-64. SMITH, R. E., 1968: Redistribution of major elements in the alteration of some basic lavas during burial metamorphism. J. Petrol., 9, 191-219. , 1969: Zones of progressive regional burial metamorphism in part of the Tasman geosyncline, Eastern Australia. J. Petrol., 10, 144-163. SMITH, R. E., & SMITH, S . E., 1976: Comments on the use of Ti, Zr, Y, Sr, K, P and Nb in classification of basaltic magmas. Earth planet. Sci. Lett., 32, 114-120. SMITH, S. E., & WALKER, K. R., 1971: Primary element dispersions associated with mineralization at Mount Isa, Queensland. Bull. Bur. Miner. Resour. Aust., 131. SUN, S . - S . , & NESBITT, R. W . , 1978a: Pedogenesis of Archaean ultrabasic and basic volcanics: evidence from rare earth elements. Contr. Mineral. Petrol., 65, 301-325. , 1978&: Geochemical regularities and genetic significance of ophiolitic basalts. Geology, 6, 689-693. VILJOEN, M . J., & VILJOEN, R. P., 1969: The geology and geochemistry of the lower ultramafic unit of the Onverwacht Group and a proposed new class of igneous rocks. Spec. Pubis geol. Soc. S. Afr., 2, 55-85.
PEARCE, J. A., & CANN, J. R., 1973: Tectonic setting of
WILSON, H . D . B . , & MORRICE, M . G . , 1978: T h e vol-
PEARCE, T . H . , GORMAN, B . E . , & BIRKETT, T . C . ,
WINCHESTER, J . A . , & FLOYD, P . A . , 1976: Geo-,
basic volcanic rocks determined using trace element analyses. Earth planet. Sci. Lett., 19, 290-300.
1975: The T i 0 - K 0 - P 0 diagram: a method of discriminating between oceanic and non-oceanic basalts. Earth planet. Sci. Lett., 24, 419-426. 2
2
2
5
canic sequence in Archaean shields. Spec. Pap. geol. Ass. Can., 16, 3 5 5 - 3 7 6 .
chemical magma type discrimination: application to altered and metamorphosed basic igneous rocks. Earth planet. Sci. Lett., 28, 4 5 9 - 4 6 9 .
THE GEOCHEMISTRY OF THE MALENE (MID-ARCHAEAN) ULTRAMAFIC-MAFIC AMPHIBOLITE SUITE, SOUTHERN WEST GREENLAND C. R. L. Friend, R. P. Hall, & D. J. Hughes Department of Geology and Physical Sciences, Oxford Polytechnic, G/psy LOT^ Headington, Oxford, OX? 0&P Department of Geology, Portsmouth Polytechnic, Burnaby Road, Portsmouth, Hants, P07 3QL 1
1
2
2
2
ABSTRACT
Malene (mid-Archaean) metavolcanic amphibolites are ubiquitous in the Archaean craton of southern West Greenland. They are juxtaposed with more ancient Amitsoq gneisses (car 3700 Ma) in the central Godthabsfjord region of the craton, but are the first formed and oldest preserved material in the northern and southern regions. The major- and trace-element chemistry of two of the best-preserved Malene supracrustal belts have been analysed by X-ray fluorescence (X.R.F.) techniques. The Ravns Store* belt, in the south of the craton, is remote from the Amitsoq gneisses. The Ivisartoq rocks occupy a central position within the Amitsoq gneiss terrain. There are minor chemical differences between the two units but both are primitive ultrabasic-basic suites and both can be subdivided into rocks which have high-and-low magnesian tholeiitic and komatiitic affinities. The composition of the Ivisartoq amphibolites tends slightly towards that of island-arc assemblages while the Ravns Stortf amphibolites are more oceanic in character, but no definitive evidence distinguishes between the formative tectonic provinces of the two belts. It is proposed that the Ravns Stord amphibolites formed primary oceanic crust while those in the Ivisartoq area possibly overlapped onto a pre-existing Amitsoq continent.
INTRODUCTION The Archaean craton of southern West Greenland is composed mainly of high-grade quartzofeldspathic gneisses and intercalated major and minor horizons of Malene supracrustal rocks (Table I). There are two main generations of gneisses; the ancient Amitsoq gneiss complex is restricted to the central portion of the craton around the major fjord northeast of Godth&b (Godthabsfjord), while the younger Nuk gneisses dominate the southern and northern regions and also invade the Amitsoq gneisses in the central area (Fig. 1). The horizons of Malene supracrustal rocks are normally composed of lithologically variable amphibolites with minor incorporated units and lenses of metasedimentary and ultrabasic material. These rocks are ubiquitous throughout the Archaean craton. This generation of supracrustal rocks is the oldest preserved material in the southern and northern regions and is invaded and disrupted by the igneous precursors of the granitic (s.l.) polyphase Nuk gneisses. In the Godthabsfjord region they have been interleaved with the older Amitsoq gneisses before their
Spec. Pubis geol. Soc. Aust., 7 (1981)
intrusion by the Nuk granitic rocks during subsequent deformation. The simplest interpretation for this arrangement is that the Malene units proximal to the Amitsoq gneisses were formed in a cover/basement relationship while those remote from . the ancient gneisses represent originally oceanic material. This possibly invokes a platetectonic model for this part of the Earth's Archaean crust. The chemistry of the remote and proximal amphibolites and the relationships between the supracrustal rocks and the gneisses are thus both critical to the understanding of the tectonic regime and processes operative in this region during the Archaean. Two of the thickest and best-preserved Malene belts have been studied in detail. The first of these is the Ravns Store* (RS) belt in the southern Fiskenaesset region which is remote from the Amitsoq gneiss (Fig. 1) (Friend, 1975) while the second is adjacent to Amitsoq gneiss and outcrops on the semi-nunatak Ivisartoq in the interior of the Godthabsfjord region (Friend & Hall, 1977). The mafic and ultramafic amphibolites in both of these belts frequently include wellpreserved, pillow-structured horizons indicating
302
C. R. L. FRIEND, R. P. HALL & D. J. HUGHES
Fig. 1. Sketch map showing the distribution of the major horizons of Malene supracrustal rocks (black) and the older, ca 3700 Ma Amitsoq gneisses (stippled) in the Archaean craton of southern West Greenland. The Malene rocks occur variably deformed throughout the craton whereas the ancient Amitsoq gneisses are restricted to the Godthabsfjord region northeast of Godthab in the centre of the craton. The Ivisartoq area is in the eastern Godthabsfjord region and Ravns Store* lies to the south of Fiskenaesset. Map compiled from data published by Allaart et a/. (1977; 1978) and Bridgwater et al (1976)
303 The Malene suite of supracrustal rocks comprises a variable succession of ultramafic and mafic metavolcanics intercalated with metasedimentary horizons and incorporated bodies, and lenses of ultrabasic (peridotitic) material. The metavolcanics are mainly amphibolites with preserved pillow structures or their deformed equivalents which presumably originated as subaqueous basic lavas. The pillows vary from 0.1 to 1 m in length and are usually composed of grey amphibolite with darker, amphibole-rich margins and thin inter-pillow selvedges of leucocratic material which effectively outline individual pillows. As the pillow structures are more intensely deformed, they develop into epidotediopside-garnet-plagioclase-hornblende mineral segregations and become compositionally-banded amphibolites which are common throughout most of the region. A few horizons of pillowed ultramafic amphibolite are locally intercalated with the more usual types. These are formed almost entirely of hornblende and are the first ultramafic pillows to be described from the Archaean of West Greenland (Hall, in press) although they occur commonly in other Archaean terrains in the sequences of komatiitic rocks (Arndt et al., 1977; Brooks & Hart, 1972; Nesbitt & Sun, 1976; Naqvi et al., 1978; Nisbet et al., 1977; Viljoen & Viljoen, 1969). The pillowed horizons are intimately associated with conformable horizons of homogeneous unpillowed amphibolite which are interpreted as flows and sills and with less common metagabbroic bodies and discordant dykes. The flows also vary in composition from normal amphibolite to ultramafic, hornblenditic varieties and have the same chemistry as the pillowed rocks. The Malene rocks on Ivisartoq define a majorfold interference structure formed by a largescale synform which is refolded about an obliquely discordant anticline (Hall & Friend, 1979). AmTtsoq gneisses occur in the cores of both the earlier synforms and complementary antiforms. This means that their pre-folding disposition was both above and below the Malene supracrustal rocks indicating an early episode of interleaving by thrusting of these two complexes (Bridgwater et al1974; Hall & Friend, 1979). There is no evidence preserved of the distances or direction of movement involved during this episode of thrusting and none to indicate whether the juxtaposition of the supracrustal rocks and ancient gneisses is an original depositional feature or purely tectonic in origin, although Chad wick & Nutman (1979) believe they have recognized a primary depositional contact between the two groups on the west coast.
G E O C H E M I S T R Y O F A M P H I B O L I T E SUITE, G R E E N L A N D
their origin as sub-aqueous lavas. The composition of the two units has been compared and the chemical relationships have been examined between these and other Archaean metavolcanic rocks and recent basic lavas from various tectonic settings. Despite a wealth of chemical data, there is little evidence to substantiate the interpretation that the proximal and remote amphibolites might have originated in different tectonic environments which could be related to different platetectonic settings. Overall, the chemistry of the two belts is similar and although each forms a highly variable suite, the compositions overlap. This reflects a similarity in formative province. The combination of chemical and structural relationships suggest that the RS amphibolites originated as oceanic crust overlying mantle (Friend & Hughes, 1978) while the proximal Ivisartoq (Iv) rocks also represent primitive basic lavas which may have overlapped onto a pre-existing continental margin. GEOLOGICAL SETTING The stratigraphy of the high-grade Archaean gneiss complex of Greenland has been described by Bridgwater et al. (1976) and is summarised in Table I. McGregor (1973) showed that the Archaean of the Godth&b district, in the centre of the craton, is composed essentially of two major generations of quartzo-feldspathic gneisses and the intermediary Malene supracrustal rocks (see Allaart et al., 1977; 1978). The older generation of gneisses, the AmTtsoq complex {ca 3700 Ma, Moorbath et al., 1972), contains enclaves of an even older suite of supracrustal rocks, the Akilia association (McGregor & Mason, 1977). PreAmitsoq supracrustal rocks also form a large belt to the northeast of the Godth&bsfjord region of Isua (Allaart, 1976) dated at ca 3750 Ma (Hamilton etal. 1978, Moorbath et al., 1977). TABLE I
Generalized portion of the stratigraphy of the Archaean craton of West Greenland Event Late Nuk g n e i s s e s Early Nuk g n e i s s e s Anorthositic
rocks
Malene s u p r a c r u s t a l r o c k s
Description
Age- (Ma)
ca
l a y e r e d g a b b r o i c and a n o r t h o s i t i c complexes pillowed ultramafic/mafic rocks p e l ! t i c metasediments
2800 3000 7
ca
3000
Heterogeneous q u a r t z o feldspathic gneisses
3700 c a 3750
tonalitic - granodioritic p e g m a t i t e banded g n e i s s e s
ca
?
AmeraUk dykes Amitsoq g n e i s s Akilfia ( = l s u a ? )
association
ca
C. R. L. FRIEND, R. P. HALL & D. J. HUGHES 304 The granitic precursors of the Nuk gneisses equivalents of the Nuk granitic gneisses (ca were injected as sheets along the AmTtsoq-Malene 3000-2800 Ma, Pidgeon & Hopgood, 1975; thrust planes and emplaced syn-tectonically as Pidgeon & Kalsbeek, 1978). There is no field bodies into developing fold cores. Numerous dis- evidence of older Amitsoq-type gneisses in this cordant igneous relationships are preserved on region and isotopic studies have shown that the Ivisartoq between the Nuk gneisses and the older Nuk gneisses are doniinantly mantle-derived and Amitsoq gneisses and Malene rocks. cannot have been formed by the reworking of an The least-deformed equivalents of the Malene older sialic basement (Moorbath, 1978; Moorsupracrustal rocks in the Fiskenaesset region, in bath & Pankhurst, 1976). Additionally, lead isothe southern part of the Archaean craton, are tope studies from Godth^bsfjord indicate that represented in the Ravns Stord amphibolite belt both anorthositic and gneissic rock (Nuk) have (Andersen & Friend, 1973; Friend, 1975). The lead isotope systems contaminated by unradiosupracrustal rocks are the oldest preserved genic Amitsoq lead, a feature which is absent material in this region and are intruded by from the Fiskenaesset region (Taylor et al., in
f ° <* «S) metavolcanic amphibolites presented a?weLftTo o iH ' , 7H and nara^Thfh 1 /i T™\> c P diagrams are drawn by eye ; S - S fi h ? r J - Open circles: Lv tholeiitic and highZ l e S S h t n ? circles: v amphibolites which have komatiitic chemical affinities; x: RS tholeiitic, high-Mg and ultrabasic rocks; + : RS komatiitic amphibolites m
t
t h C I V
rt
q (IV) a n d
R a v n s Stor
T h e t r e n d H n e s d r a w n o n t h e T i a n d
fr0n
R S
C k S
a
e q u i v a l e n t
v a l u e s
GEOCHEMISTRY OF AMPHIBOLITE SUITE, GREENLAND
305
TABLE II
XRF analyses of representative supracrustal amphibolites from Ivisartoq (1-6) and Ravns Stord (7-12) Tholeiitic :
wt % Si 02A h 03 Fe 2 Qa . MgO CaO\ Na 2 0 K2O : Ti 0.2 MnO Ms? ppm So V er Ni Ga. Zn Rb Sr V Zr Nb La Ce Nd
'
„
Komati itic
Tholeiitic
6
1 207627
2 200421
3 200888
4 200887
5 207625
45,99 12.14 12.43 16.20 9-65 2.00 0... 7§ 0.59 0.21 0.01
47.43 13-58 12.97 11.37 11.65 2.13 ND 0.70 0.22 ND
55.12 12.3 7.42 5.22 17.03 1.59 0.04 0.49 0.21 0.07
44.52 8.65 12.87 21.78 10.49 0.77 0.19 0.38 0.24 0.10
46.49 9-37 12.35 17-05 . 11.96 2.03 0.02 0.55 0.23 ND
4 7.14 13.34
37 207 878 330 59 73 53 134 12 34 ND
49 281 365 145 25 77 ND 52 15 33 1
41 218 919 181 49 52 ND 54 12 34 ND
36 191 2467 723 ND 82 9 13 7 24 ND
41 206 1765 556 ND 81 9 37 9 28 .1
52 267 1104 310 67 106 9 116
5 4 1
3 7 ND
3 6 ND
4 2 ND
2 7 1
207619
12. 6212 10. 13.52 2.02 0.16 0.69 0.29
0.02
7 119858
8 129983
9 119825
42.25 9.52 12.96 30.75 4.08 0.01 0.03 0.38 0.17 0.17
44.14 11.15 15-14 19.87 7.19
47.57 13.63 13.66 10.74
0.08 0.66
0.22
37 241 1591 575 35
60
8 . 26
43. ND 5 6 ND
ND
13 40 ND
4 4 ND
3 9 3
49.20
10.78 13.73 11.46 10.92 2.45 0.29 0.84 0.24 0.09
0.34 0.15 0.04
27
32 206 1994 705 ND
39 260 310
101
90 ND 114
107 ND
20 137 4368 1493 102 71 ND 7 7 26 ND
16 4
ND 3 ND
3 5
65 167 32 125 ND
ND
129970
46.88 9.15 13.78 19.81 8.28 1.09 0.07 0.65
1.46 0.17 0.18
250
101
12
11
129964
47.75 . 5.77 12.63 26.61 6.61 0.09
0.01
2.61 0.16
0.06
141 3375 1187 ND 77 ND 15
8
9.82
0.78
26
10
129985
208 26
6
0.22
0.07
110 73
ND 29 13 36 . ND
17
56 ND
4 14
1
11
ND: not detected; all iron presented as F e 2 0 3 ; six-figure sample number corresponds to GGU field number. Analyses recalculated on a vol ati1e-free basis.
press). Only a few inclusions of possibly older istics are illustrated in Figure 2 and representative gneisses have been identified in one area of the chemical analyses are presented in Table II. The Sukkertoppen region, in the north of the craton most notable features are the generally high MgO contents and low values of K 0 , T i 0 and P 0 , (Hall, 1978). The most significant feature of the stratigraphy and the amphibolites can be subdivided into those concerning the Malene supracrustal rocks is that which display an overall chemical resemblance to in the southern and northern regions they are the modern primitive low-K tholeiites (Engel et al., earliest-formed rocks and hence can be inter- 1965; Manson, 1967) and those which are preted as primary oceanic crust, whereas in the komatiitic in character (Arndt et al., 1977; central region they are proximal to a pre-existing Brooks & Hart, 1974; Viljoen & Viljoen, 1969). gneissic basement complex on to which they may Some of the tholeiitic characteristics are illustrated by the discriminatory diagrams proposed have been deposited. by Irvine & Baragar (1971) and Kuno (1966) (Figs 3,4). GEOCHEMISTRY OF THE AMPHIBOLITES ( K O + Na 0 ) The major- and trace-element (Sc, V, Cr, Ni, Cu, Zn, Rb, Y, Sr, Zr, Nb, La, Ce, Nd) chemistry 3 of the amphibolites has been determined by X.R.F. methods using a Philips 1410 semi2 automatic X-ray spectrometer in the Department of Geology, Portsmouth Polytechnic, and the 1 data-retrieval techniques and programmes of Brown et al. (1973). Six U.S.G.S. reference 0 samples were analysed concurrently as an 40 45 50 i0 accuracy check and show a good overall correla+ N a 0 ) versus S i 0 diagram showing tion with the values quoted by Flanagan (1973) Fig. 3. (theK 0predominant tholeiitic character of both (values may be obtained on request from the Iv and RS metavolcanic amphibolites. The R.P.H.). lines separating alkali (A), high-alumina The Malene metavolcanics from both the (HAl) and tholeiitic basalts (Th) are from Ivisartoq and Ravns Store* horizons have wide Kuno (1966) and Irvine & Baragar (1971), Dots: Iv; crosses: RS; the tholeiitic and chemical ranges from ultrabasic to basic types. komatiitic rocks are not differentiated. The principal major-element chemical character2
2
2
2
2
S
2
2
2
2
5 5
5
C. R. L. FRIEND, R. P. HALL & D. J. HUGHES
306
found in other Archaean terrains (Arndt et al., 1977; Nesbitt & Sun, 1976; Nisbet et al., 1977; Williams, 1972). No spinifex textures have been preserved in these amphibolite-facies rocks but the pillow structures indicate that they represent rapidly cooled subaqueous extrusive lavas. There is a large chemical overlap between the tholeiitic (MgO 5-16%) and komatiitic types (MgO 9-22%) and the chemical classification boundary appears to be arbitrary. The rocks which have been interpreted as komatiitic satisfy the chemical parameters M g O > 9 % , C a 0 / A l 0 > 0 . 9 , FeO*/ (FeO* + M g 0 ) < 0 . 6 (FeO* is total iron as FeO), T i 0 < 0 . 9 % and S i 0 < 5 3 % (Figs 5, 6) (see Arndt et al., 1977; Brooks & Hart, 1974). They also have relatively low alumina and high Ni and Cr contents (Figs 7, 8) compared to abyssal tholeiitic rocks (Kay & Hubbard, 1978; Miyashiro & Shido, 1975; Prinz, 1967). Apart from Ni and Cr, the trace-element chemistry of the amphibolites lends little weight to their subdivision into tholeiitic and komatiitic types but emphasises the overlap between the two groups. The absolute and relative concentrations of Ti, Zr, Y and Sr vary between those of abyssal and island-arc tholeiites (Pearce & Cann, 1973; Pearce & Norry, 1979; Jakes & White, 1972) and the entire spectrum of ratios is covered by both komatiitic and tholeiitic types (Fig. 9). This situation is reflected by most of the other analysed trace elements (Fig. 10), the ranges in values of both types of amphibolite corresponding broadly to those quoted in the literature for abyssal and island-arc tholeiites. The Iv amphibolites do not 2
2
Fig. 4. AFM diagram showing the primitive tholeiitic character of both the Iv and RS rocks. The line separating the tholeiitic (Th) and calcalkaline (CA) fields is from Irvine & Baragar (1971). Symbols as in Figure 3.
The entire chemical range of the Iv amphibolites is represented by both pillowed and unpillowed amphibolites (Hall, in press). The more mafic rocks have several chemical features which differ from tholeiitic basalts and resemble more closely komatiitic rocks which are commonly MgO
3
2
AI2O.
oN o c&0 — .
°
t!
0
G
o
(
e
10 AI 2 0 3
MgO : CaO : A1 0 diagram showing the composition of komatiitic and tholeiitic metalavas from Ivisartoq and Ravns Store*. The diagram illustrates the continuum between the two chemical types and the broad overlap in compositions between the rocks of the two regions. The RS ultrabasic rocks plot nearest to the MgO apex. The more evolved Iv gabbros and dykes are not drawn but plot toward Fig. 6. the A1 0 apex. T: average oceanic tholeiite of Engel et al (1965); O and C: average oceanic and continental tholeiite respectively of Manson (1967). Other symbols as for Figure 2. 2
2
3
3
0.3
0.4
0.5
0.6
0.8
0.7
FeO*/(FeO*+ MgO)
Plot of A1 0 versus FeO*/(FeO*+ MgO) where FeO* is total iron as FeO. Dashed line encloses field of. komatiites from Canada, Australia, India,. Rhodesia and South Africa (see Arndt et at, 1977). T, C and O as for Figure 5. Other symbols as for Figure 2. 2
3
5
GEOCHEMISTRY OF AMPHIBOLITE SUITE, GREENLAND
307
5000-
•5000
Ni ppm
Cr ppm
*
+ +xX«f +
- t
1000 -
1000
O /'o^Jb
oh
+
• "N ^
100
•
-
o
\&xO
>
0\;00y# t. x Vox O o Y, o AT xx\\ xx O \ --O. O.--'
xX
\
\
\ ox'-. \ \ o--. \ A
o
\ \
\\
x
AT
° A A x
\
\ IACM
IACM
\
,\ \
A \ -
•
i
\
\
\
10
100
\ o \
o\
x
i
d 10
FeOyMgO Fig. 7.
Plot of Ni (ppm)versus FeO*/MgO for metavolcanic amphibolites from Ivisartoq and Ravns Stored (FeO* is total Fe expressed as FeO). The RS ultrabasic units and the komatiitic rocks have Ni greater than typical tholeiitie values. The fields for abyssal tholeiites (AT) and island-arc and continental-margin tholeiites (IACM) are from Miyashiro & Shido (1975). Symbols as for Figure 2.
bear chemical similarities with continental tholeiites. The RS amphibolites make up a suite which is chemically similar to the Iv rocks. However, the most obvious difference in the RS belt is the presence of ultrabasic units (MgO 20-33%) which occur as intercalated flows and layered sills (Friend, 1975). Only one of the Iv amphibolites has a similar MgO content (22%) and this sample is from one of the pillow-structured ultramafic horizons. The chemistry of the RS ultrabasic amphibolites indicates that this suite may be even more primitive than any now occurring proximal to the ancient Amitsoq gneisses. Low-magnesian (ca 5%) pillows are present in the RS suite, but all
Fig. 8.
Plot of Cr (ppm) versus FeOVMgO showing the behaviour of chromium analogous to that of nickel.
of these samples show severe chemical alteration with respect to most elements and the analyses have, therefore, been disregarded. Unlike the Iv amphibolites, the most Mg-rich RS samples do not correspond to a komatiitic chemistry since many have a C a 0 / A l 2 0 3 ratio of ca 0.4 (Fig. 5) although there is a large overlap between the RS komatiites (MgO 11-27%) and tholeiites (MgO 5-33%) (Fig. 2). Although the ranges of values of the other m a j o r elements are similar for the amphibolites of both regions, they show slightly different petrogenetic trends in that at corresponding MgO values the incompatible elements K, Ti and P are all lower and Ca is higher in the Iv than in the RS
C. R. L. FRIEND, R. P. HALL & D. J. H U G H E S
Fig. 9. Ti/lOO:Zr:Y/3 and Ti/100:Zr:Sr/2 discriminatory diagrams for metavolcanic amphibolites from Ivisartoq (a) and Ravns Storcf (b) (after Pearce & Cann, 1973). Calc-alkali basalts plot in fields C, D and G; "within-plate" basalts in fields C and E. Both the Iv and RS rocks form suites which vary in chemical character from abyssal to low-K tholeiitic types. Symbols as for Figure 2,
rocks (Fig. 2). This is also the case for the trace elements Zr, Sr, V, Y, Ce and Nd while Sc is lower in the RS amphibolites (Fig. 10). These slight differences correlate with those found in recent basic rocks from different tectonic settings, i.e. sea floor (RS) and island arc (Iv), although it is emphasized that there is a large overlap in ranges and small variation in chemical differences between the two groups. The distribution patterns of V, Y and Sr are analogous to that of Zr which is plotted against MgO in Figure 10. The slightly higher Zr and Ti contents in the RS amphibolites mean that they are weighted toward the field of abyssal tholeiites proposed by Pearce & Cann (1973) whereas the Iv rocks tend toward the island-arc field although both groups are concentrated where the two fields overlap. Rb was detected in approximately half of the Iv amphibolites with values ranging from 1 to 200 ppm whereas only 5 of the RS rocks have detectable Rb (1, 1, 2, 2, 9 ppm) and as K is marginally higher at equivalent MgO values in the RS rocks, the K/Rb ratios consequently also tend to be higher in this suite. The rare-earth elements (REE) and Y are also slightly higher in the RS amphibolites. The relative distributions of Ce and Y are similar to that of Nd which is plot-
ted against MgO in Figure 10, although La values of the two suites overlap at between 1 and 8 ppm. The average RS rock/chrondrite value for the analysed REE is ca 10 compared to ca 6 for the Iv rocks although, as for the other trace elements, the total range in values is duplicated in the two groups. More detailed, neutron-activation REE analysis is currently under investigation. The higher trace-element concentrations in the RS amphibolites described above all correspond to a more primitive, oceanic-type origin for the Ravns Store* belt and a slight tendency toward island-arc characteristics in the rocks from Ivisartoq (Condie, 1976; Jahn et al., 1974; Jakes & White, 1972; Kay & Hubbard, 1978; Philpotts et at., 1971). However, the differences between the two groups are small and are not considered to be definitive. DISCUSSION The field relationships suggest that the RS originated as a sub-aqueous ultrabasic-basic volcanic pile which formed primary crust, probably in an oceanic environment and that the Iv belt may have been deposited either on or near to a pre-existing continental basement. However, the composition of the two groups is very similar and there is no chemical evidence for the proxi-
GEOCHEMISTRY O F AMPHIBOLITE SUITE, GREENLAND
309
100
75
50
25
0
50
40
30
20
10
15
10
5
0 MgO Fig. 10. Zr. Sc and Nd trace-element chemical-variation diagram presented as ppm versus weight °7o MgO. The behaviour of Sr, V, Y and Ce is analagous to that of Zr and Nd which are higher in the RS than the lv amphibolites at corresponding MgO values, while Sc is relatively higher in the lv rocks. Trend lines separating the two suites are drawn by eye to emphasize the differences in trace-element chemistry. Symbols as for Figure 2.
C. R. L. FRIEND, R. P. H A L L & D. J. H U G H E S
310
mal Iv rocks having been associated with continental material. Both groups form suites which vary in chemical characteristics from komatiites to primitive high- and low-Mg tholeiites. This association is common to many other Archaean metavolcanic belts (Arndt et al., 1977; Nesbitt & Sun, 1976; Williams, 1972). Komatiites have been interpreted as indicating tectonic regimes analogous to island arcs (Brooks & Hart, 1972), sea floor (Gale, 1973; Glikson, 1971), primitive crust (Viljoen & Viljoen, 1969), crustal rifting (Cox, 1978; Nisbet et al,1977) and terrestrial maria (Glikson, 1976; Green, 1972) as well as conditions and igneous processes in the mantle (Cawthorn & Strong, 1974; Cox, 1978). This complexity in interpretation is reflected in the Malene metavolcanics by the fact that those suites remote from and proximal to the older gneissic basement both contain komatiitic fractions, and the feature is not thought to indicate any one tectonic province. Brooks & Hart (1974) have postulated that analogues of island arcs, spreading ridges and local mantle plumes may all accommodate the development of komatiites during the Archaean, assuming the conditional constraints of relatively thin lithospheric plates, a high geothermal gradient and high rate of convection. However, many
authors have argued for a uniformitarian approach in interpreting the Archaean (Moorbath 1975, 1978; Talbot, 1973; Anhaeusser, 1973;' Windley & Smith, 1976). If this approach is adopted, the Malene supracrustal rocks must be regarded as primitive oceanic material which has overlapped onto a pre-existing AmTtsoq gneissic basement in the region around Ivisartoq in a situation analogous to the primitive Tertiary tholeiites which overlap onto the west coast of Greenland at the present day (Clarke, 1970; O'Nions & Clarke, 1972). There are minor chemical grounds for discriminating between the remote RS and proximal Iv amphibolites, but, despite the tendency for the lv rocks to bear certain chemical resemblances to island-arc type tholeiites, there is no definitive evidence for their formative tectonic province being any different to that of the remote primitive oceanic RS amphibolites. ACKNOWLEDGMENTS We are grateful to the Director of the Geological Survey of Greenland for providing facilities to work in Greenland and for permission to publish the results and we thank various colleagues for valuable discussion. C . R . L . F . acknowledges financial support from the Royal Society and Oxford Polytechnic.
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THE GEOCHEMISTRY OF THE EARLIEST KNOWN BASIC METAVOLCANIC ROCKS, AT ISUA, WEST GREENLAND: A PRELIMINARY INVESTIGATION R. C O. Gill, D. Bridgwater, & J. H. Allaart 1
2
3
Department of Geology, Chelsea College, University of London, U.K. Geological Survey of Greenland, Copenhagen, Denmark. Present address: Geological Museum, University of Copenhagen, Denmark Geological Survey of Greenland, Copenhagen, Denmark 1
2
3
ABSTRACT The Isua supracrustal succession, deposited 3.77 x 10 years ago, is the oldest suite of rocks known to have been formed at the Earth's surface. It forms an arcuate raft enveloped by later (3700 Myr) Amitsoq gneiss. The supracrustal rocks have undergone several phases of intense deformation, which have obliterated discordant contacts and primary minerals and textures. The belt includes 25-40% of basic amphibolites of probable igneous parentage, which fall into three categories: (a) homogeneous amphibolites forming layers up to 50 m thick in metasedimentary rocks, (b) amphibolites with garbenschiefer texture which form a 700 m-thick sheet, and (c) banded amphibolites with layering,of uncertain origin. Post-Amitsoq (Ameralik) dykes cutting the supracrustals are also described. The compositions analysed have a scattered distribution in many variation diagrams, but selection of specimens with alkali and volatile contents lower than 2.7 and 3.0% respectively yields consistent distributions which appear to be little affected by alteration. Most samples thus screened fall into two distinct areas, consisting of homogeneous amphibolites (with Ameralik dykes) and garbenschiefer amphibolites (with some banded amphibolites). The homogeneous amphibolites and dykes belong to the Archaean tholeiite suite. The observed variation in composition is quantitatively consistent with two stages of fractionation involving (i) crystallisation of olivine alone, and (ii) crystallisation of plagioclase, olivine and clinopyroxene. There is no evidence of komatiitic units among the homogeneous amphibolites. The garbenschiefer amphibolites are highly aluminous, high-Mg, basaltic rocks (A1 0 15-20%, MgO 8-19%) which are unusually depleted in Ti, Zr, P and Ba. Observed chemical variations in the series may be modelled in terms of the progressive accumulation of olivine (and minor spinel phase) in an aluminous basalt magma (AI2O3 ~ I8V0) having a moderate CaO content. The significance of such an unusually aluminous basaltic magma in an Archaean volcanic environment, and its close association with normal low-K tholeiite at Isua, is discussed. 9
2
3
GEOLOGICAL SETTING surround and intrude the supracrustal rocks The Isua supracrustal succession (Bridgwater & except in the northern part of the area where both McGregor, 1974; Bridgwater et al., 1976, 1978, old gneisses and supracrustals are truncated by a 1979; Allaart, 1976) is the oldest suite of rocks fault and brought into contact with younger known to have been deposited on the surface of gneisses equivalent to the 2800-3100 Ma Nuk the Earth. The rocks have yielded Pb/Pb, U/Pb gneisses of the Godth&bsfjord area. The Amitsoq and Sm/Nd ages of approximately 3770 Ma gneisses and the supracrustals are intruded by (Moorbath et al., 1973; Baadsgaard, 1976; two major groups of basic dykes: (i) Ameralik Michard-Vitrac et al., 1977; Hamilton et al., dykes emplaced in two major sets between 2800 and 3700 Ma; and (ii) at least two sets of Pro1978). The succession forms a raft approximately 35 terozoic dykes emplaced at some time between km long and up to 2.5 km wide folded into a cres-. 1800 and 2500 Ma. Post-Ameralik, precent about a central mass of ca 3700 Ma Amitsoq Proterozoic-dyke metamorphism reached amtonalitic gneiss. Comparable Amitsoq gneisses phibolite facies, and the post-Ameralik-dyke
Spec. Pubis geol. Soc. Aust., 7 (1981)
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R. C. O. GILL, D. BRIDGWATER & J. H. ALLAART
deformation was locally intense and may have been the cause of the present crescent-shaped outcrop of the supracrustals. The supracrustal rocks were already highly deformed before the injection of the Amitsoq gneisses ca 3700 Ma ago. Both the gneisses and the supracrustal rocks were subsequently deformed to such degree that the original discordant contacts are only preserved locally. Deformation is dominated by simple shear, with elongations X:Z seldom less than 1:10 and averaging 1:50 or more, making measurement of strain impossible. This high deformation occurred during one or more periods of thrust-slicing, during which the original supracrustal succession was dissected and locally intercalated with slices of Amitsoq gneiss. Pre-Ameralik-dyke metamorphism reached amphibolite facies with garnet-staurolite assemblages in pelitic units and anthophyllite in basic rocks. The extreme deformation and amphibolite-facies metamorphism virtually destroyed all sedimentary and igneous structures, except for gross lithological layering. Primary igneous textures are locally preserved in the centres of a few dunite pods, and in a few clasts of felsic volcanics surrounded by a less competent matrix. The Ameralik dykes cutting the gneisses within the centre of the supracrustal crescent are virtually undeformed and locally retain original igneous mineralogy. Outside the supracrustal belt the dykes are amphibolites and are cut by pegmatites, but original discordant relations to earlier structures are still preserved. Within the schistose rocks of the supracrustal belt the dykes are commonly highly deformed amphibolites with a marked schistosity and linear fabric identical to that of the rocks they cut. The supracrustal rocks have acted as a zone of high, but variable, strain between two areas of relatively competent gneiss during one or more periods of post-dyke deformation. The low strain within the gneisses enclosed by the belt is probably due to the supracrustals having taken up the regional strain. LITHOLOGIES IN THE SUPRACRUSTAL SUCCESSION General Lithology Chemical sediments, including chert, ironstone, marble and calcsilicate rocks, dominate the metasediments. The main clastic components are pelites and a suite of rather variable rocks ranging from impure calcareous units to quartzfeldspar-biotite rocks which contain clasts of Krich acid volcanic rocks (Bridgwater et al., 1978, analysis 3). There is no evidence of earlier sialic crustal material incorporated within the sedi-
ments. Major layers and pods of dunite, clinopyroxene-peridotite and their talc-schist amphibolitic derivatives occur in the southern and western parts of the belt. The ultrabasic rocks are commonly associated with calc-silicate horizons with which they locally have transitional (possibly metasomatic) contacts. They appear to have been emplaced as tectonic pods, possibly when the postulated thrust-slicing took place, Basic Meta-igneous Rocks Within the Isua Succession Basic rocks of probable igneous derivation form between 25 and 40% of the succession. Because igneous and sedimentary units are often intercalated on a small scale, the map (Fig. 1) shows dominant rock types rather than single homogeneous layers. We distinguish three types of metamorphosed basic igneous material: homogeneous black amphibolites with compositions close to tholeiitic basalts, banded amphibolites with highly variable chemistry, and "garbenschiefer" amphibolites. The last type comprises massive silver-green or pale brown rocks with a marked schistosity, in which there is a posttectonic development of anthophyllite or hornblende in characteristic radiating sheaves. The homogeneous amphibolites form layers from a few centimetres up to tens of metres thick within the metasedimentary units. They are finegrained rocks, homogeneous at hand-specimen scale, and dark green to black, typically consisting of plagioclase (25%) which is commonly zoned, quartz (25%) and hornblende (50%). Diopside is present in a few samples. A slightly "peppered" appearance is common, due to mmsize flecks of feldspar, quartz or scapolite. The hornblende needles show a marked preferential orientation defining the foliation and strong linear fabric of these rocks. Scapolitisation is local and is regarded as a very late feature since it occurs in Ameralik dykes from the same outcrops. Some variation is seen on outcrop scale in the homogeneous amphibolites. Locally, units up to a metre wide develop a schistose parting with biotite along cleavage planes. Layers some 25cm thick consisting almost wholly of amphibole (e.g. 175540) form a small percentage of some outcrops. At the margins of many units there is a gradational contact with the adjacent metasedimentary rocks, where biotite increases and hornblende decreases in a transition zone up to a metre wide. The homogeneous amphibolites may be derived either from flows or from early highlevel intrusive sheets. The amphibole-rich layers could represent mafic cumulates or a less common basic magma type. The marginal biotite-rich
EARLIEST BASIC METAVOLCANICS, ISUA
315
o X -r
Schistosity (vertical, inclined) Fault
\ M D - i Proterozoic dykes 7
^
Ameralik dykes 1
m
Garbenschiefer
pilll metasediment Amphibolite with layers
111111111 Mylonite gneiss
f t t t t i Metasedimentary rocks I : : :l with amphibolite layers
Fig.
H ^ B Talc schist, dunite . . . . . .. Acid metavolcanic rocks a n d associated volcanogenic \ * I i metasediments
r I
1A , J Amltsoq gneiss
~
Veins
Simplified map of Isua area showing main sampling localities.
rocks are interpreted as due to post-magmatic contamination with alkalis derived from the metasediments. The biotite-enriched layers within the amphibolites may be caused by secondary movement of water and alkalis along shear planes. Inhomogeneous banded amphibolites are the most widespread rocks of basic igneous origin within the succession, forming major units in the northeast part of the belt, and layers within the rest of the succession, often side by side with the massive units. The layering is defined by 1 cmwide lamellae of amphibolite alternating with lighter-coloured rock richer in epidote, scapolite, quartz, diopside and plagioclase. There is in places the appearance of grading from the lighter to the darker bands. Garnet occurs locally, some in pockets and augen which have resisted deformation. It is not known how much (if any) of the layering represents primary structures, such as tuffaceous layers or stretched-out pillows, and how much is due to post-crystallisation effects. The garbenschiefer amphibolites form a major unit up to 700 m wide which can be traced for 20 km in the centre of the belt west of the lake Imarssuaq, and on the north side of the belt east
of the lake (Fig. 1). Smaller units of the same rock type are found within the succession east of the lake. In outcrop the garbenschiefer amphibolites represent the largest single unit in the succession. They are generally silverish grey-green. There are local variations in the size, colour and distribution of the amphibole porphyroblasts, which are set in a chlorite-plagioclase-quartz matrix. The mapped unit is locally discordant to structures within the layered metasedimentary succession, possibly because of intrusive origin. It encloses small lacunae of metasedimentary rock which appear less deformed than those in the surrounding parts of the succession. There are gradational contacts between typical garbenschiefer amphibolite and thin units of garnetbiotite schist which may represent metasedimentary inclusions or be the result of metasomatic activity. The garbenschiefer amphibolites are interpreted as derived from one or more large sheets which were intruded into the succession, but which have since been highly modified by extreme deformation (possibly controlled by their chloritic nature). Many of the present contacts may be tectonic. Chemically the garbenschiefer unit is far
316
R. C. O. GILL, D. BRIDGWATER & J. H. ALLAART
less homogeneous than its field appearance suggests, having provided a sufficient spread in S m / N d ratios to allow its use in age determinations (Hamilton etal., 1978). SAMPLING CONSIDERATIONS AND ANALYTICAL METHODS The conclusions of geochemical studies on highly deformed rocks in metamorphic terrains can only be as reliable as the field interpretation upon which such work is based. Accordingly we have been selective in the choice of material for the present preliminary study, confining attention mainly to those units and localities in which original relationships appear to have been least disturbed. We have drawn on material collected by Bridgwater and McGregor in 1973, by Allaart in 1974 and 1975, and by Bridgwater in 1978. The earlier sampling sites were as far as possible revisited, and the field relations re-examined, in 1978 (Bridgwater et al., 1979). Sampling of the homogeneous amphibolite units has been concentrated on outcrops where their relationship to surrounding units is well exposed, and where in particular it is possible to identify and distinguish the later Ameralik dykes, with which they are often easily confused when deformed (c/. Appel, 1977, analyses 1906 and 2201). Biotite-rich rocks have been avoided. A number of Ameralik dyke samples cutting the supracrustals have been included in this work for comparison. The garbenschiefer amphibolites have been interpreted in the literature as comprising one or more intrusive sheets. We have therefore included this unit in the present work, though recognising that some apparent discordances with surrounding metasediments may be of tectonic origin. Less attention is given to banded rocks in the present woik owing to the difficulty of field interpretation. The banded amphibolites probably have a variety of derivations. M a j or-element contents were determined by Xray fluorescence spectrometry on borate fusion beads, either at the University of Leeds or at Greenland Geological Survey facilities in Copenhagen. In the latter case N a 2 0 and MgO have been determined by atomic absorption or related methods. Trace-element concentrations were measured by X R F at the Universities of Leeds and Nottingham. Most rare-earth elements have been analysed by instrumental neutron activation and a few by isotope dilution. Tables of analyses are available f r o m the Librarian, Geological Survey of Greenland Ostervoldgade 10, Dk-1350, Copenhagen K, Denmark. In the diagrams given in this paper, the
function 1 0 0 M g / ( M g + Fe*) is used as abscissa, where Fe* = 0.865 x L F e . G E N E R A L G E O C H E M I S T R Y AND THE QUESTION OF ALTERATION All but a few of the rocks considered here, judging f r o m S i 0 2 and MgO contents, fall in the compositional fields of basalt and high-magnesium basalt. Within this broad description, the compositions vary widely, particularly with regard to A1 2 0 3 , M g O , C a O and alkali content, and assignment to recognised basalt classes does not appear to be a straightforward task. In this context, it is important to examine the contribution of post-magmatic processes to the chemical variability seen a m o n g the. Isua metavolcanic rocks, all of which have suffered complete recrystallisation, possibly more than once. In the absence of igneous minerals, textures and structures, estimation of the degree of alteration must rest solely on chemical, rather than petrographic or structural, criteria.
Na 2 0%
Fig. 2.
K 2 0 versus Na 2 0 content (weight percent) of the Isua metavolcanics. Symbols used are described in Table L The curves labelled 1.0, etc., are contours of constant total alkali oxide content, at the percentage levels indicated.
EARLIEST BASIC METAVOLCANICS, ISUA
317
TABLE I
Subdivision by chemical alteration: symbols used in figures Increasing c h e m i c a l D e g r e e of "hydration"
Type
Homogeneous ar.iph i bol i te
(HA)
H 2 0 + C 0 2 < 3.0
Banded a m p h i b o 1ite
H 2 0 + C02
<3-0
Garbenschiefer CGS)
H 2 0 + C02
< 3.0
(AD)
Dykes cutting g n e i s s
B
A
A A
0
H 2 0 + C 0 2 < 3.0
o ®
H 2 0 + C02
> 8.0
Na20 + K20
•
A
H 2 0 + C 0 2 ^ 3.0
alter 2.1%
K 2 0 < 0.5
H 0 + co2 ^3.0 2
Dykes cutting supracr-ustals
Na20 + K20 <
>2.1%
k2O
0.5
A
K 2 O ^ 1.0
A
© ©
(D)
H 2 0 + C 0 2 < 3.0
(AD)
H 2 0 + C02
^ 3.0
4
mobility of other elements. Recent low-K and high-Mg basalts, when fresh, usually have low N a 0 and K 0 contents (less than 2.5 and 0.25% respectively) and the same seems to be true of comparable Archaean rocks. On the assumption that the higher alkali contents in a number of the Isua volcanics reflect the secondary influx of alkalis whose extent may be related to the movement of other elements, we have established three arbitrary categories of progressive chemical alteration based on alkali content, and one based on volatile content, as shown in Figures 2 and 3 and Table I. The symbol for each composition in the variation diagrams has been ornamented according to its alteration category (see Table I). Figure 4 illustrates the value of making these distinctions. T i 0 concentrations are scattered, particularly for the garbenschiefer, to an extent which obscures the influence of igneous processes. However, when attention is confined to © ® A ® the "least altered" samples (open symbols) more consistent groupings emerge, revealing plausible A © A igneous trends as shown by the fields drawn for A least altered samples in Figure 4. Altered samples with the highest K 0 contents deviate furthest from these fields, whereas rocks with less extreme A 0-01alkali enrichment lie closer to, or inside, them. The introduction of volatiles alone (symbols with vertical ornament only) seems to have had little effect on the compositions of the garbenschiefer (at least with regard to Ti, Mg and Fe), but this does not appear to be true for some of the dykes. With certain exceptions discussed below, these observations apply to variation diagrams of Fig. 3. K 0 content versus total volatile content. many other elements, in which the more severely Horizontal and vertical lines denote the per- altered samples consistently occupy anomalous centage levels used to discriminate between positions (Figs 5, 6, 7, 8 and 9). The manner in degrees of alteration.
Na 0, K 0 , H 0 and C 0 contents are extremely variable (Figs 2 and 3), and there is little prospect that their original concentrations have been preserved quantitatively in any of the Isua volcanics. Consequently the classification of these rocks using normative criteria would be inappropriate. Alkali metal and volatile contents provide a means of discriminating between more and less altered specimens, and of estimating the 2
2
2
2
2
2
10-0
k 2 o %
2
A
2
•
2
Homogeneous amphibolrte
/ \
Garbenschiefer
O
Ameralfc dyke
318
R. C. O. G I L L , D. B R I D G W A T E R & J. H. A L L A A R T
100 M g (Mg+Fe*)
Fig. 4.
Ti0 2 contents of the Isua metavolcanics plotted against Mg number, using the symbols and abbreviations given in Table I. Fields are drawn around the least altered samples (open symbols) only, omitting dyke sample 158543 and garbenschiefer sample 158475, both of which are anomalous. Arrows indicate the predicted evolution of Ti0 2 content, according to the fractionation calculations based on major elements, and assuming Ti to behave incompatibly.
T ~
i
I
A
/ /A
* > -GS 4y
^Af /
i
T
•
A A
S
y
/ /y|
P-Jt^mOn
o
D
. / /
ha+ad
O 158543 P ,175537,41
O—175540 • L an Fig. 5.
' 100 Mg/(Mg+Fe*)
I
Al 2 0 3 contents of the Isua metavolcanics versus Mg number, using the symbols given in Table I. Open arrows indicate the changes in composition implied by the least-squares fractionation calculations: the vertical " w i d t h " of the trailing end of each arrow shows the difference between observed and least-squarescalculated concentrations. Abbreviations as in Figure 4.
which these apparently substantial changes occur is not generally clear, although there is evidence in a few instances (e.g. 171757) for the migration of magnesium. On the other hand, the coherent behaviour of the open-symbol compositions in these diagrams confirms our belief that specimens with the lowest alkali contents, particularly of potassium, are the least altered specimens in the widest sense. That these samples conform to the predicted igneous behaviour of the elements concerned (discussed in the next section) suggests that in many respects they approximate to their original igneous compositions. Some elements show systematic changes of concentration even at less severe levels of alkali introduction, suggesting a more general mobility. This is so for CaO (shown as C a 0 / A l 2 0 3 in Fig. 6). Among the homogeneous amphibolites and dykes there is a consistent fall in C a 0 / A l 2 0 3 with increasing alkali content, a pattern also shown, though less clearly, by the garbenschiefer unit. In the latter, CaO also seems to have responded to the introduction of volatiles. Si0 2 contents show a similar but less clear tendency to increase with alkali alteration. One cannot express much confidence that CaO and S i 0 2 remain completely undisturbed by alteration in even the least altered compositions. The same of course applies to the alkali metals and volatiles themselves. In subsequent sections we confine attention to the least altered members of each rock type
E A R L I E S T B A S I C M E T A V O L C A N ICS, I S U A
319
TABLE II
Samples*showing unrepresentative behaviour Type
Sample numbers
Homogeneous amphi boli te
l75537 1755*0 175541
Low AI2O3 High incompatible element concentrations: Ti ~ 1.2 x rest Zr ~ 1,7 x rest La ~ 2.0 x rest
158403
Low Y, Zr High Ca, Cr
Slight accumulation of cpx?
Ameralrk dyke
158543
Very low Si, ti, Zr, Y, A1 Very high Ni, Cr
Accumulation of olivine and spinel?
Garbenschiefer
175550 175551 175559 175560
Low Si, A], Zr High Ca, Cr, Ni
Slight accumulation of cpx + olivine?
Differences from other members of group
1 nterpretat ion
-
Only "least altered" compositions are considered.
(represented by open symbols), believing these to be close to their primary compositions for many elements (though less close for Ca, Si and the alkali metals). We have included the mostmagnesian members of the garbenschiefer series in this class in spite of their high volatile contents, because no better preserved representatives of this composition are available, and their behaviour appears to be consistent with other members of the series. The primary geochemistry of the homogeneous amphibolites appears to be identical to the later
Ameralik dykes, of which we now recognise a wider range of composition than previously reported from Isua (Gill & Bridgwater, 1979). Most rocks in these two groups appear to belong to a single line of descent (Figs 4, 5, 7, 8 and 9). A small number of samples (among those having low alkali and volatile contents) do however exhibit divergent behaviour to varying degrees, as outlined in Table II. Such departures might have arisen from incipient alteration, but we believe they are more likely to be primary features in view of the consistency seen between petrogeneti-
L e a s t altered H A
CaO/AI203
Altered H A
L e a s t altered G S
80
70
100 M g / ( M g + F e ' ) Volatile-rich altered G S
ANcali-rich altered G S a n d B A
Fig. 6. C a 0 / A I 0 ratios plotted against Mg number. Fields are drawn to show the dependence on degree of alteration, as inferred from Figures 2 and 3. Symbols as in Table I. 2
3
320
R. C . O . G I L L , D. B R I D G W A T E R & J . H . A L L A A R T
cally related elements. Some compositions noted in Table II are consistent with the slight accumulation of phenocrysts in otherwise unmodified magma. The garbenschiefer amphibolites have very distinctive compositions quite different from the homogeneous amphibolites. They appear to have been highly aluminous basic rocks (Fig. 5), very depleted in incompatible elements such as Ti, Zr, P and Ba( Figs 4 and 9) compared to the homogeneous amphibolites, but also with somewhat lower contents of Ni and Cr (Figs 7 and 8), these comparisons being made at a fixed Mg number. The aluminous nature of these rocks is unusual in the Archaean context, and is discussed later. A few members again have compositions consistent with accumulation of phenocrysts (Table II). The banded amphibolites have variable compositions. One member (175513) lies in the garbenschiefer amphibolite field in many variation diagrams, but the other samples are compositionally more isolated. Their compositions (high Ni and Cr, low Ti etc.,) could be derived from the garbenschiefer type by the secondary loss of major amounts of magnesium, but there is no positive evidence for this interpretation.
h o m o e e n L o u f amnhfhnv t tenTs Arrow , E t th
g
" T ^ " ^ C h
omen C
a f t ^ n ^
IGNEOUS FRACTIONATION In spite of the locally severe effects of alteration, deformation and metamorphism evident in the Isua succession, the geochemistry appears to retain a record of igneous processes. We have conducted least-squares mixing calculations to formulate and test the details of these processes semi-quantitatively. The homogeneous amphibolites and dykes cover a wide interval of Mg number, which we have attempted to model in two arbitrary stages: (1) from Mg number 70 to 60 and (2) from 60 to 53. A third stage would be required to produce the most evolved compositions seen at Isua, but such a calculation has been published elsewhere (Gill & Bridgwater, 1979) and need not be duplicated here. To minimise the disruptive effects of alteration on individual samples, liquid compositions have been derived by averaging groups of three rock analyses, each group representing either the beginning or end of one of the fractionation stages. The calculations assume reasonable "phenocryst" compositions, as described previously (Gill & Bridgwater, 1979). The first stage of differentiation, in which A1203 rises (Fig. 5), is dominated by the crystallisation of
F l d d s are d r a w n
for ^ least altered garbenschiefer, least altered and f r ° ^ - a l k a l i h o m o g e n e o u s amphibolites with low Ni c o ^ r g e 0 f N i c o n t e n t w i t h crystallisation as determined by the least° f t h e o l i ™ e / l i q u i d partition coefficient (based on melt MgO Rayleigh f r a c t i o n a t i o n equation: garbenschiefer^;
EARLIEST
BASIC METAVOLCANICS,
olivine (11%), with only trivial amounts of clinopyroxene and plagioclase involved. In the secondstage solution, the major influence is plagioclase (10%), accompanied by small amounts of olivine (4%) and Ca-rich clinopyroxene (3%). The result is a slight reduction of A1 0 content. The calculations predict the behaviour of relevant trace elements with sufficient accuracy to suggest that the observed geochemical trends are of igneous origin. Ni content falls at a rate consistent with the crystallisation of olivine, although Figure 7 suggests that the stage 2 calculation has slightly exaggerated the amount involved. The predicted rise in incompatible element contents (Figs 4 and 9) agrees with observation, and the same applies to La and V (not plotted). We have not attempted to use Sr or Eu behaviour to estimate plagioclase fractionation, for reasons discussed elsewhere (Gill & Bridgwater, 1979). The calculated fractionation scheme, which parallels that seen in recent ocean-floor basalts, offers no explanation for the anomalous behaviour of specimens 175537, 175540 and 175541 (Table II), which perhaps represent a separate magma succession. Their higher Ca0/Al 0 ratios (Fig. 6) suggest an affinity with basaltic komatiite, a contention which is not however supported by their higher incompatible-element contents (Table II; cf. Arndt et al., 1977). Similar calculations show that the garbenschiefer series could be generated by the crystallisation of olivine from a melt corresponding to the most magnesian members of the series (Table III). In common with the homogeneous amphibolite calculations the residual errors obtained are somewhat larger than might be expected, especially for Si0 (0.9%), A1 0 , CaO and Na 0, but given the appreciable mobility of some of these elements, and the uncertainty in mineral compositions, they are not unacceptably large. A model with these parameters (Table III) predicts the observed trace-element behaviour with acceptable accuracy (Figs 4, 5, 7 and 9), confirming that the essential igneous geochemistry of the least altered specimens of garbenschiefer amphibolite has not been seriously distorted by alteration. The principal difficulty in ascribing the observed chemical variation among the garbenschiefer to fractional crystallisation is to explain the composition of the supposed parent liquid, which is exceptionally aluminous (A1 G -15%) for such a magnesian composition (MgO ~ 18%, Mg number -80). Such a magma would undoubtedly lie beyond the present compositional limits of known natural melts. A more acceptable interpretation is that the garbenschiefer amphibolites represent melt compositions that have been modified by crystal accu2
2
2
TABLE III
A fractional crystallization model for the garbenschiefer amphibolites "Parent liquid" composition
3
2
3
2
2
3
321
ISUA
I"Res idual | liquid" compos i t ion
Crystallizing minerals
Observed 1
Calculated 2
Olivine 3
Magnetite
Plagioclase''
Observed 5
Si0 2
49.22
50.02
41.9
-
47.20
51.32
A1203
15.98
15.08
-
-
33.05
18.09
Fe203
1.40
1.42
-
69.0
0.82
1.48
FeO
8.68
8.57
7.50
31-0
-
8.60
MnO
0.22
0.31
-
-
0.03
0.25
MgO
10.16
17.55
17.52
51.21
-
-
CaO
6.59
7.23
-
-
17.06
Na20
0.67
1.05
-
-
1.78
1.26
K20
0.08
0.15
-
0.05
0.18
0. IS
82.}%
Solution 6
17.5%
0.3%
8.67
1
Hypothetical parent -1iquid. Mean of analyses 171756, 171758, 175562 and 175563 recalculated volatile-free. Oxidation ratio has been adjusted to a standard value (FeO = 0.865 x E FeO).
2
Calculated by least-squares mixing of crystallizing minerals and residual composition, in proportions shown insolution.
3
Calculated assuming Fe/Mg in equi1ibrium with parent liquid (after Roeder & Ems lie, 1970).
**
Bytownite analysis from Deer, Howie & Zussman Table 17).
5
Hypothetical residual liquid. Mean of analyses 158483, 175554 and 175557 recalculated volatile-free. Oxidation ratio has been adjusted to a standard value (see note x ).
6
Proportions in which residual liquid and phenocryst phases must be combined to regenerate the supposed parent 1iquid.
3
(1963,
mulation. One possibility is that the high A1C>3 contents reflect the addition of plagioclase, but this hypothesis is inconsistent with the low observed CaO concentrations and negative Eu anomalies. Plutonic rocks enriched in basic plagioclase generally have CaO contents well in excess of 16% and positive Eu anomalies. Calculations show (Table IV) that the garbenschiefer series may be reproduced by adding olivine to a liquid having the composition of one of the Mg-poor members of the series. The olivine composition is assumed to be in equilibrium with such a liquid. It is therefore more Fe-rich than the olivine whose removal was considered in the fractional crystallisation model, and a correspondingly greater proportion (25%) must be added to generate the most magnesian compositions. Table IV shows that an acceptable match with the average of magnesian members (cf. Table III) may be obtained, though the error in Al 03 is marginally greater. Although the amount of olivine involved is greater than in Table III, the change in Ni content is approximately the same, because it is controlled by a linear rather than non-linear (Rayleigh) process. The observed behaviour of the incompatible elements (e.g. Ti0 in Table IV) is found to be consistent with either of the two models. That of Cr 2
2
2
R. C. O. GILL, D. BRIDGWATER & J. H. ALLAART
322
suggests that a minor amount of an oxide phase such as magnetite must be involved in addition to olivine. Two conclusions may be drawn regarding the garbenschiefer amphibolites. In the first place, their internal geochemical variation may be explained quantitatively in terms of either of two igneous processes. It would be highly fortuitous for such consistent relationships to arise spuriously from non-igneous processes (sedimentary, metasomatic), and we judge their contribution to be insignificant, except for elements of established mobility. Secondly, if our olivine-accumulation model is correct, the melt in which the olivine has accumulated (the 'host melt'', Table IV) evidently had a composition much more aluminous than basaltic rocks commonly associated with Archaean volcanic environments. 4
TABLE I V
A linear mixing model for the garbenschiefer series Host melt c o m p o s ition Si02
51.3
A1203
18.1
0 1 i v i n e in equi1ibrium3
ko.k -
Host melt plus 25% olivine
Average of four magnesian members
49.1
49.2
14.5
15.7
Fe203
1.5
-
1.61
1.4
FeO
8.6
12.1
9.01
8.7
MnO
0.25
0.1
0.2
0.22
MgO
10.2
47.5
17.6
17.6
CaO
8.7
-
6.9
6.6
Na20
1.3
-
1.0
0.7
K20
0.18
Ti02 Ni
0.31 (ppm)
181.
-
2170.
0.14 0.25 580.
0.08 0.22 534.
1
Adjusted
2
cf.
3
Calculated from column 1 using distribution constant of R o e d e r 6 Ems lie ( 1 9 7 0 ) .
Table
to s t a n d a r d o x i d a t i o n
ratio.
III.
MAGMATIC AFFINITY AND SIGNIFICANCE The igneous geochemistry of the homogeneous amphibolites, which occur in a volcanic setting within metasedimentary units at Isua, is indistinguishable from that of the later intrusive phase of magmatism represented by the widespread Ameralik dykes. This identity of character extends to the rare-earth elements (Fig. 10; cf. Gill & Bridgwater, 1979). Thus magmas of identical composition have been produced at one locality on two different occasions, before and after the plutonic magmatism represented by the AmTtsoq gneisses. The Isua homogeneous amphibolites, like the Ameralik dykes, are members of the Archaean low-K tholeiite type, and share its chemical distinctiveness (Gill, 1979; Gill & Bridgwater, 1979). Although the Isua meta-
Fig. 8. Cr contents (ppm) versus Mg number. Fields are drawn around least altered garbenschiefer (but note anomalous groups) and for homogeneous amphibolites and dykes, omitting the samples detailed in Table II.
volcanic occurrence is significantly older than other greenstone belts, there appears to be no evidence for age-related differences in composition, except that the Isua metabasalts and dykes have slightly lower Ti, Zr and P contents than many Archaean tholeiites (cf. Gill, 1979). Archaean tholeiites are usually the most abundant basic magma type in greenstone belts. This does not appear to be true of Isua, as now exposed, where they appear to be subordinate in volume to the amphibolites of garbenschiefer type. The geochemistry of this unit seems to be unique, combining highly aluminous and magnesian tendencies. The rare-earth patterns are similarly unusual (Fig. 11), but preliminary determinations by INAA methods suggest that the samples shown may not be typical of the least altered category. The most plausible igneous interpretation of the geochemistry of the garbenschiefer amphibolites views them as a series of magma (not liquid) compositions produced by accumulation of olivine (with minor magnetite) in a host melt of the composition shown in Table IV, column 1. This composition is noteworthy for its very high A1 0 content and low CaO content. Such magmas, hitherto unreported in Archaean volcanic successions, are today associated with island-arc environments. It is perhaps significant 2
3
EARLIEST BASIC METAVOLCANICS, ISUA
323
100 175537 175541 0 •
o
175540
50-
Zrp.p.m.
Y P'grfll. 20-
10*
80
"to
~eS
50"
40
100 Mg/(Mg+Fe ) #
9. Zr and Y contents (ppm) versus Mg number. The fields are drawn around the least altered specimens, omitting those specified in Table II. The arrows show the calculated evolution of Zr content, as described in the caption to Figures 4, 7 and 8.
R. C. O. GILL, D. BRIDGWATER & J. H. ALLAART
324
to some common value of Mg number or MgO content). Reference to Figures 4, 5, 7, 8 and 9 shows that the same differences exist between the homogeneous amphibolites at Isua (or the Archaean tholeiite association in general) and the garbenschiefer "host melt" composition. It i now generally accepted that absolute geochemical comparisons between Archaean and recent basalt geochemistry provide no indication of the tectonic environment in which Archaean volcanism occurred. It is however possible that relative differences among Archaean volcanics may have significance in relation to the environment of melting and perhaps eruption. The garbenschiefer amphibolites possibly represent melting in mantle conditions similar to those in which island-arc magmas are generated today, but there are obvious discrepancies with regard to REE and Mg/Fe values which indicate that this is not an exact analogy, and discussion in terms of presentday subduction processes would be unwarranted. s
Fig. 10. Abundances of rare-earth elements for the homogeneous amphibolites. Analyses by instrumental neutron activation methods at the University of London Reactor Centre.
that picritic basalts are not uncommon in such associations (Carmichael et al., 1974), indicating the occurrence of olivine accumulation of the kind we infer for the Isua garbenschiefer amphibolites, although modern picrite basalts are rarely as aluminous. Tholeiites in present-day island arcs, though broadly similar to mid-ocean ridge basalts, have certain differences in detail. They are generally more aluminous, and are impoverished in Ni, Cr, Ti, Zr, Nb and P (the comparison being referred
ACKNOWLEDGMENTS R. G. thanks Drs P. N. Taylor and N. W. Rogers for guidance and help with REE determinations, and Dr S. Moorbath for permission to use the facilities of the Age and Isotope Laboratory at the University of Oxford. Dr B. E. Gorman and Mr R. Hardy have provided access to computer programmes. R. G. gratefully acknowledges the award of a travel grant by The Royal Society, without which attendance at the Perth symposium would not have been possible. This paper is published with the permission of the Director of the Geological Survey of Greenland.
15-
10 -
"Oc rO0 5 o0 DC
Ce
Nd
Sm Eu Gd
Dy
Er
Yb
Fig. 11. Abundances of rare-earth elements for three garbenschiefer samples 171755, 1 7 1 7 5 7 , 171759. Analysis by isotope dilution at University of Oxford. Figures beside the specimen numbers denote T i 0 contents (%). 2
EARLIEST BASIC METAVOLCANICS, ISUA
REFERENCES
ALLAART, J. H., 1976: The pre-3760m.y. old supra-
crustal rocks of the Isua area, central West Greenland, and the associated occurrence of quartzbanded ironstone; in Windley, B. F. (Ed.) The
Early History of the Earth, 177-189. Wiley, London.
APPEL, P. W. U.,
1977: Aeolian differentiation of basaltic tuffs in the early Precambrian Isua supracrustal belt, West Greenland. Neues Jb. Miner.
Mh1977, 521-528.
ARNDT, N. T . , NALDRETT, A . J . , & PYKE, D. R., 1977:
Komatiitic and iron-rich tholeiitic lavas of Munro Township, Northeast Ontario. J. Petrol., 18, 319-
369.
BAADSGAARD, H., 1976: Further U-Pb dates on zircons
from the early Precambrian rocks of the Godthaabsfjord area, West Greenland. Earth planet.
Sci. Lett33, 261-267.
BRIDGWATER, D . , ALLAART, J . H . , BAADSGAARD, H . , COLLERSON, K. D . , ERMANOVICS, I., GORMAN, B. E., GRIFFIN, W . , HANSON, G . , MCGREGOR, V. R . , MOORBATH, S., NUTMAN, A . P . , TAYLOR, P . , TVETEN, 1., & WATSON, J., 1979: International
field work on Archaean gneisses in the Godthaabsfjord—Isua area, southern West Greenland. Rep.
geol. Surv. Greenland, 95, 66-71.
BRIDGWATER, D . , COLLERSON, K. D . , & MYERS, J. S.,
1978: The development of the Archaean gneiss complex of the North Atlantic region; in Tarling,
0. H. (Ed.) Evolution of the Earth's Crust, 18-69.
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BRIDGWATER, D., & MCGREGOR, V. R., 1974: Field
work on the very early Precambrian rocks of the Isua area, southern West Greenland. Rep. geol.
Surv. Greenland, 65, 49-54.
CARMICHAEL, I. S. E., TURNER, F . J., & VERHOOGEN,
J., 1974: Igneous Petrology. McGraw-Hill, New
York.
DEER, W . A . , HOWIE, R. A . , & ZUSSMAN, J., 1963:
Rock Forming Minerals, 4. Longmans, London.
GILL, R. C. O., 1979: Comparative pedogenesis of Archaean and modern low-K tholeiites. A critical review of some geochemical aspects; in Ahrens,
L. H. (Ed.) Origin and Distribution of the
Elements, 2, 431-447. Pergamon, Oxford. 1979: Early Archaean basic magmatism in West Greenland: the geochemistry of the Ameralik dykes. J. Petrol., 20, 695-726.
GILL, R. C. O., & BRIDGWATER, D.,
HAMILTON, P . J., O'NIONS, R. K., EVENSEN, N . M . , BRIDGWATER, D., & ALLAART, J. H., 1978: Sm-Nd
isotopic investigations of Isua supracrustals and implications for mantle evolution. Nature, Lond., 272, 41-43. IRVING, A. J., 1978: A review of experimental studies of crystal/liquid trace element partitioning. Geochim.
cosmochim. Acta, 42, 743-770.
MICHARD-VITRAC, A . , LANCELOT, J., ALLEGRE, C. J . , &
MOORBATH, S., 1977: U-Pb ages on single zircons from the early Precambrian rocks of West Greenland and the Minnesota River Valley. Earth planet.
Sci. Lett., 35, 449-453.
Academic Press, London.
MOORBATH, S., O'NIONS, R. K., & PANKHURST, R. J.,
MYERS, J. S., 1976: Archaean gneiss complex of Greenland; in Escher, A., & Watt, W. S. (Eds) Geology of Greenland, 18-75. Geol. Surv. Greenland, Copenhagen.
ROEDER, P. L., & EMSLIE, R. F., 1970: Olivine-liquid
BRIDGWATER, D . , KETO, L . , MCGREGOR, V. R., &
1973: Early Archaean age for the Isua iron formation, West Greenland. Nature, Lond., 245, 138-
139.
equilibrium. Contr. Mineral. Petrol., 29, 275-289.
ARCHAEAN GRANITOIDS AND GNEISSES
THE ANCIENT SAND RIVER GNEISSES, LIMPOPO MOBILE BELT, SOUTH AFRICA R. E. P. Fripp Department of Geology, University of the Witwatersrand, 1 Jan Smuts Avenue, Johannesburg 2001, South Africa Present Address: International Nickel Australia Ltd, 7205 Hay St., West Western
Australia
Perth,
6005
ABSTRACT The Sand River Gneisses, which are at least ~ 3800 Ma old, occur within the Central Zone of the Limpopo Mobile Belt, and are particularly well exposed near Messina in the northern Transvaal, South Africa. They form the main component of a high-grade migmatitic complex which has been studied in particular detail within an area of ~ 2 k m 2 . The Sand River Gneisses consist of two components, a mesocratic grey quartz dioritic gneiss ("grey gneiss") and a leucocratic granodioritic gneiss ("leucogneiss") which are interlayered as units commonly a metre or less in thickness, and rarely up to 10m thick. Together these define a - 3 7 9 0 M a Rb-Sr whole-rock isochron, which is regarded as the age of a tectono-metamorphic mineral fabric (including hypersthene) superimposed on the Gneisses. Field relationships indicate that the Sand River Gneisses were intruded at least five times by dykes of various compositions and ages. Two of these are pyroxene amphibolites of tholeiitic composition and they cut both the mineral fabric and layering of the Gneisses. The older of the two, called the Causeway Dykes, have been dated at - 3 5 7 0 Ma, and are cut by the younger, called the Stockford Dykes, dated at ~ 3070 Ma. The latter contain fragments of anorthosite in places, clearly indicating that they post-date anorthositic gneiss which occurs as trains of boudins in the Sand River Gneisses, and which presumably represent genetic equivalents (feeders?) of the Messina Layered Intrusion. Including the early tectono-metamorphic fabric, the Sand River Gneisses have urfdergone six distinguishable deformation events, two of which, the second and last, involved the formation of small-scale shear zones, and are considered to indicate periods of basement uplift. The adjacent high-grade orthogneisses and paragneisses which comprise the "supracrustal" gneisses in the Central Zone of the Limpopo Mobile Belt, and are collectively called the Beit Bridge gneisses, have undergone the last four deformation events and d o not contain the Causeway Dykes, or the second deformation shear zones. Therefore, the Sand River Gneisses underwent a period of uplift before, or during, the development of the Beit Bridge gneisses, and it is possible that the Sand River Gneisses formed a true basement to the latter, so that the Beit Bridge gneisses are possibly supracrustal in the strict sense.
INTRODUCTION Previous Work The Limpopo Mobile Belt is an approximately linear, ENE-trending, distinctive tectono-metamorphic province which separates the Kaapvaal and Rhodesian cratons and is up to 320 km wide (Fig. 1, inset). In contrast to the low-grade, mainly greenschist facies of metamorphism associated with the granite-greenstone terrains of the adjacent cratons, the Limpopo Belt is characterized by high-grade, mainly granulite-facies gneisses in which hypersthene and/or ortho-
amphibole are present (Macgregor, 1947; du Toit & van Reenen, 1977). A distinctive 4 'grey granitic gneiss" within the Central Zone (Cox et al., 1965; Mason, 1973; Fig. 1) of the Limpopo Belt, was first described by Sohnge (1946). He noted that the "gneiss" is locally migmatitic, and contains folded pegmatite dykes and layers of intensely deformed "amphibolite". Subsequently the amphibolite layers were shown by Bahnemann (1971) to be deformed and metamorphosed mafic dykes within a predominantly "hypersthene tonalite gneiss". Two distinct ages* for these dykes have since
* All ages unless otherwise stated, are Rb-Sr whole-rock isochrons, using the decay constant for 8 7 Rb 10- n y r - i .
Spec. Pubis geol. Soc. Aust., 7 (1981)
= 1.42 x
330
Fig. 1.
R. E. P. FRIPP
Inset: the general structure of the Limpopo Mobile Belt is that of a Central Zone (CZ) flanked by Northern and Southern Marginal Zones (NMZ, SMZ). Locality of the main figure is annotated. Main Figure: the general structure of part of the Central Zone showing the area of Figure 2 southeast of Messina (M).
been recognized, ~3570Maand ~3070Ma (Barton et al., 1977). The gneisses into which these dykes were intruded have been dated at ~ 3790 Ma (Barton et al., 1978) and have been referred to as the Sand River Gneisses because they are particularly well exposed as pavements in the bed of the Sand River. In this paper the general nature of the ancient migmatite complex is summarized, and the tectonic and metamorphic history of the various components, their field relations and their compositions are outlined. Geological Setting The Sand River migmatite complex, as it is currently understood and mapped, covers an area of about 40km2 around the Sand River (Fig. 2). Here, these distinctive gneisses are tightly infolded and possibly in-thrust with the Beit Bridge gneisses. The informal term 4'Beit Bridge gneisses" is given to a suite of high-grade, polydeformed paragneisses and orthogneisses composed mainly of quartzite, pyroxene amphibolite, garnet-sillimanite-cordierite-biotite gneiss (meta-
pelite), and various leucocratic, commonly garnetiferous and biotitic quartzo-feldspathic gneisses, together with lesser amounts of banded magnetite quartzite, calc-silicate gneiss and marble. Also intimately associated, infolded and metamorphosed with all these gneisses are the anorthositic, gabbroic and peridotitic gneisses of the Messina Layered Intrusion, which has been dated a t - 3 1 5 0 Ma (Barton et al., 1919a) and may be as old as ~ 3350Ma (Barton, 1979). These intrude both the Sand River Gneisses and the Beit Bridge gneisses and have a similar tectonic history to the latter. None of the older ( - 3 5 7 0 M a ) mafic dykes has been discovered in the Beit Bridge gneisses, which are therefore apparently between about 3570 and 3350 Ma in age, The contact between the Sand River migmatites and the Beit Bridge gneisses is exposed at several localities, the most widely known (Bahnemann, 1971) and most intensely studied being in the bed of the Sand River about 10 km southeast of Messina, some 200 m south of the MessinaTshipise road bridge (Fig. 2). Sand River migmatites are also excellently exposed in the river bed
S A N D RIVER GNEISSES, LIMPOPO MOBILE BELT
331
EXPLANATION FAULT BANDED
—' MAGNETITE
[';'-•;•;"} MESSINA BEIT BRIDGE ~
] SAND RIVER
LAYERED
QUARTZITE INTRUSION
GNEISSES MIGMATITE
ROAD FARM
TRACK
©
CAUSEWAY
LOCALITY
©
STOCKFORD
LOCALITY
COMPLEX
Fig. 2. Generalized geology of the Sand River area (mainly after Fripp and partly P. C. Horrocks), showing main lithostratigraphic units and the two main pavement exposures of the Sand River migmatite, in the bed of that river. Locality 1 is detailed in Figure 3.
about 1.5 km to the southwest. The two localities are known as the Causeway and Stockford localities respectively. At the Causeway locality (Fig. 2) the rock exposures have been mapped at a scale of 1:1000, and in part at 1:100, and generalized maps are reproduced here (Figs 3 and 4). The mapping, together with detailed observations at the Stockford locality (Fig. 2), has shown that the Sand River migmatite is made up mainly of mesocratic, grey, biotitic granitoid gneiss, and leucocratic granitoid gneiss, loosely referred to as the "grey gneiss" and the "leucogneiss" respectively. Together they define the— 3790 Ma isochron and are formally referred to as the Sand River Gneisses (Barton etal., 1977, 1978). The grey gneiss consists mainly of plagioclase, quartz and biotite, with minor hypersthene, hornblende and opaque minerals (Table I, column 1). Potassium feldspar is notably absent and the potash revealed by geochemical analysis is therefore mainly in the biotite. The average majorelement chemistry of the grey gneiss is very similar to the average quartz-diorite quoted (Table I, column 2) so that the gneiss is, both geochemically and mineralogically, a quartzdiorite (c/. Streckeisen & le Maitre, 1979). The leucogneiss component of the Sand River Gneisses is composed mainly of plagioclase, quartz and potash feldspar, with lesser biotite, and minor hypersthene, hornblende and opaque minerals (Table I, column 3). Geochemically the
rock is very like average granodiorite (Table I, column 4), which is consistent with its classification as a granodiorite (Streckeisen & le Maitre, 1979). The grey gneiss and the leucogneiss are interlayered and infolded as units commonly a metre or less thick, and rarely up to 10m thick (Fig. 4), and both have a gneissic fabric superimposed on the lithologic layering. Fabric and layering are cut by the mafic dykes at the Causeway locality, and at the Stockford locality. Igneous Events In addition to the Messina Layered Intrusion, the Sand River Gneisses have been intruded at least four times by dykes of various compositions and ages (Table II). Cross-cutting relationships and included fragments which demonstrate the time-sequence of these igneous events are exposed at the Causeway and Stockford localities. Early, light grey granodiorite-gneiss dykes are cut by very leucocratic granite-gneiss dykes and these, together with their host, the Sand River Gneisses, are cut by the deformed and metamorphosed mafic dykes of two ages (Barton et al., 1977). The older mafic dykes, dated at — 3570 Ma, are particularly well exposed at the Causeway locality (Fig. 4) and are called the Causeway Dykes. The younger (-3060 Ma) dykes are, in addition to the Causeway Dykes, exposed at the Stockford locality (Fig. 5), and are called the Stockford Dykes. They contain frag-
332
R. E. P. TABLE I
Geochemistry and modal mineralogy of the Sand River Gneisses* . Si0 2
1
2
3
4
61.01
61.59
72.82 *
68.97
Ti0 2
0.75
0.66
0.27
0.45
A1 2 0 3
18.24
16.21
14.41
15.47
Fe 2 0 3 t
5.86
6.73
2.39
3.40
MnO
0.20
0.10
0.08
0.06
MgO
2.44
2.80
0.64
1.15
CaO
5,94
5-38
2.81
2.99
Na 2 0
1.56
2.10
2.37
3.16
P2O5
0.34
0.26
0.08
0.19
H20
0.50
1.22
0.48
0.70
TOTAL
100.49
100.42
100.17
100.23
Quartz
23.3
37.4
K-feldspar
NIL
13.5
Plag ioclase
61.2
42.2
'(% An)
(a) 45.6
(b) 40.0 5.6
Biotite
9.5
Hornblende 6 hypersthene
5. 1
1.0
Opaques
0.9
0.3
i.
2.
34. (a) (b) t
Grey gneiss, average of 20 analyses, Sand River area. Average quartz-diorite (Daly, 1933). Leucogneiss, average of 14 analyses, Sand River area. Average biotite granodiorite (Nockolds, 1954). Average of 8 analyses by ARL SEMQ microanalyser Average of 2 analyses University of the Witwatersrand. Total iron as Fe203. Major-element geochemistry by XRF; University of the Wi twatersrand.
ments of, and transect, trains of boudins of anorthositic gneiss which occur in the Sand River Gneisses, and which are considered to represent genetic equivalents of, and possible feeders of, the Messina Layered Intrusion (Barton et al., 1979a). The sequence of igneous events and their possible timing are summarized in Table II.
FRIPP
quently about NE upright axial surfaces (D5). These fold events (D4 and D5) were of similar scale (frequency) and produced the elongate domes and basins which define the dominant NE trend in the area. Such structures can be seen on Figure 1 about 5 km southeast of the southeastern edge of the area annotated as Figure 2, and also on Figure 2., defined by banded magnetite quartzite in the eastern part of the area. The last fold event produced large-scale, heterogeneously developed folds with axial surface traces trending NW and having wavelengths and amplitudes of about 20 km. In the rocks immediately to the east of Messina, Fripp et al. (1979) considered the D6 folds (called D4 by them) and the young shear zones to be genetically related, and these are given the same annotation. In the Sand River Gneisses, the fold events subsequent to D2 are recognized because they deform both the lithologic layering and the D1 gneissic fabric, and as indicated on the map of the contact between the Sand River Gneisses and the Beit Bridge gneisses at the Causeway locality (Fig. 3), the fabric is apparently conformable EXPL AN A TION 4{ CS
Outcrop / Inferred Beit Bridge gneisses Mafic "Causeway" Dykes Outcrop/ Inferred Sand River Gneisses inferred /
contact
synform/antiform -axial trace metres
Tectonic Events
Including the early gneissic fabric called Dl, which is the oldest recognizable tectonic event in the area, the Sand River Gneisses have undergone six distinguishable deformation events. Two, called D2 and D6, are heterogeneous and involved the formation of small-scale shear zones. The D2 shear zones are within the Sand River Gneisses, but absent from the adjacent Beit Bridge gneisses, whereas the late (D6) shear zones are within both suites of gneiss, and are the last (Archaean) ductile deformation event recognised in them. Subsequent to D2, the tectonic histories of the Sand River Gneisses and the Beit Bridge gneisses in this area are apparently similar (Table II). After the mainly nappe-forming event (D3), the gneisses were folded about upright, mainly NW-trending axial surfaces (D4) and subse-
Fig. 3. Generalized geology of the Sand River pavements at the Causeway locality showing the main structures and broad relationships between the Sand River Gneisses and the Beit Bridge gneisses.
SAND RIVER GNEISSES, LIMPOPO MOBILE
BELT
333
TABLE II
Summary of igneous and tectonic events recognized within the Sand River area Igneous events
Probable age (Ma)
Tectonic Events Young shear zones and large-scale folds " Upright folds (D s ) Upright folds and fabr ics
(D 6 )
(DO
Fold nappes, folds and local fabrics (D3) Intrusion Old shear zones
(D 2 )
Causeway Dykes
~ 3060 ?
Barton et al.
- 3350 ?
Barton (1979)
~ 3570
Barton et al.
(1977)
Barton e t al.
(1978)
Granitic Dykes
?
Granodioritie Dykes
1 Major metamorphic fabr ic (Di)
Sand River Gneisses
with the Iithologic layering in the latter. This is true at all known exposures of the contact, and deformation (strain) generally becomes more intense as the contact is approached. DISCUSSION The Sand River Gneisses, which have been deformed and metamorphosed together with the Beit Bridge gneisses, are clearly the older, as is particularly well demonstrated by the lack of
Fig. 4.
Reference Barton et a l . (1979) Fripp et al. (1979)
?
Stoekford,Dykes
Messina Layered
~ 2700 ?
- 3790
(1977)
>3790
Causeway Dykes in the latter and the presence of at least two additional deformations in the former. This relationship is supported by the very ancient age for the Sand River Gneisses ( - 3 7 9 0 Ma; Barton et al., 1978), and can be interpreted in two ways: (i) either the Sand River Gneisses represent an older fragment of sialic crust which was brought into contact with the Beit Bridge gneisses by mainly horizontal tectonic translation and crustal shortening or (ii) the Sand River
Generalized detailed map of part of the Sand River pavement at the Causeway locality. The Causeway Dyke shown is about 3570 Ma old and clearly cuts the Iithologic layering of the gneisses.
334
Fig. 5.
R. E . P . F R I P P
P h o t o g r a p h taken at the S t o c k f o r d locality which shows the cross-cutting relationship between a — 3570 Ma old Causeway Dyke and a - 3 0 6 0 M a old S t o c k f o r d Dyke. T h e latter also clearly transects the banding in the gneisses, although the f o r m e r is apparently c o n c o r d a n t at this o u t c r o p .
Gneisses represent a sialic basement, in the strict sense, on which the precursors to the Beit Bridge gneisses were deposited. The early (D2) shear zones in the Sand River Gneisses probably formed during relatively rapid (bulk) strain and at relatively high temperatures such as occur in rocks under a supracrustal load and undergoing relatively rapid uplift (e.g. Ramsay, 1976). Furthermore, these shear zones are younger than the - 3 5 7 0 Ma old Causeway Dykes and older than the Beit Bridge gneisses; that is, at least older than - 3 3 5 0 M a . Therefore they may have formed during an uplift- deposition process involving the deposition of the Beit Bridge gneiss precursors into a depository floored at least in part by the Sand River Gneisses. It is reasonable to suggest therefore that the Beit Bridge gneisses, which were formed some time between - 3 5 7 0 and - 3350 Ma ago (Table II), are supracrustal in the strict sense. The original nature of the Sand River Gneisses is obscure. The main gneissic fabric-forming event (Dl) was accompanied by a relatively highgrade metamorphic event about 3790 Ma ago (Barton et al., 1978) beyond which no geochronological method has penetrated (Barton, 1979). A model age for their initial Rb-Sr ratio of about
0.70122 (Barton et al., 1978) indicates that they cannot be older than about 4200 Ma (Barton, 1979). Their whole-rock geochemistry (Table I) is not diagnostic of their original nature and they may have been greywackes of epiclastic or volcanic-clastic origin, or felsic volcanic or plutonic rocks. The mutual disposition of the grey gneiss and the leucogneiss in the field as thin interlayers, rarely up to 10m thick (Fig. 4), throughout the known extent of the Sand River migmatite, that is, over distances of about 50km or more, argues against a plutonic origin. The suggestion that the leucogneiss is an ancient pre-Dl anatect of the grey gneiss is supported by plots of their major, refractory, inter-element correlations, and negated by their trace (Rb, Sr, Ba) interelement correlations (Fripp, unpubl. data), which are very poor. At this stage, therefore, the original nature of the Sand River Gneisses is unresolved, and must await the findings of other investigations such as rare-earth element and Nd-Sm isotopic studies. CONCLUSIONS The Sand River Gneisses represent a remnant of ancient sialic crust within the Central Zone of the Limpopo Mobile Belt. Other migmatites,
SAND RIVER GNEISSES, LIMPOPO MOBILE BELT 335 similar in aspect to those at Sand River, are covered within the Kaapvaal Craton. The Sand regarded as basement to the supracrustal gneisses River Gneisses therefore may provide a useful elsewhere in the Limpopo Belt (du Toit & van chronostratigraphic, lithostratigraphic and tecReenen, 1977) and the "grey gneiss" is described tonic link between the Rhodesian and Kaapvaal by Sohnge (1946) as being very widespread in the cratons, and help to provide further clues to early region/ Furthermore, granitoid gneisses and crustal evolution in this part of southern Africa. migmatites of perhaps comparable age, and at least older than about 3500 Ma, have been recog- ACKNOWLEDGMENTS nized on the adjacent Rhodesian Craton (Fig. 1) This paper is contribution number 55 of the (Hawkesworth et al. 1975; Moorbath et al., South African contribution to the International 1977), Somewhat similar bimodal gneisses cut by Geodynamics Project, and is sponsored by the mafic dykes have been observed north and south- Council for Scientific and Industrial Research. west of the Barberton greenstone belt (van J. M. Barton, Jr., T. N. Clifford, P. C. HorNierop, pers. comm.; Anhaeusser & Robb, 1978; rocks, A. J. Martin and M. J. de Wit are thanked Robb, 1981). Accordingly, it is likely that rem- for their comments and criticisms of an earlier nants of ancient sialic crust will also be dis- draft of this paper. y
REFERENCES
ANHAEUSSER, C. R . , & ROBB, L. J . , 1978: Regional and
detailed field and geochemical studies of Archaean trondhjemitic gneisses, migmatites and greenstone xenoliths in the southern part of the Barberton Mountain Land, South Africa. Inf. Circ. econ. geol. res. Unit, Univ. Witwatersrand, Johannesburg, 125. BAHNEMANN, K. P., 1971: in Morrison, E. R., & Wilson, J. F. (Compilers), Symposium on Granites, Gneisses and Related Rocks: Excursion Guidebook, 21. BARTON, J. M . JR., 1979: Crustal evolution clues. Nuclear Active, 21, 16-19.
BARTON, J. M . J R . , FRIPP, R . E . P . , & HORROCKS,
P. C., 1979: Effects of metamorphism on the RbSr and U-Pb systematics of the Singelele and Bulai Gneisses, Limpopo Mobile Belt, southern Africa. Trans, geol. Soc. S. Afr., 82, 259-269.
BARTON, J . M . JR., FRIPP, R . E . P . , HORROCKS, P . C . , & MCLEAN, N . 1979A: The geology, age and tec-
tonic setting of the Messina Layered Intrusion, Limpopo Mobile Belt, southern Africa. Am. J.
5c/., 279, 1108-1134. BARTON, J . M . JR., FRIPP, R . E . P . , & RYAN, B., 1977:
Rb/Sr ages and geological setting of ancient dykes in the Sand River area, Limpopo Mobile Belt, South Africa. Nature, Lond., 267, 487-490.
BARTON, J . M . J R . , RYAN, B., & FRIPP, R . E . P . , 1978:
The relationship between Rb-Sr and U-Th-Pb whole-rock and zircon systems in the -3790 Ma old Sand River Gneisses, Limpopo Mobile Belt, southern Africa; in Zartman, R. E. (Ed.) Short papers of the Fourth International Conference on Geochronology, Cosmochronology, Isotope Geology. Open File Rep., U.S. geol. Surv., 78-701, 27-28.
Cox, K. G . , JOHNSON, R . L . , MONKMAN, L . J . , STILIMAN, C . J . , VAIL, J . R . , & WOOD, D . N . , 1965: T h e
geology of the Nuanetsi igneous province. Phil. Trans. R. Soc. Lond., A257, 72-218. of the Earth. McGraw-Hill, New York.
DALY, R. A . , 1933: Igneous Rocks and the Depths
DU TOIT, M . C . , & VAN REENEN, D . D . , 1977: T h e
southern margin of the Limpopo mobile belt, northern Transvaal, with special reference to metamorphism and structure; in Ermanovics, I. F., Key, R. M., & McEwen, G. (Eds) The Proceedings of a Seminar Pertaining to the Limpopo Mobile Belt, 83-97. Bull. geol. Surv. Botswana, 12.
FRIPP, R . E . P . , LILLY, P . A . , & BARTON, J . M . JR,
1979: The structure and origin of the Singelele Gneiss at the type locality near Messina, Limpopo Mobile Belt. Trans, geol. Soc. S. Afr., 82, 161-167.
HAWKESWORTH, C . J . , MOORBATH, S., O'NIONS, R . K., & WILSON, J . F . , 1975: Age relationships between
greenstone belts and "granites" in the Rhodesian Archaean craton. Earth planet. Sci. Lett., 25,
251-262. MACGREGOR, A. M . , 1947: An outline of the geological
history of Southern Rhodesia. Bull. geol. Surv. Sth Rhod., 38.
MASON, R., 1973: The Limpopo mobile belt—southern Africa. Phil. Trans. R. Soc. Lond., A273, 463-485. MOORBATH, S., WILSON, J . F . , GOODWIN, R . , & HUMM,
M. 1977: Further Rb-Sr age and isotope data on early and late Archaean rocks from the Rhodesian craton. Precamb. Res., 5, 229-239. NOCKOLDS, S. R . , 1954: Average chemical compositions of some igneous rocks. Bull. geol. Soc. Am., 65,
1007-1032. RAMSAY, J . G . , 1976: Displacement and strain. Phil.
Trans. R. Soc. Lond., A283, 3-25.
ROBB, L. J . , 1981: Detailed studies of select migmatite
outcrops in the region southwest of the Barberton greenstone belt and their significance concerning the nature of the early Archaean crust in the region. Spec. Pubis geol. Soc. Aust., 7, 337-349. SOHNGE, P. G . , 1946: The geology of the Messina copper mines and surrounding country. Mem. S. Afr. geol. Surv., 40.
STRECKEISEN, A . , & LE MAITRE, R . W . , 1979: A chemical approximation to the modal Q A P F classi-
fication of the igneous rocks-. N. Jb. Miner. Abh., 136, 169-206.
DETAILED STUDIES OF SELECT MIGMATITE OUTCROPS IN THE REGION SOUTHWEST OF THE BARBERTON GREENSTONE BELT AND THEIR SIGNIFICANCE CONCERNING THE NATURE OF THE EARLY ARCHAEAN CRUST IN THE REGION* L. J. Robb Economic Geology Research Unit, University of the Witwatersrand, 1 Jan Smuts Avenue, Johannesburg 2001, Republic of South Africa ABSTRACT Detailed studies of migmatites in the region southwest of the Barberton greenstone belt have been carried out in the hope that they may elucidate primary relationships between early granitic (.sensu lato) and greenstone crust in the area. Certain migmatites are formed by the intrusion of t o n a l i t i c / t r o n d h j e m i t i c m a g m a into segments of pre-existing greenstone (ensimatic) crust; in others the mafic component of the migmatite clearly intrudes well-foliated tonalite or t r o n d h j e m i t e gneisses; a third type is characterized by anatectites intruding both gneiss and greenstone c o m p o n e n t s . Detailed m a p p i n g of the migmatite outcrops has shown that tonalite and trondhjemite gneiss plutons are characterized by discrete, episodic emplacement. Where recognized, the older trondhjemite gneiss phases have an equivocal relationship with the greenstones; the younger t o n a l i t e / t r o n d h j e m i t e gneiss plutons, however, clearly intrude both greenstones and older gneiss phases.
INTRODUCTION Ever since the pioneering work of Sederholm was responsible for the selective description of migmatites and the recognition of their role in granite pedogenesis, these " m i x e d r o c k s " have been the object of innumerable studies. In recent times migmatites are recognized as having been formed by in situ anatexis (von Platen, 1965; Ashworth, 1976), by subsolidus metamorphic differentiation (White, 1966; Misch, 1968; Hedge, 1972; Amit & Eyal, 1976; Yardley, 1978), by metasomatism (King, 1965; Viljoen & Viljoen, 1969a) and by physical emplacement of an externally derived m a g m a ( L o w m a n , 1965; Van Schmus & Anderson, 1977). Workers in highly deformed areas have also demonstrated the relationship between intensity of d e f o r m a t i o n and genesis of migmatites (Mackenzie, 1957; Myers, 1978). Inasmuch as A r c h a e a n terrains the world over are characterized by controversy concerning the nature and development of the primordial crust, it was decided to examine migmatites in the Barberton region, in the hope that their characteristics might shed light on this problem. The well-exposed terrain southwest of the Barberton greenstone belt (Fig.. 1) was selected, as the * South African G e o d y n a m i c s Project P a p e r No. 58
Spec. Pubis geol. Soc. Aust., 7 (1981)
regional mapping here is virtually complete (Anhaeusser & Robb, 1978; Anhaeusser, 1980). The results of detailed mapping (1:50) here have led to the recognition of three types of migmatite. This paper describes them, and attempts to explain their formation in terms of the regional geology of the Barberton Mountain Land. A description of features only recently observed provides information that has led to a better understanding of the nature and development of the early Archaean crust in the area. MIGMATITE DESCRIPTION A number of migmatite outcrops were mapped as part of this study, all but one in the region outlined in Figure 1. The pertinent characteristics of a representative selection of these migmatites are presented in schematic f o r m in Figure 2. The seven diagrams (A-G) presented in this figure depict portions of individual, geographically discrete, migmatitic outcrops, and their approximate localities are described in the captions to Figure 2. The migmatites described in this paper are all related to the oldest suite of granites (sensu lato) in the area (Barton, 1981) and differ f r o m the migmatites related to younger granitoids, which
338
L. J. R O B B
1
D00RNH0EK PLUTON
2
BATAVIA REGION
3
ROOIHOOGTE REGION -(Situated 20km west
4
KAAP VALLEY PLUTON
of the d i a g r a r ^ *
5
STOLZBURG PLUTON
6
WEERGEVONDEN REGION
7
THEEBOOM REGION
POST-TECTONIC SYENITIC AND ADAMELLITIC PLUTONS TONALITESXTRONDHJEMITES BARBERTON GREENSTONE BELT
Fig. 1.
AREA WITHIN WHICH THE MIGMATITE OUTCROPS DESCRIBED^ ARE LOCATED
M a p of the southwestern portion of the Barberton greenstone belt showing the region in which the migmatites described in the text are located. Also shown are the localities of some of the discrete tonalite/ trondhjemite plutons or "cells" referred to.
are described elsewhere (Anhaeusser & Robb, 1981). As a result, the migmatites in the study region can be interpreted in two ways; firstly, in terms of their often complex style of deformation they could represent a sialic basement upon which the Barberton greenstone belt was deposited, or secondly, they may have formed in response to the physical interaction between pre-existing greenstones and marginally younger, or coeval (according to isotopic data; Barton, 1981) tonalitic and trondhjemitic gneisses. Regional mapping in the area, however, shows that the migmatites are not randomly distributed but are invariably close to the numerous greenstone xenoliths in the region (Anhaeusser & Robb, 1978). This suggests, albeit superficially, that the migmatites are more likely to represent the second of the above possibilities.
Detailed mapping of migmatite outcrops indicates that three types of migmatite can be described: (a) Type 1, where tonalitic or trondhjemitic gneisses clearly intrude mafic (and felsic) units which are correctable with recognized greenstone lithologies (Figs 2A and D); (b) Type 2, where both gneiss and greenstone components are present in addition to often significant proportions of anatectic material that has intruded all pre-existing components, thereby obscuring the primary relationships between them (Figs 2B and C); and (c) Type 3, where the mafic component of the migmatite intrudes well-foliated tonalite or trondhjemite gneisses (Figs 2E and G). In terms of this subdivision Type 1 and Type 3 migmatites represent opposite extremes, as the
MIGMATITE SOUTHWEST OF BARBERTON 339 former has tonalitic or trondhjemitic gneiss that anatectites arranged in a systematic lit-par-lit is younger than mafic xenoliths that can often be array (Fig. 3B). Again, the primary relationship correlated with greenstone lithologies, whereas between gneiss and amphibolite is obscured and the latter consists of gneisses that are intruded by the interaction of anatectite and amphibolite is mafic dykes not directly correctable with the responsible for the development of a migmatite. greenstones. Type 2 migmatites may represent Typical Type 3 migmatites are shown in Figure either extreme, but invariably contain a signifi- 2E and G where amphibolitic dykes of tholeiitic cant anatectic component that obscures all composition clearly intrude well-foliated tonalitic primary relationships. or trondhjemitic gneisses. These dykes have been Migmatites of Type 1 are formed by the intru- involved in, at least, the later stages of dynamosion of tonalite/trondhjemite gneiss into pre- thermal metamorphism that affected the region existing greenstones of variable composition. For and have been deformed and migmatized so that example, in Figure 2A and Figure 3A a tonalitic some appear similar to mafic xenoliths described gneiss intrudes a heterogeneous (migmatized) from Type 1 migmatites (Fig. 3C). As a result felsic meta-tuff. The felsic meta-tuff has been their recognition depends on the presence of a extensively recrystallized as well as now contain- clearly discordant relationship between them and ing innumerable anatectic veinlets. Where un- pre-existing trondhjemite/tonalite gneisses. affected by these veinlets this unit appears to have Where intense structural deformation has been retained its original chemistry (Sample A2, Table responsible for the concordance of many linear 1) which is markedly similar in both major and mafic bodies with respect to enveloping gneisses, trace elements to that of unaltered acid tuffs in a the identification of possible dykes becomes large nearby greenstone xenolith (Anhaeusser & equivocal and, hence, they are not widely Robb, 1978). In places the felsic meta-tuff has recognized in the region. been tightly folded and it is here that the maxiMany of the migmatite outcrops encountered mum development of anatectic veinlets occurs cannot be unequivocally interpreted principally (Fig. 2A). The anatectic material has apparently because of the effects of deformation. In Figure migrated preferentially from fold limbs to fold 2C, for example, the foliation in a boudinaged hinge zones over a relatively small area. This out- lens of trondhjemite gneiss is truncated by crop, therefore, demonstrates a small-scale amphibolite in places. If the orientation of maxiexample of where the extent of in situ migmatiza- mum principal strain during deformation were tion in metamorphic rocks is closely related to slightly oblique to the contact plane of the two both style of deformation and magnitude of units, the truncation could be due to the rotation of boudinaged units (i.e. lozenge-shaped strain (Mackenzie, 1957). In Figure 2D an amphibolitic meta-basalt is in- boudins; Rast, 1956; Charlesworth & Evans, truded by a well-foliated trondhjemite gneiss. 1962). On the other hand, a close examination of Both units have been deformed twice so that a the amphibolite-banded trondhjemite gneiss consuperimposed fold pattern prevails (i.e. small tact shows that the leucocratic banding in the dome and basin-like features occur). Many of the gneisses is also truncated by the amphibolite (Fig. mafic xenoliths in this category have been exten- 3E). This could be construed as evidence for the sively assimilated by the intrusive tonalites or intrusion of gneiss by amphibolite in spite of the trondhjemites. Despite the deformation and stretching and concordancy that has generally assimilation, clearly cross-cutting and intrusive been applied to the outcrop. For example, relationships are preserved, and these serve to Dearnley & Dunning (1968) have, in some instances, used only short lengths of discordant characterize the Type 1 migmatites. Type 2 migmatites characteristically contain contact when identifying the highly deformed anatectites that intrude both pre-existing am- dykes of the Lewisian in Scotland. In Figure 2F, dominantly linear bodies of amphibolitic and gneissic components. In Figure 2B an amphibolitic meta-basalt of komatiitic com- phibolite of tholeiitic composition are enveloped position has been intruded and fragmented by a by well-foliated trondhjemite gneiss such that felsic anatectite. Here the anatectite is responsible mutual contacts are generally concordant. The for migmatization (i.e. the formation of amphibolite has been slightly assimilated by felsic agmatite) and the relationship between trondhje- material and possibly represents xenoliths within mite gneiss and amphibolite is not preserved. the gneiss. However, in view of the assimilated Figure 2C illustrates a broad contact zone be- nature of the dykes described above and judging tween well-foliated banded trondhjemite gneiss by the dyke-like "horned" structure shown in and amphibolitic meta-basalt, again of komatiitic Figure 2F as well as the occasional small lengths composition. The contact zone is represented by of discordant contact between gneissic schlieren the same amphibolite intruded by a series of felsic and amphibolite demonstrated in Figure 3D, it is 4
L- J-
340
R O B B
A
FAULTS AND FRACTURES PEGMATITIC VEINS ANATECTIC VEINS/VEINLETS MIGMATIZED FELSIC META-TUFF AMPHIBOLITIC XENOLITHS/DYKES TONALITIC GNEISS TRONDHJEMITIC GNEISS
GRASS
Fig. 2.
A : Exploded diagram illustrating the intrusion o f tonalitic gneiss into a heterogeneous (migmatized) felsic meta-tuff in the Weergevonden region. A l , A 2 etc are the localities o f analysed samples presented in Table I. B: Fragmentation o f amphibolite by anatectite in the R o o i h o o g t e region. C : Contact zone between trondhjemite gneiss (partly covered by grass on the left o f the diagram) and amphibolite where a lit-par-lit array o f anatectic veins has intruded both the trondhjemite and amphibolite. O n the right o f the diagram the foliation in a boudinaged lens o f trondhjemite gneiss is clearly dis-
MIGMATITE SOUTHWEST OF BARBERTON
341
cordant with respect to the amphibolite with which it is juxtaposed; from the Theeboom region. D: Deformed xenolith of amphibolite within well-foliated trondhjemite gneiss in the Batavia region. E: Dykes of amphibolite intruding well-foliated trondhjemite gneiss in the Batavia region. F: Linear dyke-like bodies of amphibolite within trondhjemite gneiss from an area east of the Weergevonden region. G: Deformed amphibolitic dyke intrusive into tonalitic gneiss from the Batavia region.
E ;. 3.
A: Intrusive relationships between tonalite gneiss (left) and heterogeneous (migmatized) felsic meta-tuff (right) in a Type 1 migmatite. B: Lit-par-lit array of anatectic veins intrusive into amphibolite in a Type 2 migmatite. C: Portion of an assimilated tholeiite dyke intrusive into well-foliated trondhjemite gneiss. D: Discordant contact between trondhjemite gneiss and a possible tholeiite dyke in a Type 3 migmatite from the outcrop in Figure 2F. The emplacement of this dyke may have been controlled, in part, by preexisting fractures. E: Photograph illustrating the discordance between amphibolite and boudinaged banded trondhjemite gneiss at the outcrop shown in Figure 2C.
M I G M A T I T E SOUTHWEST OF BARBERTON
343
TABLE I
Chemical analyses of select samples from the migmatite outcrops described in Figure 2 AT
^
2
3
4
5
6
7
8
9
10
11
12
61
13
14
A2
CI
G2
C3
C4
D1
D2
D3
El
E2
F1
G1
71.70
71.50
71.54
73.43
50.54
67.61
48.95
75.21
69.85
49.13
51.80
wt % Si0 2
:
73.47
68.77
Ti0 ?
D.52
V
A1203
1,7.18
. 13.05
Fe203
3.22;
.,2.18
:
0.20 .
50.26
0.20
0.15
0.38
0.12
0.73
0.30
1.22
0.03
0.31
0.88
1 » 25
15.62
15.60
16.48
14.96
9.40
17.86
14.97
13.52
15.56
16.60
13.66
12.10
2.19
1.30
2.43
0.88
13.47
3.08
11.40
0.76
3.23
10.31
15- 76
11.62
0.46
0.06
-
1.13
MnO
0„0l
- 0,02
0.01
0.01
0.01
0.01
0.17
0.01
0.01
0.14
0.23
0.09
MgO
1;. |4 ,... 0 . 9 0
0.85
0.40
1.04
0.44
10.47
1.20
6.14
0.22
1.23
8.10 ,
4.00
11.40
CaQ
0.77
3.05
1.79
2.55
1.12
11.52
4.35
11.49
1.94
2.93
9.70
6.72
4.45
3.57
4.29
5-90
4.26
4.59
1.51
4.61
0.91
2.12
1.92
4.10
0.79
Na?0 k2o
.
0.18
-
.
1-75
'
9.10
4.69
3-33
5.53
4.08
3.14
2.30
3.71
0- 99
0.58
0.72
1.22
1.46
1.38
0.66
0.07
0.06
0.04
0.15
0.07
0.01
0.09
0.08
0.01
L0I
0.85
0,48
0.63
0.42
0.36
0.60
1.02
0.48
0.53
0.54
0.75
0.97
1.53
0.65
TOTAL
100.91.
99-32
99.51
99-23
101.11
100.32
100.15
100.74
98.86
98.49
99-24
100.85
99.30
100.78
P2O5
ppm
0.07
YS
0.42
0.13
0.06
46
60
97
14
17
5
12
30
49
Sr
13k
122
434
604
810
339
184
549
148
231
386
434
100
190
Ba
22:0
993
120
620
584
655
81
159
10
82
237
228
248
34
Rb
58
76
Columns 1, 3 , 4 , 5 , 8 , 11 Columns 7 , and 9 Column 2 Columns 1 2 , 13 a n d 14 Columns 6 a n d 10
-
64
Tona1ite/trondhjemite gneisses Komatiitic and tholeiitic xenoliths Felsic meta-tuffaceous xenolith Tholeiitic dykes Anatectites
(i) Samples a n a l y s e d by X - r a y f l u o r e s c e n c e in the D e p a r t m e n t of G e o l o g y , U n i v e r s i t y of the W i t w a t e r s r a n d , by L . J . Robb (i; i) S a m p l e l o c a l i t i e s f r o m F i g u r e 2 e x c e p t C2 a n d C4 w h i c h c o m e f r o m n e a r b y the o u t c r o p shown in F i g u r e 2 C . (i i "f) T o t a l fron as F e 2 0 a .
equally likely that the amphibolites here may be intrusives. This interpretation is equivocal, however, as this migmatite outcrop is in an area where no other dykes of this nature have been recognized and where extensions of the dominantly tholeiitic upper portions of the Onverwacht Group in the Barberton greenstone belt were previously considered to have been intruded and migmatized by younger trondhjemitic gneiss. The brief descriptions above indicate that a variety of processes have been responsible for the development of the migmatites in the study area. On the one hand tonalitic/trondhjemitic gneisses have intruded and migmatized portions of recognizable greenstone (ensimatic) crust; on the other hand a suite of now deformed and migmatized mafic dykes has intruded tonalitic/trondhjemitic (sialic) crust. All the rocks involved have been deformed, often to the extent that primary relationships between migmatite components are obscured. The remainder of the paper discusses in more detail some of the characteristics of the three dominant phases of the migmatites, namely the tonalitic/trondhjemitic gneisses, the anatectites and the amphibolites.
trude the Barberton greenstone belt were largely considered to be cogenetic as well as broadly coeval (Viljoen & Viljoen, 1969a; Oosthuyzen, 1970). Recent work in the Barberton Mountain Land has, however, shown that this suite of rocks varies considerably in age between ~ 2.9b.y. and « 3.45 b.y. (Barton, 1981). In addition, geological mapping has revealed the presence of discrete older and younger trondhjemitic gneisses where the latter clearly intrude the former (Anhaeusser & Robb, 1978). This relationship can be clearly demonstrated in another section of the outcrop
TONALITE AND TRONDHJEMITE .. ,, GNEISSES Fig- 4. The relationship between older, banded trondhjemite gneiss and a younger, less wellIn the past, the numerous tonalite and foliated, trondhjemite gneiss in the vicinity of trondhjemite gneiss plutons that diapirically inthe outcrop shown in Figure 2C.
L. J.
344
ROBB
gneiss unit must intrude and, therefore, contain xenoliths of, combined older gneiss and amphibolite. In terms of the regional geology, younger, and volumetrically dominant, tonalite/ trondhjemite gneiss plutons similarly intrude and deform pre-existing greenstones (Fig. 2A and D); they also, however, intrude any older sialic crust, whose relationships with respect to greenstones may be equivocal principally because of the deformation subjected to both units. The development of migmatites usually takes place along the margins of homogeneous tonalitic/ trondhjemitic plutons and it is these places, therefore, that are most likely to preserve
illustrated in Figure 2C. The banded trondhjemite gneiss illustrated here, a boudinaged lens of which is directly juxtaposed with the amphibolitic unit in Figure 2C, is clearly intruded by a younger trondhjemite gneiss some distance from the amphibolite. The younger gneissic unit, illustrated in Figure 4, does not contain the leucocratic banding characteristic of the older unit and transects both the banding and foliation in the latter. Both units have similar major-element compositions but are characterized by different Sr contents (Anhaeusser & Robb, 1978; Table I and Fig. 5). The relationships seen at this outcrop indicate that the younger, and volumetrically dominant,
•1000
-800
I-600 ppm
400
1
D 0 0 R N H 0 E K P L U T O N - 9 SAMPLES (ROBB, UNPULISHED DATA)
2
BATAVIA REGION -
8 SAMPLES
(ANHAEUSSER AND ROBB. UNPUBLISHED DATA )
3 R00IH00GTE REGION-7 SAMPLES KAAP VALLEY PLUTON -
4
41 SAMPLES
(ROBB. UNPUBLISHED DATA)
200 5
S T O L Z B U R G P L U T O N - 4SAMPLES
6
WEERGEVONDEN REGION-6 SAMPLES
7
THEEBOOM REGION - 6 SAMPLES
(ANHAEUSSER AND ROBB. UNPUBLISHED DATA,
TONALITE 60 - L
62
64
66 -L.
68
76
Si02 wt. % Fig.
5.
Plot o f S i 0 2 vs Sr for tonalitic a n d trondhjemitic gneisses from the region southwest o f the Barberton greenstone belt. Individual samples plotted are from Table I.
MIGMATITE SOUTHWEST OF BARBERTON 345 primary relationships between the greenstones particular, Type 2 migmatites (Figs 2B and C) are and remnants of sialic crust that may have pre- characterized by intrusion of significant propordated them. tions of anatectic material into both trondhjeA major problem in recognizing the episodic mitic and amphibolitic components. The anatecemplacement of tonalitic/trondhjemitic plutons tites are generally syntectonic and have been inis the.similarity in their major-element chemistry volved in all, or the greater part, of the deformaand the rarity of outcrops such as that shown in tional history of the migmatites (Figs 2A and D). Figure 4. Available data suggest, however, that They are invariably devoid of a significant mafic Sr may be used in "fingerprinting" geographi- component although they may contain small cally discrete tonalite/trondhjemite plutons or amounts of muscovite or chlorite and, subse"cells", some of which are shown in Figure 1. In quently, have high S i 0 (usually >70%) and low Figure 5 a plot of S i 0 vs Sr shows that certain F e 0 + M g O contents ( = 1% or less). plutons may be wholly trondhjemitic (where a Qtz trondhjemite is described as a leucotonalite and is defined here as having S i 0 > 7 0 % ; e.g. the Doornhoek Pluton), others may be wholly tonalitic ( S i 0 < 7 0 % ; e.g. the Kaap Valley Pluton) whereas most are characterized by both tonalitic and trondhjemitic compositions. However, each pluton, or geographically discrete cell, is characterized by a distinctive Sr content that varies between 100 and lOOOppm. Although the data indicate that each pluton or cell has a distinctive Sr content it is obvious that, with more information, overlaps will occur and only low-, intermediate- and high-Sr plutons will be recognized. The S i 0 vs Sr diagram is significant in terms of the description of migmatites because it demonstrates that the tonalite or trondhjemite component in the migmatites is generally similar in composition (particularly its Sr content) to that An of the trondhjemite pluton or cell with which it is associated. For example, samples A l , Bl, CI and El (from Table I and which correspond to the migmatite outcrops in Figs 2A, B, C and E respectively) all have compositions which correspond to the plutons or cells with which they are associated, namely the Weergevonden cell, the Rooihoogte cell, the Theeboom cell and the Batavia cell (Fig. 1). The older trondhjemite gneiss (CI), which occurs as a xenolith together with the amphibolite unit in the younger gneiss (C2) in the outcrop shown in Figure 2C, has a different composition to the latter (i.e. lower Sr) and falls in the compositional field of the Stolzburg Pluton. The tonalite gneiss D1 (Table I and Fig. 2D) comes from a migmatite outcrop associated with the Batavia cell and has a composition very close to it in Figure 5. It is apparent, therefore, that the tonalite/trondhjemite plutons are inti- Fig. 6. Ternary plots of mesonormative Qtz-Ab-Or (A) and An-Ab-Or (B) showing the composimately associated with the process of migmatizational fields of a number of anatectites in adtion such that the development, at least of Type 1 dition to those from Table I. Isotherms and and Type 2 migmatites, is genetically related to projected cotectic surfaces are from Luth et the emplacement of these bodies. al. (1964) and Kleeman (1965). Vertical hatching is the field of Or-depleted anatectites (7 samples); horizontal hatching is the field of ANATECTITES Or-rich anatectites (11 samples). All migmatite outcrops show evidence of varying degrees of anatexis during their formation. In 2
2
2
3
2
2
Projected eutectic points and thermal minima
2
346
L. J. ROBB
The anatectites can be divided into two categories, on the basis of their major-element compositions. The first type, which is characteristic of anatectites from the migmatites shown in Figures 2B and C, have high Si0 2 , low F e 2 0 3 + MgO and high K 2 0 values such that K 2 0 / N a 2 0 = 1 (Column 6, Table I). The second type is similar except for the K 2 0 content which is much lower such that K 2 0 / N a 2 0 = 0.1 - 0.2 (Column 10, Table I). These compositional characteristics reflect simply the presence, or lack, of microcline. Because of their dominantly quartz-Kfeldspar-plagioclase mineral assemblage, the anatectites are suitably represented in the granite system Qtz-Ab-An-0r-H 2 0, and the melt origin of these rocks can be assessed in terms of these components. The compositions of a number of anatectites additional to those listed in Table I (that come from the migmatites represented in Fig. 2) are plotted on Qtz-Ab-Or and An-Ab-Or ternary diagrams ir Figures 6A and B respectively. The anatectites are plotted on two different ternary diagrams so that misrepresentation of their composition in terms of only a portion of the granite tetrahedron is minimized. Isotherms are plotted in Figure 6 to give an estimate of the temperatures of melting for P h 2 o = 5kb (after Luth et al., 1964 and Kleeman, 1965). They demonstrate that the total temperature range over which isobaric cotectic melting may take place (i.e. where quartz-plagioclase solid solution—K-feldspar solid solution and melt are in equilibrium) is less than 50°C and, as Winkler & Lindemann (1972) have pointed out, is probably between 5-25 °C depending on the melt composition. In Figure 6, the composition of the orthoclase (microcline)-rich anatectites falls in a field that roughly coincides with minimum-melt (or eutectic, at higher Ph 2 o) compositions in the granite "system". In Figure 6A the field of Or-rich anatectites coincides with the envelope of projections of ternary eutectic points and thermal minima for a range of P H 2 Q and Ab/An as defined by Luth et al. (1964) and von Platen (1965). In Figure 6B likewise, they coincide with the projected cotectic surface as defined by Kleeman (1965). In the granite tetrahedron, therefore, these anatectites will plot on, or near, the cotectic line where melt is in equilibrium with three coexisting phases (i.e. quartz-plagioclase solid solution-orthoclase). As such, these anatectites probably represent relatively small degrees of partial melt where the temperature of anatexis did not rise significantly above that where melting commenced. It can also be deduced that the parent rock must have contained plagioclase, quartz and, at least some orthoclase in order that
eutectic compositions could form (alternatively the parent may have contained a phase such as biotite that melted incongruently thereby resulting in K 2 0-enriched melt). Such a parent might well have been a tonalite/trondhjemite gneiss which commonly contains — 5% microcline. The compositions of orthoclase (microcline)depleted anatectites all plot in a field that is somewhat removed from minimum melt or eutectic compositions in Figure 6. In the granite tetrahedron these anatectites will undoubtedly lie in the orthoclase-depleted region where, according to projected isotherms, melt compositions have higher (by 40-50 °C at P H 2 o = 5kb) temperatures. Winkler & Lindemann (1972) and Winkler & Breitbart (1978) have stressed that compositions of this nature cannot be precluded from having a magmatic or partial melt origin. This is because the composition of early melts will only remain on the cotectic line in the granite tetrahedron until one of the equilibrium phases is used up and, therefore, the criterion that defines a magmatic melt is not its proximity to the isobaric cotectic. If a trondhjemitic parent with 5% orthoclase is considered and a partial melt consisting of equal proportions of quartz, plagioclase and orthoclase is derived from it, then as soon as melting exceeds 15% the melt composition will leave the cotectic and change in a direction away from the Or apex of the tetrahedron. Here the melt is in equilibrium with only two phases, namely plagioclase and quartz. Orthoclasedepleted anatectites may, therefore, represent higher degrees of partial melt of the same tonalitic/trondhjemitic parent from which the Or-rich anatectites were derived. Alternatively, they may have been derived from a dominantly plagioclase + quartz parent (plagiogranite), but such rocks are volumetrically insignificant in the region. These considerations support the field impression that the anatectites were derived by localized melting, probably of a tonalitic or trondhjemitic precursor, during formation of the migmatitic zones. The degree of melting was variable and probably facilitated by dehydration of mafic greenstones in zones where magmatic activity prevailed.
AMPHIBOLITIC XENOLITHS AND DYKES Most of the migmatitic outcrops studied are characterized by a mafic component, with the exception of that in Figure 2A where a metamorphosed felsic tuffaceous rock occurs as one discrete component of the migmatite. Mafic components of the migmatites can all be described as amphibolites as hornblende invariably constitutes the dominant mineral phase. In some of the more
MIGMATITE SOUTHWEST OF BARBERTON 347 magnesian xenoliths poikiloblastic clinopyroxene occurs, whereas the tholeiite dykes of Type 3 migmatites are characterized by minor biotite in addition to hornblende. Most amphibolites contain small, but variable, amounts of plagioclase and quartz and some contain accessory chlorite and sphene. The following section compares the chemistry of amphibolites (meta-basalts) from the migmatite outcrops with that of basalts and tholeiites from the Barberton greenstone belt itself. Viljoen & Viljoen (1969Z?) originally defined a series of basaltic komatiites (i.e. specifically Barberton-, Badplaas- and Geluk-type basaltic komatiites with = 10% MgO, = 15% MgO and ^ 2 0 % MgO Fig. 7. Plot of CaO/A1 0 vs K 0 for amphibolites from the migmatite outcrops described in the respectively) in the greenstone belt that were text. Also shown are the fields of altered and characterized by C a 0 / A l 0 > 1 . Other workers unaltered Ameralik dykes from West in Archaean greenstone belts have subsequently Greenland (after Gill & Bridgwater, 1979). shown, however, that komatiites are not rigorMT—average meta-tholeiite; BBK—average ously defined by C a 0 / A l 0 > 1 and have used Barberton-type basaltic komatiite; GBK— other features, such as MgO content, and distincaverage Geluk-type basaltic komatiite; PK— average peridotitic komatiite (after Viljoen & tive skeletal textures. to recognize these rocks Viljoen, 19696). Individual samples plotted (Arndt et al., 1977; Sun & Nesbitt, 1978). In the are from Table I. Barberton greenstone belt too, it has now been demonstrated that unaltered basaltic komatiites have a C a 0 / A l 0 ratio that ranges between I). By comparison other tholeiite dykes have 0.86-2 5 whereas tholeiitic lavas vary between lower MgO (4-8%, Table I) but have composi0.66-0.85 (Smith & Erlank, 1978). Nevertheless, tions suggestive of assimilation or residence conthe Barberton basaltic lavas generally have a tamination. For example, samples E2 and F1 higher C a 0 / A l 0 ratio than those, for example, have higher K 0 contents (1.45% and 1.38% from the Abitibi greenstone belt in Ontario respectively) with concomitantly higher Rb con(Arndt et al., 1977) and the distinction between centrations (64 and 30ppm) than those usually komatiites and tholeiites in this study, is still observed in tholeiites. These compositions are taken, for convenience, to be at C a 0 / A l 0 = 1. probably also due, in part, to the more evolved Many of the amphibolitic xenoliths from the chemistry of the tholeiites,as well as to assimilamigmatites have the chemical attributes of tion and contamination. The effects of assimilation and contamination Barberton-type basaltic komatiites (Table I). Amphibolites from Type 1 and Type 2 migmatites have been noted in West Greenland in the prehave MgO contents ranging from 10-14% MgO 3.0b.y. Ameralik dykes (Gill & Bridgwater, with C a 0 / A l 0 > 1 and low (<0.8%) T i 0 . 1979). In Figure 7, the fields of unaltered (A) and Amphibolites from these categories may, altered (B) Ameralik dykes are shown on a plot of however, be tholeiitic with lower MgO and K 0 VS C a 0 / A l 0 and suggest a similarity in the C a 0 / A l 0 and distinctly higher T i 0 (sample composition of these dykes with those of the D2, Table I). All xenoliths analysed are little Type 3 migmatites. The relatively unaltered altered apart from the effects of a retrogressive sample G1 plots close to A whereas the assimimineral assemblage and, where sampled, compo- lated samples E2 and F1 plot in the field of B. sitions had not been markedly changed by intrud- Gill & Bridgwater (1979) have shown that Ameralik dykes range between i.5-5.5% K 0 when ing felsic magma. By comparison, the amphibolitic dykes that intrusive into Amitsoq gneisses, by comparison form the Type 3 migmatites have usually been with a range of 0.88-1.00% K 0 when intrusive partially assimilated and probably affected by into mafic supracrustals. They attributed this difresidence contamination (Fratta & Shaw, 1974). ference (which is only significantly noticeable in Where relatively unaltered, the amphibolitic terms of K 0 ) to residence contamination, but dykes have a composition similar to an average obviously where the dykes have been deformed tholeiitic lava from the Barberton greenstone belt and metamorphosed these effects could equally (Fig. 7). S a m p l e d (from the outcrop shown in be attributed to assimilation and migmatization. In terms of similarities in age, syntectonic Figure 2G) has a higher MgO content (11.4%) than most tholeiites but C a 0 / A l 0 < 1 (Table character, chemical composition, effects of 2.5-
A
FIELD OF W E L L PRESERVED AMERALIK DYKES
B
FIELD OF A L T E R E D A M E R A L I K D Y K E S
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Ca0/AI203 -1
Ca0/AI203 - 1
1.0-
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AMPHIB0LITE FROM FIGURE 2B
.
BBK
GBK •
MT
0,6
0.8
.
1.0
.
A 1.2
pi/
*
1.4
1.6
1.8
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2
2
3
3
2
3
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3
3
2
2
2
2
3
2
2
3
3
2
3
2
2
2
2
2
3
348
L. J. R O B B
alteration and even mutual presence of biotite (Gill & Bridgwater, 1979) it appears that the tholeiitic dykes of the Type 3 migmatites have a close analogue in the Archaean Ameralik dykes of West Greenland. These dykes, together with counterparts in the Lewisian of Scotland, have been described as the oldest known mafic magmatic event to have taken place within a crustal sialic environment. CONCLUSIONS The study of migmatites in the region southwest of the Barberton greenstone belt has led to the recognition of three distinct types. The first type occurs when tonalitic or trondhjemitic magma has intruded and interacted with preexisting greenstone (ensimatic) crust; a second type is characterized by the generation of significant amounts of anatectic material that intrudes both the gneiss and greenstone component and is, subsequently, responsible for migmatization; the third type occurs when tholeiitic dykes intrude pre-existing tonalite/trondhjemite (sialic) crust. All migmatites encountered have been somewhat affected by deformation. In places this has not been intense enough to obscure the primary relationships often preserved in the migmatite outcrops, but elsewhere high strains have resulted in a tendency to parallelism of previously discordant contacts so that the distinction between deformed dykes and xenoliths is obscure. Geochronological and field evidence indicates that the tonalite and trondhjemite gneiss plutons in the region are characterized by discrete, episodic emplacement, in the interval between = 3.45b.y. and = 2.9b.y. ago. The younger and volumetrically dominant gneiss plutons clearly intrude, and deform, the pre-existing ensimatic crust. They also intrude any older sialic crust whether it pre-dated or post-dated the greenstones. Primary relationships between early ensiAMIT, O., & EYAL, Y., 1976: The genesis of Wadi
Magnsh migmatites (N-E Sinai). Contrib. Mineral
Petrol., 59, 95-110.
ANHAEUSSER, C . R., 1980: A geological investigation of
matic and sialic crust are only likely to be preserved, therefore, in migmatites that formed from the interaction of older tonalite/trondhjemite gneisses with the earliest greenstone successions. The study did not reveal any unequivocal relationships indicating that primitive komatiitic basalts were extruded onto, or that komatiitic feeder dykes were intrusive into, an earlier sialic crust. Although their age is not known, mapping reveals the presence of tholeiitic dykes which are clearly intrusive into pre-existing sialic crust. Although these dykes are relatively insignificant, their recognition has been obscured by their concordance with trondhjemite gneissosity in areas of high strain, and because they are often assimilated in the same way as mafic xenoliths in the region. These dykes may well represent an intrusive, magmatic event that was coeval with the extrusion of the dominantly tholeiitic lavas in the upper (and younger) portions of the Onverwacht pile in the Barberton greenstone belt. Like the Ameralik dykes of West Greenland, they are not considered to have been feeders to the greenstone assemblages, and the sial which they intrude cannot, therefore, be considered as basement upon which the ensimatic crust was deposited. ACKNOWLEDGMENTS This work was undertaken as part of South Africa's contribution to the International Geodynamics Project and the financial assistance of the Council for Scientific and Industrial Research is acknowledged. The constructive criticisms of Dr L. Bettenay are gratefully acknowledged and Professor C. R. Anhaeusser kindly read an earlier version of this paper. Mrs L. Tyler, Mrs D. Amaler and Mrs W. Job are thanked for secretarial and drafting assistance. , 1981: Magmatic cycles and the evolution of the Archaean granitic crust in the eastern Transvaal and Swaziland. Spec. Pubis geol. Soc. Aust7, 457-467.
the Archaean granite-greenstone terrane south of the Boesmanskop syenite pluton, Barberton Mountain Land. Inform. Circ. econ. Geol. res. Unit,
ARNDT, N . T . , NALDRETT, A . J . , & PYKE, D . R . , 1977:
ANHAEUSSER, C . R., & ROBB, L. J., 1978: Regional and
ASH WORTH, J. R., 1976: Pedogenesis of migmatites in the Huntley-Portsoy area, north-east Scotland.
Univ. Witwatersrand, Johannesburg, 143.
detailed field and geochemical studies of Archaean trondhjemitic gneisses, migmatites and greenstone xenoliths in the southern part of the Barberton Mountain Land, South Africa. Inform. Circ. econ.
Geol. res. Unit, Univ. Witwatersrand, Johannesburg, 125.
Komatiitic and iron-rich tholeiitic lavas of Munro Township, Northeast Ontario. J. Petrol., 18, 319-369.
Mineral. Mag., 40, 661-682.
M. JR., 1981: The pattern of Archaean crustal evolution in southern Africa as deduced from the evolution of the Limpopo Mobile Belt and the Barberton granite-greenstone terrain. Spec.
BARTON, J .
Pubis geol. Soc. Aust., 7, 21-31.
MIGMATITE SOUTHWEST OF BARBERTON CHARLESWORTH, H . A . K . , & EVANS, C . R . , 1962:
Cleavage-boudinage in Precambrian rocks of Jasper, Alberta. Geol. Mijn., 41, 3 5 6 - 3 6 2 . W., 1968: Metamorphosed and deformed pegmatites and basic .dykes in the Lewisian complex of the Outer Hebrides and their geological significance. Q.J. geol. Soc. Lond., 123, 3 5 3 - 3 7 8 . FRATTA, M., & SHAW, D. M., 1974: "Residence" contamination of K, Rb, Li and T1 in diabase dikes. Can. J, Earth Sci., 11, 422-429. GILL, R. C. O., & BRIDGWATER, D., 1979: Early ; Archaean basic magmatism in West Greenland: the geochemistry of the Ameralik dykes. J. Petrol., 20, 695-726. HEDGE, C, E., 1972: Source of leucosomes of migmatites in the Front Range, Colorado. Mem. geol. Soc. Am., 135, 6 5 - 7 2 . KING, B. C., 1965: The nature and origin of migmatites: metasomatism or anatexis; in Pitcher, W. S., & Flinn, G. W. (Eds) Controls of Metamorphism. Geol. J. spec. Issue, 1, 2 1 9 - 2 3 4 . KLEEMAN, A . W . , 1965: The origin of granitic magmas. J. geol. Soc. Aust., 12, 3 5 - 5 2 . LOWMAN, P. D., 1965: Non-anatectic migmatites in Gilpin County, Colorado. Bull. geol. Soc. Am., 76, 1061-1064. DEARNLEY, R . , & DUNNING, F .
LUTH, W . C . , JAHNS, R . H . , & TUTTLE, O . F . , 1964:
The granite system at pressures of 4 to 10 kilobars. J. geophys. Res., 69, 759-773.
MACKENZIE, D. H., 1957: On the relationship between
migmatization and structure in mid-Strathspey. Geol. Mag., 94, 117-186. MISCH, P., 1968: Plagioclase compositions and nonanatectic origin of migmatitic gneisses in Northern Cascade Mountains of Washington State. Contrib. Mineral. Petrol., 17, 1-70. MYERS, J. S., 1978: Formation of banded gneisses by deformation of igneous rocks. Precamb. Res., 6, 43-64. OOSTHUYZEN, E. J . , 1970: The Geochronology of a Suite of Rocks from the Granitic Terrain surrounding the Barberton Mountain Land. Ph.D. Thesis, Univ. Witwatersrand [unpublished].
349
1956: The origin and significance of boudinage. Geol. Mag., 93, 401-408. SMITH, H. S . , & ERLANK, A. J., 1978: Aspects of the geochemistry of Archaean volcanic rocks: Lower Ultramafic Unit, Swaziland Sequence, South Africa; in Smith, I. E. M., & Williams, J. G. (Eds) Archaean geochem. Conf. Abs., Univ. Toronto, Ontario. SUN, S - S . , & NESBITT, R . W . , 1978: Pedogenesis lof Archaean ultrabasic and basic volcanics: evidence from rare earth elements. Contrib. Mineral. Petrol., 65, 3 0 1 - 3 2 5 . VAN SCHMUS, W. R . , & ANDERSON, J. L., 1977: Gneiss and migmatite of Archaean age in the Precambrian basement of central Wisconsin. Geology, 5, 45-48. RAST, N . ,
VILJOEN, M . J . , & VILJOEN, R . P . , 1969A: T h e g e o -
chemical evolution of the granitic rocks of the Barberton region. Spec. Pubis geol. Soc. S. Afr., 2, 189-219.
, 196%: The geology and geochemistry of the Lower Ultramafic Unit of the Onverwacht Group and a proposed new class of igneous rocks. Spec. Pubis geol. Soc. S. Afr., 2, 55-86. VON PLATEN, H . , 1965: Experimental anatexis and genesis of migmatites; in Pitcher, W. S., & Flinn, G. W. (Eds) Controls of Metamorphism. Geol. J. spec. Issue, 1, 2 0 3 - 2 1 8 . WHITE, A. J . R . , 1966: Genesis of migmatites from the Palmer region of South Australia. Chem. Geol., 1, 165-200.
WINKLER, H . G . F . , & LINDEMANN, W . , 1972: T h e
system Qtz-0r-An-H 0 within the granite system Qtz-0r-Ab-An-H20: application to granitic magma formation. N. Jb. Mineral Mh., (H4), 49-61. 2
WINKLER, H . G . F . , & BREITBART, R . , 1978: N e w
aspects of granitic magmas. N. Jb. Mineral. Mh.,
( H 10), 4 6 3 - 4 8 0 .
YARDLEY, B. W. D . , 1978: Genesis of the Skagit gneiss
migmatites, Washington, and the distinction between possible mechanisms of migmatization . Bull, geol. Soc. Am., 89, 9 4 1 - 9 5 1 .
THE FISKENAESSET ANORTHOSITE COMPLEXA STRATIGRAPHIC KEY TO THE TECTONIC EVOLUTION OF THE WEST GREENLAND GNEISS COMPLEX 3000-2800 M.Y. AGO John S. Myers Geological Survey of Greenland, Oster Voldgade 10, 1350 Copenhagen, Denmark ABSTRACT The Fiskenaesset Complex is a layered sequence of igneous cumulates 500 m thick, comprising anorthosite, leucogabbro, gabbro, peridotite, dunite and chromitite. It is interlayered with granitoid gneiss which veins it and makes up 80% of the region. The igneous stratigraphy and small-scale way-up structures of the anorthosite complex provide a framework by which it can be seen that the granitoid gneiss was intruded as sheets associated with thrusting, sub-parallel with the. igneous layering of the anorthosite complex. Large, recumbent, nappe-like folds of these rocks were refolded by two sets of folds with steep axial surfaces perpendicular to each other. Metamorphism reached granulite facies as tectonic activity waned and the region became part of a stable craton. The combined intrusion of granitoid sheets with thrusting and nappe-like folding led to considerable crustal thickening and appears to have been a major process in the rapid formation and stabilisation of large regions of sialic crust 3000-2800 m.y. ago.
INTRODUCTION Inclusions of metamorphosed anorthosite and leucogabbro are widespread throughout the Archaean quartzo-feldspathic gneiss complex of Greenland (Bridgwater et al., 1976). They are especially abundant in the Fiskenaesset region (Fig. 1) where they have been derived from a single major layered intrusion called the Fiskenaesset Complex. This complex occurs over an area of 5000 sq km as layers and trains of inclusions in a terrain of amphibolite- and granulitefacies gneisses. Although it is fragmented by granitoid intrusions and repeatedly deformed and metamorphosed, the layered anorthosite complex extensively preserves relic igneous textures and structures, and a detailed stratigraphy. It provides stratigraphic marker horizons and way-up structures, such as are rarely seen in Archaean gneissose terrains, that can be used to interpret the intrusion mechanisms of granitoid gneisses and the regional structure and tectonic history of an otherwise typical segment of the Archaean gneiss complex of Greenland. STRATIGRAPHY OF THE FISKENAESSET COMPLEX The original igneous stratigraphy of this layered intrusion is disrupted by thrusting and the associated intrusion of granitoid sheets and veins, and is distorted by deformation. The complete stratigraphy is not exposed at any one locality
Spec. Pubis geol. Soc. Aust., 7 (1981)
and its primary thickness is unknown, but a generalised stratigraphic column (Fig. 2) can be constructed from a large number of outcrops of parts of the stratigraphy where subsequent disruption and deformation are minimal. Most rocks are completely or partly recrystallised from igneous assemblages but, for simplicity in description, igneous names are used without the prefix "meta" regardless of the degree of metamorphic recrystallisation. Most rocks consist of metamorphic plagioclase and hornblende: but where igneous assemblages have survived they show that the ultramafic rocks are olivineorthopyroxene-clinopyroxene-spinel cumulates, the gabbros are olivine-plagioclase cumulates with corona structures of orthopyroxene-clinopyroxene-amphibole-spinel-garnet between cumulus olivine and plagioclase (Myers & Piatt, 1977), and the leucogabbros and anorthosites are plagioclase cumulates. A variety of igneous structures provide evidence of the original way-up of the stratigraphic succession. They include mineral-graded layering (Fig. 3), size-graded layering, crossbedded layering, trough layering (Fig. 4), columns of trough layers, comb layering (Fig. 5) and slump structures (Fig. 4). In addition certain unique kinds of deposits provide marker horizons within some of the major stratigraphic units. The most spectacular of these deposits is a 10 m thick layer of plagioclase-aggregate cumulate at the base of the Upper Leucogabbro Unit. This con-
F2 , Granitoid gneiss
H
U p p e r ->
LowerJ
r Fiskenaesset
Complex
If.. 1 A m p h i b o l i t e F 2 A x i a l trace F 3 A x i a l trace
Fig. 1.
Simplified geological map and cross-section of the Fiskenaesset region.
FISKENAESSET ANORTHOSITE COMPLEX UPPER GABBRO UNIT
ooo
oooo
oooo oooo
ANORTHOSITE UNIT OOQOOO OOO O^OOOOOOOO O 0
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o o»o o
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MIDDLE GABBRO UNIT
LOWER LEUCOGABBRO UNIT
ULTRAMAFIC UNIT 100 m LOWER GABBRO UNIT
" J Anorthosite
l l Chromitite xxxxx
loo ooi
Leucogabbro coarse grained | Leucogabbro fine grained | Gabbro Ultramafic rocks
Fig. 2. General stratigraphy of the Fiskenaesset Complex compiled from a large number of measured incomplete sections.
sists of 20 cm diameter aggregates of plagioclase crystals 1-5 cm in diameter in a matrix of plagioclase and hornblende (Fig. 6). It is interlayered with trough-shaped ultramafic channel deposits, etch about 1-3 m thick, 100-500m wide and of great but unknown length. These trough layers consist of olivine-orthopyroxene-clinopyroxenespinel-chromite cumulates which were probably
353
deposited from turbidity currents (Myers, 1976a). Similar ultramafic trough layers occur sporadically throughout the Upper Leucogabbro, Middle Gabbro and Lower Leucogabbro Units but are only abundant and interlayered with plagioclaseaggregate cumulates at the base of the Upper Leucogabbro Unit. Smaller trough layers of plagioclase-chromite cumulate occur scattered throughout the Anorthosite Unit in the central part of the Fiskenaesset region (the region shown on Fig. 1), but are of less stratigraphic use. However, in the western part of the Fiskenaesset region (Fig. 1) chromite is mainly concentrated in a few (between 1 and 6) more extensive layers, each generally between 50 cm and 3 m thick, which occur in the upper part of the Anorthosite Unit at the bottom of a sub-unit of garnet anorthosite (Windley et al., 1973; Ghisler, 1976). Other features which can be used as stratigraphic markers within some of the major units are plagioclase-magnetite-ilmenite cumulates which are restricted to the Upper Gabbro Unit and are especially abundant in the lower part of that unit, and a 10 m thick horizon of plagioclasechromite cumulate at the top of the Upper Leucogabbro unit. The upper part of the Middle Gabbro Unit is characterised by vertical columns of well-defined mineral-graded trough layers in uniform gabbro (Fig. 3), in contrast to poorly defined centimetre-scale mineral-graded layers of great lateral extent in uniform gabbro in the lower part of that unit. In the central and eastern part of the Fiskenaesset region the whole 60 m thick Upper Leucogabbro Unit is size-graded with cumulus plagioclase increasing in size upward from 2cm to 10cm in diameter (Fig. 7). In contrast, cumulus plagioclase in the Lower Leucogabbro Unit is generally 1-5 cm in diameter, but some layers contain mixed grain sizes of cumulus plagioclase between 1 mm and 2 cm in diameter, and some layers show sizegrading with plagioclase decreasing in size upward from 5-10cm to less than 1 cm in diameter (Fig. 8). The combination of the overall igneous stratigraphy of 7 major units, the marker horizons and other distinctive features within some of these units, and the small-scale way-up structures provides a detailed stratigraphic framework with which to interpret the intrusion mechanism of the granitoid gneisses and the structure and tectonic history of the region. The Fiskenaesset Complex was intruded as a sheet into basic volcanic rocks, which are now amphibolites, before metamorphism about 2800 m.y. ago (Alexander et al., 1973; Black et al., 1973). Some amphibolites preserve pillow lava, pillow-lava breccia and pyroclastic structures
354
Fig. 3.
J . S. M Y E R S
Mineral-graded layering in u n i f o r m g a b b r o of the Middle G a b b r o Unit. An u n d e f o r m e d but completely recrystallised rock of hornblende, diopside and calcic plagioclase. All the rocks illustrated by p h o t o g r a p h s consist of similar metamorphic mineralogy with cumulus crystals completely or partly recrystallised to m e t a m o r p h i c aggregates and, unless stated otherwise, are f r o m M a j o r q a p qava (Fig. 12). The h a m m e r is 28 cm long. Original way-up toward the t o p of the p h o t o g r a p h .
Fig. 5.
Comb-layering marked by elongate h o r n b l e n d e crystals. Original way-up toward the left. T h e pen is 13 cm long.
Fig. 6.
Layers of plagioclase aggregate cumulate in u n i f o r m g a b b r o at the base of the Upper L e u c o g a b b r o Unit. Original way-up toward the top right.
(Escher & Myers, 1975), and are associated with thin siliceous metasediments. All these rocks were deformed before the emplacement of granitoid intrusions, now gneisses, which make up 80% of the region. G R A N I T O I D SHEETS AND THRUSTING
Fig. 4.
Mineral-graded trough layering filling a large channel in the Middle G a b b r o Unit. N o t e the magmatic slump structure to the right of the h a m m e r . Original way-up toward the right.
The granitoid intrusions split the anorthosite complex and metavolcanic amphibolites into thin layers and trains of inclusions (Fig. 1), some of which can be followed over well-exposed ground for 30 km or more even when they are no more than 1 m wide. Most thin layers and inclusion trains are parts of individual stratigraphic units
FISKENAESSET ANORTHOSITE COMPLEX
Fig. 7.
Leucogabbro with mainly 10 cm diameter cumulus plagioclase, typical of the upper part of the Upper Leucogabbro Unit.
of the anorthosite complex rather than mixtures of different parts of the anorthosite stratigraphy. The granitoid intrusions were thus emplaced as sheets, sub-concordant with the igneous stratification of the anorthosite complex (Fig. 9), rather than as large discordant plutons. The widespread evidence of brittle fragmentation of the amphibolites and anorthosite complex during the emplacement of the gneiss, the lack of major chemical reaction between fragments of amphibolite and anorthosite and the enclosing
Fig. 9.
Fig. 8.
355
Size-graded layering shown by cumulus plagioclase in leucogabbro in the Lower Leucogabbro Unit. The pen is 13 cm long and is parallel to the igneous layering. Original way-up toward the top right.
gneiss, and the presence of gneiss sheets separating previously adjacent layers of anorthosite leave little doubt that most of the gneiss was intrusive and was not derived by metasomatic alteration or remobilisation of older sedimentary, volcanic or plutonic layers at the level now exposed. In many places where the granitoid sheets cut at low angles across the stratified anorthosite, it
Schematic section showing the magmatic-tectonic stratigraphy in the Fiskenaesset region before the formation of F1 recumbent folds. It illustrates how the granitoid gneiss occurs as intrusive sheets associated with thrusts, sub-parallel to the igneous layering of the Fiskenaesset Complex (shown with three hypothetical divisions A, B and C) and bedding of metavolcanic amphibolite.
J. S. MYERS
anorthosite was rigid. Angular fragments of anorthosite were detached from both sides of the granitoid sheets and were disorientated. The granitoid sheets are strongly deformed and have a planar fabric parallel with their margins. This evidence, together with the composition of these granitoid rocks and the scattered preservation of relic igneous textures in them (Myers, 1978), suggests that much of the quartzofeldspathic gneiss of the region was intruded as sheets of tonalite, grariodiorite or granite, associated with thrusting. This mode of emplacement was first recognised in Greenland by McGregor (1973) in the Godthib region, 150 km north of Fiskenaesset, where sheets of 3000-280Qm.y. old Nuk gneiss are interlayered with metavolcanic amphibolite and 3800-3600 m.y> old Amitsoq gneiss. The importance of this sub-horizontal tectonic regime was elaborated by Bridgwater et ai (1974) and Myers (19766), and further evidence was presented from the GodtMb region by Hall & Friend (1979).
t-euc°9abbro
Middle
Upper
356
Lower Leucogabbro Unit
.
Middle Gabbro Unit
.
Upper Leucogabbro Unit
.
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Granitoid gneiss
Leucogabbro
REGIONAL FOLD STRUCTURES The igneous layering of the Fiskenaesset Complex and sub-parallel layers of amphibolite and granitoid gneisses were folded into large-scale, recumbent isoclinal F1 structures (Fig. 1). The known stratigraphy and way-up of the anorthosite layers enables the facing directions of these folds to be determined (Fig. 11). A large-scale example occurs at Majorqap qava (Fig. 12) where the oldest fold structure is an isoclinal syncline which is refolded by a tight F2 synform. The F1 syncline is downward-facing on the northwest side, and upward-facing on the southeast side of the F2 synform. A similar F1 synclinal structure
Gabbro Gabbro with mineral-graded layering Ultramafic rocks
Fig. 10. Measured stratigraphic section across the lower part of the Fiskenaesset Complex on the summit crag of Majorqap qava. A relatively intact section of the Lower Leucogabbro Unit, Middle Gabbro Unit and lower part of the Upper Leucogabbro Unit, showing a number of granitoid sheets intruded parallel to the cumulate layering, and a tectonic repetition of the upper part of the Middle Gabbro Unit with a thrust plane along a thin granitoid sheet.
can be seen that the granitoid sheets were emplaced along active thrust zones which cut out or duplicated parts of the anorthosite stratigraphy (Figs 9 & 10). When these movements occurred, the granitoid sheets were ductile and the adjacent
Fig. 11. A minor downward facing F1 anticline in gabbro near the hinge region of the major F1 syncline at M a j o r q a p qava. The hammer shaft is parallel to a pronounced F1 schistosity, subparallel to the mineral- graded layering to the left.
FISKENAESSET ANORTHGSITE COMPLEX
Fig. 12. Simplified geological map and section of the Fiskenaesset Complex at Majorqap qava.
357
358
J. S. MYERS
was postulated by Windley et al. (1973) for the more deformed anorthosite layers north of Fiskenaesset (Fig. 1), based on the symmetrical repetition of the overall stratigraphy and using geochemical trends as indicators of the original way-up. The F2 synform at Majorqap qava extends westward to the sea north of Fiskenaesset and eastward to the inland ice cap, and most layers of anorthosite associated with this structure are part of a major Fl syncline. In contrast, most anorthosite layers in a second major belt of anorthosite southeast of Fiskenaesset and Majorqap qava are part of a single stratigraphic sequence which is inclined to the southeast and is the right way-up. This southern belt of anorthosite outcrops could be interpreted as the lower limb of the major recumbent F1 syncline of Majorqap qava (Fig. 1 section). In detail the structure is much more complicated because of the disruption of the anorthosite stratigraphy by granitoid intrusions and thrusting before Fl, and later refolding of the recumbent Fl folds by major F2 and F3 folds with steep axial surfaces at high angles to each other (Fig. 1). These F2 and F3 folds produced large scale dome-and-basin interference patterns and the main structural grain of the region (Fig. 1). Macroscopic particles such as pillow-lava structures in amphibolites, equant cumulus plagioclase in anorthosite (Fig. 13), and porphyritic textures, amphibolite fragments and vein networks in the granitoid gneisses, provide measures of both local and regional variations in the intensity of deformation (Myers, 1978), but metamorphism generally outlasted deformation and led to xoarse-grained equigranular textures and the obliteration of microscopic tectonic fabrics. Metamorphism locally reached granulite facies toward the end of the tectonic activity, and F3 deformation is more brittle and concentrated in shear zones and faults in the region of granulite-facies metamorphism near Fiskenaesset than in the region of amphibolite-facies metamorphism to the east.
Fig. 13. Fl schistosity (centre) developing from chromite-bearing leucogabbro with equant cumulus plagioclase (right), sub-parallel to a mineral-graded layer (left). Northeast of Fiskenaesset; original way-up toward the top right.
1976), and may be much more common in other Archaean terrains than is yet recognised. The intrusion of enormous volumes of granitoid sheets associated with thrusting and recumbent folding appears to be a major process during one of the most marked episodes of continental growth on Earth 3000-2800 m.y. ago. The Fiskenaesset region is 150 km south of a 75 km wide zone in which 3800-3600m.y. old sialic crust, the Amitsoq gneisses (McGregor, 1973, 1979), is extensively preserved. Layers of anorthosite and amphibolite similar to those of the Fiskenaesset region are interleaved with the Amitsoq gneisses and all are intruded by sheets of 3000-2800m.y. old Nuk gneisses (McGregor, 1973, 1979). It is not known whether these metavolcanic amphibolites and the anorthosites originally accumulated on an older sialic crust of Amitsoq gneisses, or represent a portion of oceanic crust or island arc that was technically IMPLICATIONS FOR LATE ARCHAEAN interleaved with the Amitsoq gneisses during colCONTINENTAL GROWTH lision between an Amitsoq gneiss continent and The detailed stratigraphy of the anorthosite oceanic crust or island arc 3000-2800m.y. ago complex provides a framework for recognising (Fig. 14): the scanty evidence is controversial that the granitoid gneisses which form 80% of the (McGregor, 1973; Chadwick & Coe, 1976), If the Fiskenaesset region were intruded as sheets sub- interleaving of these rocks is the result of plate concordant with the igneous layering of the collision then this collision process was different anorthosite complex and bedding of the meta- from recent subduction zones in that the oceanic volcanic host rocks. This mode of emplacement or island-arc crust (now amphibolite and appears to apply to much of the 3000-2800 m.y. anorthosite) was extensively interleaved with old gneisses of Greenland (Bridgwater et al., sialic crust (Amitsoq gneisses) and new (3000-
FISKENAESSET ANORTHOSITE COMPLEX
^
Mantle' . " .
M
Basalt
•
Anorthosite
OCQ]
Gneiss
Basicmagma
|
1
Granitoid magma
Fig. 14. Schematic section showing a model for the interleaving of oceanic and continental crust with granitoid magma, here derived by subduction processes below the edge of the continent during an early stage of plate collision about 3000 m.y. ago. 2800m.y. old) magmas rather than mostly being subducted and melted. The Fiskenaesset Complex was intruded into volcanic rocks which were deposited in water at the Earth's surface. These rocks do not show evidence of partial melting and so were probably not carried to great depth ( > 3 0 km) during any collision with sialic crust such as the Amitsoq gneisses. There is n o sign of marked chemical reaction between fragments of amphibolite and anorthosite in the granitoid gneisses, and so the granitoid intrusions were probably emplaced rapidly and at fairly shallow depth (probably <15 km). The F1 recumbent folds with amplitudes of tens of kilometres which predate highgrade metamorphism could not easily have formed with steep axes, and probably also formed at shallow depth as nappes. These folded rocks underwent prograde metamorphism to a maximum temperature and pressure of about 800°C and 10kb (Myers & Piatt, 1977; Wells, 1979), suggesting that they were then buried to a maximum depth of a b o u t 30 km. This temperature and pressure marks the peak of metamorphism and probably indicates the maximum depth of burial. The lack of chemical reaction between the amphibolite, anorthosite and gneiss suggests that any subduction process was shallow and rapid and was followed by rapid uplift and cooling. Phanerozoic intrusions of similar composition are most voluminous at convergent plate margins, and Windley & Smith (1976) suggest that the Archaean rocks of the Fiskenaesset region are ancient equivalents of the epizonal Mesozoic-Tertiary batholiths of the western Americas. The f o r m of most granitoid intrusions in the two regions is however quite different, being thin sheets in the Archaean of Greenland and mainly large globular or rectangular plutons in the western Americas. Most granitoid intru-
sions in the Fiskenaesset region were intruded as sub-horizontal sheets during sub-horizontal tectonic movements, whereas most exposed epizonal granitoid intrusions of the western Americas were intruded vertically as diapirs or by cauldron subsidence. The granitoid intrusions emplaced along convergent plate margins such as those of the western Americas form narrow linear belts. If the extensive areas of granitoid gneisses in Greenland which indicate m a j o r late Archaean continental growth formed in a similar way it might be expected that these gneisses would consist of a large number of linear belts of slightly different rock types and ages. But most late Archaean gneisses in Greenland show a monotonously similar range of rock types and ages, and do not appear to have formed in narrow linear belts. McGregor (1979) suggests that the late Archaean Nuk gneisses of West Greenland formed by partial melting of basaltic material above a zone of mantle downwelling but, unless this downwelling was patchy, it might also be expected to have produced linear belts of granitoid gneiss. The environment which produced the late Archaean rocks of Greenland is still a subject of speculation, but perhaps the main difference with Proterozoic and Phanerozoic granitoid rocks reflects a more vigorous mobility of plates and a higher geothermal gradient during the Archaean. This could have led to shallower subduction, much more extensive interleaving of plates beyond plate margins, and the associated production and intrusion of granitoid magmas over extensive areas in an environment of subhorizontal tectonic movements. ACKNOWLEDGMENTS The paper is published with the permission of the Director of the Geological Survey of Greenland.
J. S. MYERS
360
REFERENCES
ALEXANDER, E . C . , EVENSEN, N . M . , & MURTHY, V . R . ,
1973: Ar- Ar and Rb-Sr studies of the Fiskenaesset complex, West Greenland [Abstr.]. Trans. Am. 40
39
geophys. Un., 54 (11), 1227.
BLACK, L . P . , MOORBATH, S . , PANKHURST, R . J . , & WINDLEY, B. F., 1973 : 2 0 7 P b / 2 0 6 P b whole rock age
of the Archaean granulite facies metamorphic event in West Greenland. Nature, Phys. Sci., 244, 50-53.
BRIDGWATER, D . , MCGREGOR, V . R . , & MYERS, J . S . ,
1974: A horizontal tectonic regime in the Archaean of Greenland and its implications for early crustal
thickening. Precamb. Res., 1, 179-197.
BRIDGWATER, D . , KETO, L . , MCGREGOR, V . R . , & MYERS, J . S., 1976: Archaean gneiss complex of
Greenland; in Escher, A., & Watt, W. S. (Eds) Geology of Greenland, 18-75. Grdnlands Geol. Unders., Copenhagen. CHADWICK, B., & COE, K., 1976: New evidence relating to Archaean events in southern West Greenland; in
Windley, B. F. (Ed.) The Early History of the
Earth, 203-211. Wiley, London.
ESCHER, J . C., & MYERS, J . S., 1975: New evidence con-
cerning the original relationships of early Precambrian volcanics and anorthosites in the Fiskenaesset region, southern West Greenland. Rapp.
Grdnlands geol. Unders., 75, 7 2 - 7 6 . GHISLER, M., 1976: The Geology, Mineralogy and Geo-
chemistry of the Pre-orogenic Archaean Stratiform Chromite Deposits at Fiskenaesset, West Greenland. (Monograph Series on Mineral Deposits, 14).
Gebriider Borntraeger, Berlin.
HALL, R . P . , & FRIEND, C. R . L . , 1979:
Structural evolution of the Archaean rocks in Ivisartoq and the neighboring inner GodthSbsfjord region, southern West Greenland. Geology, 7, 3 1 1 - 3 1 5 .
1973: The early Precambrian gneisses of the GodthSb district, West Greenland.
MCGREGOR, V . R . ,
Phil. Trans. R. Soc. Lond.,A273,
343-358.
, 1979: Archaean gray gneisses and the origin of the continental crust: evidence from the GodthSb region, West Greenland; in Barker, F. (Ed.)
Trondhjemites, Dacites, and Related Rocks,
169-204. Elsevier, Amsterdam. 1976a: Channel deposits of peridotite, gabbro and chrornitite from turbidity currents in the stratiform Fiskenaesset anorthosite complex, southwest Greenland. Lithos, 9, 281-291. , 1976Z?: Granitoid sheets, thrusting, and Archaean crustal thickening in West Greenland.
MYERS, J . S . ,
Geology, 4, 265-268.
, 1978: Formation of banded gneisses by deformation of igneous rocks. Precamb. Res., 6, 43-64. MYERS, J. S . , & PLATT, R. G., 1977: Mineral chemistry of layered Archaean anorthosite at Majorqap qava, near Fiskenaesset, southwest Greenland. Lithos, 10, 59-72. WELLS, P. R. A., 1979: Chemical and thermal evolution of Archaean sialic crust, southern West Greenland.
J. Petrol., 20, 187-226.
WINDLEY, B . F . , H E R D , R . K . , & BOWDEN, A . A . , 1973:
The Fiskenaesset Complex, West Greenland, part 1: a preliminary study of the stratigraphy, petrology, and whole rock chemistry from Qeqertarssuatsiaq. Bull. Gr(inlands geol. Unders106 (also
Meddr. Gr(inland, 196, 2).
WINDLEY, B. F . , & SMITH, J. V . , 1976: Archaean high
grade complexes and modern continental margins.
Nature, Lond., 260, 6 7 1 - 6 7 5 .
EVOLUTION OF THE SHAW BATHOLITH—AN ARCHAEAN GRANITOID-GNEISS DOME IN THE EASTERN PILBARA, WESTERN AUSTRALIA
L. F. Bettenay, M. J. Bickle, C. A. Boulter, D. I. Groves, P. Morant, T. S. Blake, & B. A. James Department of Geology, University of Western Australia, Ned lands, Western Australia 6009 ABSTRACT
The Shaw Batholith is a composite ovoid dome containing analogues of high-grade gneiss terrains, centrally-developed migmatites representing modifications of the gneiss terrain, several variably deformed intrusive granitoid suites, and later, discrete post-tectonic granites. Greenstone intercalations that can be traced from gneiss terrains into enveloping greenstone sequences are not xenoliths but were incorporated by "Dj" tectonic interleaving, probably by thrusting, followed by isoclinal, probably recumbent, " D " folding. Interleaving in one area at least apparently involved stratigraphically high greenstones. The batholith evolved by late-stage uplift of the gneiss terrains whose early " D ' - " D " structures are unrelated to the present margin. Regional fabric orientations result from superimposition on "Dj" and "D2" of late shear belts and upright " D " folding that was syn-post adamellite/granodiorite plutonism. Tectonic crustal overthickening during gneiss terrain formation may have directly caused the relatively high-pressure type metamorphism and the later initiation of domal uplift. This study illustrates the complexity and protracted tectono-thermal history of the Shaw Batholith and highlights the difficulties in resolving a priori relationships in granite-greenstone terrains, a prerequisite to conceptual models for Archaean crustal development. Previous models appear oversimplified or inappropriate in the light of new constraints imposed by this study. 2
1
,
2
3
INTRODUCTION Elongate to ovoid batholiths comprising gneisses, migmatites and intrusive granitoids characteristically dominate Archaean granitegreenstone terrains. Their origin and subsequent evolution is widely debated because of the important consequences for the character of early crust and the tectonic setting of greenstone belts. Fundamental problems in these terrains include the possible existence of pre-greenstone sial within the batholiths, the source of component granitoids, and the relative contributions of magmatic intrusion, solid-state diapirism, superimposed folding and late differential uplift to the present configuration of batholith-greenstone contacts. In this early report of a continuing study of the composite Shaw Batholith, Eastern Pilbara, we document the internal complexity of a multiplydeformed and subsequently modified gneiss terrain and its interaction with surrounding greenstones. Previous conceptual models for batholith emplacement and tectono-metamorphism of greenstones are examined and discussed. 1s. is shown that the evolution of this batholith was
Spec. Pubis geol. Soc. Aust., 7 (1981)
more protracted and complex than is portrayed in most models, and that critical relationships such as those between greenstones and the oldest granitoid rocks are far less easily demonstrated than some earlier studies imply. REGIONAL SETTING The Shaw Batholith has an exposed area of some 3000 km (Figs 1, 2), but the entire southern and parts of the northern margin are unconformably overlain by the Proterozoic Fortescue Group. Archaean greenstones of the Warrawoona Group form a discontinuous, high-grade, high-strain envelope that contrasts with weakly metamorphosed, relatively undeformed sequences characterizing much of the Eastern Pilbara (c/. Barley et al., 1979). Hickman & Lipple (1978) correlate the eastern marginal greenstone sequences with the lowermost Taiga Taiga Subgroup and those of the western contact with the stratigraphically higher Salgash Subgroup. The base of the Warrawoona greenstones has yet to be clearly recognized. Isotopic ages of 3520 ± 60 Ma (Sm-Nd whole-rock method, Hamilton et al1980) and 3453 ± 16 Ma (zircon U-Pb 2
362
L. F. BETTENAY, ET AL. INDIAN
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Fig. 1. Geology of the Eastern Pilbara showing location of the study area within the Shaw Batholith.
method, Pidgeon 1978tf) are considered to be depositional events for the lower part of the succession. GROSS FEATURES OF THE BATHOLITH Hickman & Lipple (1978) recognized major components of the batholith as (i) older "migmatites", (ii) relics of greenstone lithologies, (iii) the North Shaw Tonalite, (iv) 'late- to syn-tectonic" intrusive adamellite/granodiorite plutons of finegrained (Agl) and porphyritic (Agp) types, and (v) post-tectonic (Sn-Ta related) granites. The last appear to be a discrete late event, typically yielding 2600-2700 Ma whole-rock Rb-Sr ages (de Laeter & Blockley 1972; de Laeter et al., 1975), although Pidgeon (19786) reported "consistent"' muscovite Rb-Sr ages of 2830 ± 30 Ma for tin4
bearing pegmatites. "Migmatites" have yielded ages of 3417 ± 40 Ma (zircon U-Pb method; Pidgeon, 19786), 2936 ± 9 Ma (Pb-Pb mineral; Oversby, 1976) and 2951 ± 83 Ma (Rb-Sr wholerock; de Laeter et al., 1975), but intrusive plutons later than the migmatites but predating the posttectonic suite have yet to be dated. Our study has confirmed the complexity and composite nature of the Shaw Batholith. The distribution of major units is schematically shown on Figure 3 but it is emphasised that this scale of presentation requires some simplification: most regional contacts are transitional or intercalated, many discrete but volumetrically minor phases cannot be shown, and few outcrops contain a single homogeneous phase. Neglecting the posttectonic granites, the batholith can be conveni-
EVOLUTION OF SHAW BATHOLITH 363 mineral fabrics. Hornblende tonalites and quartz diorites (North Shaw Tonalite) dominate in the north, whereas to the south these rocks are transitional to variably recrystallized biotite (± hornblende) granodiorites (Fig. 3). Biotite-bearing leuco-adamellites and leuco-granodiorites form several major plutons and related, widespread minor phases disrupt the more mafic granitoids to produce "intrusive agmatites". The granitoids display a single foliation that is oriented subparallel to the regional contact, and a moderately to shallowly plunging mineral lineation. Compatible structures are developed in adjacent greenstones. The absence within the granitoids of mesoscopic folds deforming the foliation, or of evidence for more than one foliation-producing event, contrasts with the situation in the remainder of the batholith. In general, the intensity of deformation and recrystallization increases away from greenstone contacts. This is particularly evident near the northeastern contact, where a high-strain zone contains foliated and interbanded orthogneisses derived from "intrusive agmatites" and contrasts markedly with adjacent marginal zones where recognizable pillowed ocelli-bearing meta-basalts are in contact with weakly foliated igneoustextured granitoids. Greenstone enclaves, mainly mafic-ultramafic rocks, are minor but widely dispersed. NearFig. 2. Regional subareas of the Shaw Batholith contact areas of the North Shaw Tonalite (Fig. 3) described separately in the text, showing loca- contain partially assimilated small angular basic tion of the Cooglegong Adamellite, largest of xenoliths whereas larger greenstone masses are the post-tectonic granite plutons. interpreted to be isolated roof pendants from a flat-lying contact. Elsewhere, isolated elongate ently subdivided into four subareas (Fig. 2), bodies of greenstone, some demonstrably faultnamely: bounded, parallel the granitoid foliation both (i) Northern subarea of variably deformed, near and remote from the batholith margin. homogeneous intrusive granitoids including the EASTERN SUBAREA North Shaw Tonalite. (ii) Eastern subarea of gneisses and homoThis poorly exposed marginal zone has not yet geneous deformed granitoids. been examined in detail but laminated metasedi(iii) Central subarea of intrusive plutons and mentary units up to 200 m in thickness, comprismigmatites comprising variably foliated neo- ing marls, calc-silicates and pelites, occur at or somes and multiply deformed gneissic palaeo- near the batholith contact. Homogeneous, folisomes. ated porphyroblastic adamellite gneiss that occufiv) Western subarea of complex gneiss terrains pies a prominent structural embayment in the with local deformed intrusives. south of the mapped area (Fig. 3) is transitional away from the margin to foliated porphyritic NORTHERN SUBAREA adamellite that intrudes multiply-deformed, The Northern subarea is typified by relatively banded granodiorite gneisses. Immediately south homogeneous granitoid units mappable at the of the post-tectonic granite (Fig. 3), mafic granoregional scale. Hornfelsed intrusive or tectonized diorites intrude the greenstones although a variintrusive contacts occur against the predomi- ably developed, superimposed foliation has pronantly mafic-ultramafic marginal greenstone duced augen textures within the granitoids. These sequences, and near-contact granitoids typically deformed intrusive phases are apparently gradaretain igneous textures despite partial recrystal- tional to interbanded leucogranite-granodiorite lization and the superimposition of tectonic orthogneisses lying remote from the contact,
364
L. F. B E T T E N A Y , ET
AL. Post-tectonic granites Variably deformed intrusive granitoids Medium grained adamellite/ granodiorite (Agl type) Biotite-muscovite adamellite Porphyritic adamellite/ granodiorite (Agp type) Leucogranite-adamellite Mafic granodiorite Hornblende tonalite (North Shaw Tonalite) Migmatite (gneiss remnants shown diagram malic ally) Biotite granodiorite gneiss
V - +\ V - ( i + + Aa 30 + T^Tjt-^ )/ M + + + + + + + + + + + + + + •+ + + + + + + + IJLLA /,,++ + •+ + + + + + + + + . + + + + + + + + + H > 'VVJ X'7 + + + + + + + + + + + + + + + + + + + + + -%/ I "/C\ * + + ' + + + + + + + + ) \ ' "„ tc* +
+
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Banded hornblendic gneiss M P M \ Greenstone intercalations J Greenstone envelope
j
Batholith margin Foliation Mineral lineation r
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/
Schematic distribution o f m a j o r units within the Shaw Batholith, showing orientation o f main fabric elements and areas o f detailed m a p p i n g . The granitoids are listed in the legend in order o f inferred age with the youngest at the top.
whereas migmatites with foliated and recrystallized neosomes developed farther to the west. The contact is analogous to, and presumably a continuation of, the high-strain zone of the Northern subarea.
CENTRAL
SUBAREA
The Central subarea contains anatectic migmatites in which palaeosomes of granitoid gneiss, which are more widespread than can be shown on Figure 3, represent minor phases in most outcrops except where major intrusive plutons are
developed. Two contrasting, although locally transitional and intercalated, gneiss types occur: " h o m o g e n e o u s " gneisses of biotite granodiorite composition in which banding is defined by layerparallel intercalations of pegmatite-aplite in a compositionally uniform host dominate (Figs 5a, c, d). Complex hornblendic gneisses (Fig. 5b) contain similar concordant lamellae but the host is heterogeneous and comprises hornblendebearing granitoids of monzo-diorite to leucoadamellite compositions. Their regular and continuous banding is defined by alternate diffuse mafic and felsic lamellae. Concordant or boudin-
EVOLUTION OF SHAW BATHOLITH
365 o
| Quaternary
m
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Proterozoic dolerites
Metajaspillites and minor •••'A calc-silicate& Pelitic and quartzo-feldspathic schists
H Z Ultramafics |V V V
Amphibolites Foliated intrusive granodiorites, adamellites, pegmatites Biotite granodiorite gneiss j Banded leucogneisses Banded hornblende gneisses Major synform, antiform Minor folds ^
Generated fabric
N
Other lineations
D2
^
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Lineations
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Gross lithological layering Fabric Shear zone
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Axial trace Fabric
Geology and schematic cross-section of the Fairwig Well area. Location shown on Figure 3.
aged amphibolite layers occur particularly within zones dominated by distinctive clinopyroxenebearing hornblende monzonite gneisses. The migmatites contain neosomes of leucoadamellite or leuco-granodiorite, typically accompanied by abundant podiform pegmatites. Some externally derived magmatic components were probably introduced, but in situ melting, including total reactivation of concordant leucoeratic lamellae, also took place. Although neosomes have extensively intruded and disrupted adjacent gneisses, and wholesale obliteration of axial zones in some folds of the gneisses attests to localization of partial melting or intrusion, structural continuity is largely retained between gneiss remnants in the migmatites. However, recrystalHzation has normally produced granofelsic tex-
tures and loss of microscopic layer-parallel foliations in gneiss palaeosomes. ~ Several ages of migmatite development are suspected: foliated neosomes dominate the southern areas and in one locality these are demonstrably intruded by weakly foliated porphyritic granodiorite. Elsewhere, massive, aplitic to pegmatitic leuco-adamellite neosomes possibly record a younger partial melting event. Intrusive plutons comprise a major proportion of the Central subarea. Apart from the post- tectonic granites, Hickman & Lipple (1978) distinguished two types of massive to foliated biotite adamellite/granodiorite intrusions, namely, porphyritic (Agp) types, and fine-grained aphyric (Agl) types. The former comprise internally homogeneous regional units (Fig. 3) whereas the
366
Fig. 5.
L
F. B E T T E N A Y , ET
AL.
Gneisses of the Shaw Batholith. Scale in all p h o t o g r a p h s indicated by 0.5 m long hammer. (a) near-isoclinal " D 2 " fold in homogeneous granodiorite gneiss with a b u n d a n t concordant leucocratic lamellae. Mesoscopic " D 3 " folds oblique to banding occur elsewhere in the o u t c r o p (from Western subarea). (b) Complex hornblendic gneiss intruded by weakly foliated pegmatites ( f r o m Central subarea). (c) Disrupted, fine-graine v d hornblende tonalite dyke intruding h o m o g e n e o u s granodiorite gneiss. Note truncation of linear aplite dykes by tonalite, but intrusion of later, d i f f u s e pegmatite veins into tonalite ( f r o m Western subarea). (d) Strongly foliated linear dyke of Agl-type granodiorite cutting folded homogeneous granodiorite gneiss and concordant leucocratic lamellae. Later leucocratic veins in older gneiss are continuous with sub-parallel leucocratic lamellae within dyke ( f r o m Western subarea).
Agl-types only f o r m m a j o r plutons near the n o r t h e r n m a r g i n of the C e n t r a l s u b a r e a a n d to t h e s o u t h o f t h e m a p p e d a r e a ( F i g . 1). H o w e v e r , both types also occur as widespread minor
p l u t o n s a t o u t c r o p s c a l e (less t h a n 1 k m 2 ) . A g p types g e n e r a l l y p r e d a t e A g l - t y p e s a l t h o u g h transitional textural varieties also occur a n d definitive cross-cutting relationships are u n c o m m o n . Con-
EVOLUTION OF SHAW BATHOLITH 367 tacts with the surrounding gneisses are normally all intrude the gneisses. These intrusives are precomplex. served in places as broadly folded and disrupted Greenstone lithologies (dominantly amphibo- dykes at a high-angle to gneiss foliation (e.g. Fig. lites) abound in the migmatites where most 5c) but normally form subconcordant or lowsequences are extensively disrupted by weakly angle discontinuous lenses within gneisses. In foliated to massive pegmatites and leuco-granite some places (e.g Fig. 5d) heterogeneous late related to neosome phases. Hornfelsing and par- deformation has transposed pegmatites into tial retrogression are common but there is no parallelism to produce zones of well-banded evidence that the greenstones underwent partial "intrusive" gneiss cutting contorted older melting; rather, it appears that structurally con- gneisses. cordant greenstone intercalations were invaded Amphibolite and amphibole-chlorite bodies and disrupted by leucocratic melt phases derived forming of homogeneous en-echelon by anatexis of surrounding gneisses. Neither boudins orseries sub-concordant bifurcating lenses assimilation of greenstones nor contamination of within gneisses are considered to be early intruthe invading phases is prominent. sive dykes although as a consequence of repeated deformations cross-cutting relationships can WESTERN SUBAREA rarely be established. In one locality, a discordant amphibolite dyke cropping out discontinuously The Western subarea mainly comprises homogeneous biotite granodiorite gneisses (Fig. 3) with for some 200 m post-dates folded leucocratic bands within the gneiss, but is itself foliated and locally-dominant intercalations of pegmatoidleucogranite gneiss, porphyroblastic adamellite lineated, displays open folding, and is cut by gneiss, and hornblendic gneisses: the last form mylonitized granodiorite dykes. The amphibolite major units in the Fairwig Well area (Figs 3, 4). dyke extends to within 50 metres of the greenAbundant concordant greenstone intercalations stone contact and is represented in the contact typically compose major multi-lithological units zone by on-strike concordant lenses of amphiboup to 1 km in width and 20 km in length that high- lite within mylonitic gneisses. Pre-deformation contact relationships between light regional fold patterns in the gneisses. The Western subarea shows strong similarities to gneisses and greenstone intercalations of the gneiss belts in Archaean high-grade terrains Western subarea are equivocal, but there is no (Bickle et al1980) and this is particularly evi- evidence that the greenstones within the batholith dent from relationships in the Fairwig Well area originated as large xenoliths (cf. Glikson, 1979): neither the homogeneous granodiorite nor the as shown in Figure 4. Upright '03" folds deform complex hornblendic gneisses display intrusive a concordant, diverse sequence of gneisses and contacts contain blocky enclaves of greenstone greenstone intercalations. Normally layer- material. orSome leuco-gneisses exhibit severely parallel, moderate to strong foliations, that are modified, local intrusive contacts against hornalso deformed by " D " , are a composite fabric blendic gneisses, and possibly represent relathat can be locally resolved into "S ", related to tively late additions tothus this gneiss terrain. isoclinal folds, and " S j " parallel to layering and folded by " F " . Weakly foliated biotitemuscovite-garnet adamellites intruding the gneiss terrain display only D "-related structures REGIONAL PLUTONIC AND whereas later, minor Agl-type phases are STRUCTURAL CORRELATIONS An assessment of the development of the Shaw massive. Elsewhere in the Western subarea, regional Batholith requires that the sequence of granitoid fold patterns are difficult to recognize despite phases and structural patterns developed at the abundant mesoscopic folds within both gneisses outcrop scale be correlated throughout, particuand greenstone intercalations. Low-angle mylo- larly between the regional subareas. Such correlanites and shear zones are common and may relate tions have proved difficult during reconnaisto attenuated fold limbs within the gneisses. sance-level mapping because of the following: Superimposed foliations and recrystallization (i) The scarcity of undisputed intrusive contacts between phases that can be related to major fabrics in both gneisses and cross-cutting granicomponents of the batholith. toid phases frustrate structural correlation on a (ii) The limited overall compositional contrast regional scale. Despite this, local intrusive and between major phases in relation to their cross-cutting relationships are commonly preinternal heterogeneity, and the presence of served; pegmatite-leucogranitoid phases, finecompositionally similar phases of several grained Agl-type granodiorites, abundant horndifferent structural ages within the one outblende-bearing granodiorites, tonalites and crop. quartz diorites, and porphyroblastic adamellites 4
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L. F. BETTENAY, ET AL. 368 (iii) The difficulty in delineating regional fold METAMORPHISM patterns in the absence of distinctive marker Most greenstones marginal to and within the horizons within the gneisses, and in relating Shaw Batholith have undergone high-grade metaminor fold structures to fabric elements that morphism associated with heterogeneous highare commonly sub-parallel to rotated older strain in contrast with low-grade (prehnitefoliations, preferentially developed in some pumpellyite to greenschist) assemblages characoutcrop phases, and partly obliterated by re- terizing much of the Eastern Pilbara. Insufficient crystallization and superimposed deforma- data are yet available to resolve fully grade varition fabrics. ations throughout the batholith, variations in the Nevertheless, the following correlations appear "regional" grade of surrounding greenstones, or valid: the nature of their boundary with the envelope. (i) The concordant gneisses (and greenstone Narrow amphibolite-facies thermal aureoles intercalations) of the Western subarea, can characterize the northern contact, where be traced continuously from the Fairwig "regional" grade is greenschist facies (actinoliteWell area (Fig. 4) into the Central subarea, albite in mafic rocks). On the western margin, the where the greenstones and gneisses become high-grade envelope (up to 3-5 km wide) is in progressively disrupted by leucocratic places fault-bounded but apparently transitional phases. Hence the migmatites of the Central with "regional'' sequences near the Emerald subarea represent a modified portion of the Mine area (Fig. 3). Pelitic assemblages here gneiss terrain. record mid to upper amphibolite-facies condi(ii) Granodiorite/adamellite plutons of Agl- and tions (kyanite-staurolite-white mica, stauroliteAgp-types, that intrude migmatites in the garnet-biotite) and imply relatively higher metaCentral subarea, are represented near the morphic pressures (>5kb) than normally western margin by sub-concordant, highly reported from greenstone terrains (e.g. Archibald strained lenses within the gneisses. These et al., 1978; Jolly, 1978). marginal zones therefore reflect areas of Upper amphibolite-facies assemblages charachigher composite strain. terize greenstone intercalations in the Central and (iii) The conspicuous N-S to NE-SW trending Western subareas, with sillimanite-white micaupright folds are assumed to be of the same garnet-biotite in pelites, quartz-magnetitegeneration throughout the Western subarea. grunerite (rarely with orthopyroxene, clinoThese " D " folds (The " D " folds of Hick- pyroxene, and fayalite) in banded iron formaman, 1975) deform normally layer-parallel tions, and anthophyllite- or forsterite-bearing foliations that are a composite of " D f and ultramafic assemblages. Widespread partial " D " (see also Bickle et al., 1980). Evidence retrogression is commonly for deformation events prior to " D j " is at greenschist-facies related to shear zones or proximity to migmatite this stage tentative. neosomes. Although there may have been It has not yet been possible to derive an unequithermal peaks, maximum prograde vocal sequence relating all of the phases and multiple metamorphism post-dated the earliest foliationstructural events within the Shaw Batholith; critical relationships, such as the structural age of production but predated upright " D " folding. amphibolite dykes, are not yet sufficiently resolved. The relative age of granitoids intruding GRANITOID-GREENSTONE INTERACTION the gneiss terrains as shown on Figure 3 preMarginal relationships between the Shaw supposes that the few compositionally and petro- Batholith and enveloping greenstones are varilogically distinctive phases (e.g. Agp- and Agl- able. Thus hornfelsed, discordant intrusive contypes) are of the same structural age throughout tacts in parts of the Northern subarea contrast the batholith so that a composite of age relation- with tectonized contacts elsewhere. On the ships from many different outcrops can be western margin the contact is a major tectonic erected. Even this may be an over simplification feature marked either by prominent faults or by a (cf. Hickman & Lipple, 1978). For example, Agl- zone of mylonites that commonly display comtype phases in the Fairwig Well area post-date plexly re-folded and transposed laminae. Open congruent minor folds related to " D " whereas mesoscopic folds also occur in contact mylonites in the Emerald Mine area (Fig. 3, see also Bickle and rarely the greenstone contact is folded by et al., 1980) " D " structures deform all minor upright, generally shallowly-plunging folds with phases including sub-concordant dykes correlated partially attenuated common limbs. Large-scale with the Agl-type plutons. Isotopic data are upright " D " folds deform the contact in the essential to place absolute time references on the Emerald mine area (Fig. 3), and Bickle et al. plutonic/structural sequence when finally it is (1980) have demonstrated that the greenstone unravelled. 3
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EVOLUTION OF SHAW BATHOLITH 369 belts structurally underlie the gneisses in part. tions that record several deformation events. This Greenstone intercalations within the batholith renders difficult the compilation of regional folican be traced into the envelope in this area. A ation trends, and hence the structural pattern major " D " structure folding earlier foliations in over the batholith as a whole, but it is clear that both greenstones and gneisses is involved in an the complex local deformation sequences are difinterference fold with the upright " D " folds, ficult to reconcile with diapiric uprise of previand Bickle et al. present evidence that greenstones ously undeformed granitoids. Finally, the present were interleaved with gneisses within the batho- attitude of both contact and fabric elements near lith during an early period of sub-horizontal tec- the western margin is clearly the result of upright tonics involving thrusting and recumbent folding. N-S to NE-SW trending "D " folds that postFurther studies are required to establish whether dated adamellite-granodiorite intrusions and tie enveloping greenstones are continuous with caused reorientation of at least two earlier fabrics the main greenstone sequences, and to establish in both gneisses and greenstones (Bickle et al., by structural correlation and mapping in the 1980). Hence in this area marginal deformation intervening areas whether the conclusions of predominantly reflects superimposed folding. Bickle et al. apply to gneiss terrains throughout If the upright folding is considered to reflect the batholith. the doming event (see, for example Hickman, A distinctive suite of finely intercalated felsic 1975), then it is clear that the present form of the schists, including quartzites and calc-silicate Shaw Batholith must have developed late in terms rocks, occurs on many mylonitized contact zones of the evolution of internal components. This and as an abundant lithological association in the doming event must have been contemporaneous greenstone intercalations within the batholith. with, or after, the emplacement of adamelliteFurther detailed studies will test the possibility granodiorite plutons and was post-dated only by that these rocks represent highly strained, basal the post-tectonic "tin granites". Within the gneiss terrain, a major recumbent " D " structure sedimentary units to the greenstones. described by Bickle et al. (1980) faces towards the DEVELOPMENT OF THE BATHOLITH batholith rather than away from its crest indiIn previous models for development of Eastern cating that the development of this gneiss terrain Pilbara batholiths, both Glikson (1979) and occurred earlier and was unrelated to doming of Hickman (1975) appeal to diapiric mechanisms the Shaw Batholith: early nappe structures but the former invokes magmatic diapirs and the developed as a consequence of doming would latter solid-state density inversion related to the face away from the dome and towards interdomal troughs (e.g. Rutland & Nicholson, 1965). deposition of greenstones on sialic basement. Theoretical and model studies of diapirism Bickle et al. (1980) speculate that the vertical (e.g. Stephansson, 1972; Ramberg, 1973; Dixon, movements responsible for doming developed as 1975; Schwerdtner et al., 1978) predict steeply a consequence of density instability within a tecdipping foliations and radially plunging maxi- hnically overthickened crust; the relatively highmum elongation lineations at the margins of pressure metamorphism is consistent with this diapiric intrusions, progressively becoming near- model. Our evidence accords in general with the horizontal in the centre. Such considerations led Archibald et al. (1978) to conclude that synkine- sequence proposed by Hickman (1975). However, matic intrusive granitoids in the Yilgarn were the evidence cited by Hickman to infer that greenemplaced as predominantly solid-state diapirs stone intercalations within the batholith were whereas other studies (e.g. Snowden & Bickle, interleaved with gneisses before deposition of the 1976; Fyson, 1978) conclude that fabric elements main greenstone sequence (his Table 1), is now in some domal plutons developed as a conse- equivocal following the discovery of pre-doming fabrics in greenstones of the envelope in the quence of superimposed deformations. Several considerations indicate that the Emerald Mine area. Similarly, whereas Hickman development of the Shaw Batholith was not solely attributed metamorphism/major deformation to a consequence of diapirism. Firstly, near the doming, our evidence indicates that the earlier northern margin, the intensity of a single defor- deformations were more severe in terms of fabric mation event within granitoids increases away development, and that metamorphism peaked from the greenstone margin rather than towards before upright folding and doming. It is not yet clear to what extent our concluit. Furthermore the orientation of lineations is at high angles to the radial plunge predicted for sions regarding the development of the Shaw Batholith as a structural dome by late-stage uplift diapirism (see Fig. 3). Secondly, elsewhere in the batholith, different phases within a single out- apply to other major granitoid-gneiss domes of the Eastern Pilbara (e.g. see Fig. 1). However, crop commonly show contrasting fabric orienta2
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L. F. BETTENAY, ET AL. 370 one of us (B. James) has demonstrated that the (iv) Gneiss complexes are dominated by granodiorites: the lesser adamellitic and tonalitic Tambina Pluton, an elongate dome located gneisses occur in roughly equal proportions. within greenstones some 15 km west of the Most gneiss phases could not be directly Emerald Mine area (see Fig. 3), also displays generated by anatexis of mafic-ultramafic fabric orientations incompatible with predicted sequences without extensive subsequent fracstrain distribution in simple solid-state diapirism. tionation and total removal of residue from Irrespective of whether the Tambina Pluton can exposed profiles. be used as a small scale analogue of the Shaw Batholith, it is clear that in both, superimposed (v) Granitoid gneisses are widespread throughout central zones of the batholith. In condeformation has played a major role in greenstone-granite interaction. trast, major plutons of igneous-textured tonalites are confined to discrete marginal A contrasting model for the development of environments. Eastern Pilbara granitoid batholiths elaborated by Glikson (1979) commands discussion and can (vi) Discordant amphibolite dykes cutting gneisses have a compatible tectonobe summarized as follows: A primitive maficultramafic crust of 'lower greenstones" experimetamorphic history to the high-grade ences subsidence and crustal anatexis to yield the greenstone envelope, and hence cannot be oldest granitoids as near-liquidus to superheated feeders to a yet younger greenstone setonalite and trondhjemite magmas. These intrude quence. "lower" greenstones as "polydomal multi- (vii) If, as inferred by Glikson (1979), greenstone lobate" diapirs, leading to extensive assimilation metamorphism was directly caused by the of greenstones at inter-domal contacts marked by intrusion of vast quantities of near-liquidus xenolith screens; gneisses develop in marginal to superheated tonalite magma (at 1100° + ; high-strain zones. Following unconformable see Wyllie, 1977), assemblages characteristic deposition of "upper greenstones", crustal of intermediate to low pressure, and high to anatexis of "lower greenstone"-"tonalite/ very high temperature, should occur on the trondhjemite" complexes then yields a much margin and within greenstone intercalations. younger, post-tectonic, adamellite-dominated Such a prediction contrasts with our intrusive suite. evidence that metamorphism occurred under Evidence accumulated in the Shaw Batholith relatively high pressures (>5kb), but temdirectly conflicts with most aspects of Glikson's peratures did not exceed ca 750 °C, even proposals. For example: within greenstone intercalations near the centre of the batholith. (i) Demonstrably xenolithic greenstones occur Any assessment of the development of Aronly in massive granitoid phases near contact zones, or as disrupted intercalations in chaean granitoid batholiths, and their role in the younger intrusive suites that also contain evolution of granite-greenstone terrains, must take into account the composite nature and comgneiss enclaves. plex internal relationships of some batholiths, as (ii) Major greenstone intercalations are concor- documented here. Such batholiths are not simply dant with, and have a comparable tectonometamorphic history to, surrounding homogeneous, genetically related magma pulses emplaced during the one tectonic stage into gneisses. These intercalations represent tectonic slices which, together with gneisses, are greenstone stratigraphies, and evidence that extensively disrupted by younger granitoids, granitoid components (notably pegmatites) ingiving the false appearance of "xenolith- trude greenstones implies nothing concerning the screens". The present distribution of inter- relation between greenstones and major batholith calations results from superimposed major units. Similarly the presence of greenstone xenolate fold structures and does not relate to liths within igneous-textured tonalites has no recognizable intraplutonic contacts or to the bearing on the relative age of greenstones and the oldest granitoids in the crust unless it is estabbatholith margin. lished that coeval intrusive tonalites were paren(iii) The intercalations show no preferred spatial tal to the oldest gneiss components. association with migmatites, with areas of study (see also Bickle et al., 1980) implies more intense deformation, or with any par- thatOurthe development of Archaean graniteticular gneiss type. The adjacent gneisses greenstone and high-grade gneiss terrains does show no evidence of contamination, nor do not necessarily independently but that the the greenstones show any indications of par- latter may in thisoccur have developed as an early tial melting or the effects of assimilation by stage of greenstonecase and granite inter-reaction (ef. tonalitic phases. Gorman et al., 1978), irrespective of which of 4
EVOLUTION OF SHAW BATHOLITH
these components contains the oldest crustal units. Our data suggest that Pilbara greenstones were most likely deposited on sialic basement, relics of which occur within the gneiss complexes of the Shaw Batholith, but unequivocal supporting structural evidence has not yet been obtained. These gneiss terrains may contain intercalated components derived not only from sialic basement, but also from older supracrustal sequences, basal sediment units, greenstones and early intrusive granitoids, but the severity and complexity of deformation in this environment severely restricts reconnaissance-level attempts to decipher the original relationships between such gneiss units. High priorities for future work are structural studies in key areas to resolve which if any of the gneiss components preserves evidence of older deformation events. Geochemical data are required to assess likely source rocks for both intrusive granitoids and gneisses, and further isotopic data are required to provide a geochronological framework for batholith development and to constrain petrogenetic models. However, we stress that careful field control is essential in such a complicated terrain if geochronological studies are to be of any significance in establishing absolute time controls on major, recognizable events in crustal development. For example the interpretation of available ages for "migmatites" is unclear in terms of the diversity and complex origin for granitoid phases allocated to these broad rock suites (for example, compare Figure 3 with Hickman & Lipple, 1978). CONCLUSIONS Preliminary conclusions emerging from this study of the Shaw Batholith are: (i) The oldest events yet deciphered in this crustal segment involved interleaving of extensive granitoid rocks with supracrustal components including significant metasedimentary units. Subsequent, probably recumbent folding, and metamorphism of the interlayered sequence, formed a typical highgrade gneiss terrain in which the oldest components are equivocal. (M) The early gneiss terrain was extensive and the present batholith outline bears no relationship to the distribution of gneiss components or to the orientation and intensity of early deformation structures. The broad impression of deformation decreasing from margin to centre of the batholith is brought about by a combination of late-stage events. Partial melting occurred in central zones with the development of younger weakly deformed granitoid phases, and this resulted
371 in partial obliteration by recrystallization of tectonic fabrics in centrally-located gneisses. In addition, further deformation affected composite gneiss-intrusive granitoid complexes near some greenstone contacts. (iii) Syn- to post-"D " high-grade metamorphism, under apparently higher-pressure conditions than normally encountered in graniteterrains, affected gneiss terrains and greenstones in the envelope. (iv) The batholith is an elongate structural dome developed by uplift relative to the adjacent greenstone belts. However, lineation orientations and the spatial variation in deformation intensity on the northern margin do not conform with strain patterns and distribution predicted for solid-state diapirism. On the western margin, where structures within gneisses are composite, the upright folds post-date adamellite-granodiorite intrusives. Hence, the present domal geometry of this batholith was attained during a late-stage structural event. (v) Models for batholith formation involving deformation and metamorphism of greenstones as a consequence of the emplacement of magmatic tonalite diapirs (e.g. Glikson, 1979), are not applicable to the development of the Shaw Batholith. Similarly, suggestions that interfolding of granitoid and greenstones occurred prior to the deposition of the main greenstone succession (ef. Hickman, 1975) are now rendered equivocal by our data. Enveloping greenstones (Upper Warrawoona, Salgash Subgroup of Hickman & Lipple, 1978) on the western contact have undergone the deformations so far unravelled in the gneiss terrain, so that gneissgreenstone interleaving must have postdated greenstone deposition unless major as yet unrecorded, tectonic discontinuities or unconformities separate the high-grade envelope from the main greenstone succession. If such discontinuities are present then existing regional stratigraphic correlations within the East Pilbara must be reconsidered, and granitoid-greenstone interrelationships that are a direct consequence of stratigraphic position (e.g. Glikson, 1979) should be regarded as even more speculative. ACKNOWLEDGMENTS This study was funded largely by the Australian Research Grants Committee. T. S. Blake and P. Morant held University of Western Australia Research Studentships and B. A. James a Commonwealth Postgraduate Scholarship during the study. 2
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THE GEOCHEMISTRY OF THE MOUNT EDGAR BATHOLITH, PILBARA AREA, WESTERN AUSTRALIA R. Davy & J. D. Lewis Geological Survey of Western Australia, Mineral House, 66 Adelaide Terrace, Perth, Western Australia 6000 ABSTRACT The Mount Edgar Batholith consists of a large mass of granitoid rock, mainly granodiorite in composition but ranging from tonalite to adamellite, intruded by minor plutons of similar rocks of the same general age. The batholith is also intruded by the younger Moolyella Adamellite. Though the compositions of each minor pluton are different, they are all contained within the range of composition shown by the main granodiorite mass. Regional geochemical maps show widely divergent patterns between different components with no evidence for systematic fractionation or differentiation. The batholith probably formed as a nested group of plutons from a magma derived by hightemperature partial melting of a lower-crustal felsic granulite layer. Chemical trends in the batholith are independent of foliation and textural variety. The Moolyella tinfield, defined by haloes of high Li and F, is probably underlain by a stock of younger granitoid, partly exposed as the Moolyella Adamellite. Anomalous Li in the northeast of the batholith is unexplained.
INTRODUCTION This paper presents a summary of a geochemical study of the Mount Edgar Batholith in the Pilbara region of Western Australia. The batholith is ideal for providing an overview of an Archaean batholith since it is almost wholly enclosed by Archaean greenstone. It forms a discrete unit, uncomplicated by either the loss of a major portion of outcrop beneath overlying younger rocks, or by the merging of adjacent batholiths. The Mount Edgar Batholith is located immediately east of Marble Bar (lat. 119° 45 'E, long. 21° 10'S). The main economic interest in this batholith is the Moolyella tinfield, centred 13 km east-northeast of Marble Bar. Part of this study was an investigation of the economic potential of the remainder of the batholith. The batholith crops out on parts of four 1:250000 geological sheets. Regional mapping has been carried out by Hickman (1978—Nullagine), Hickman & Lipple (1978—Marble Bar), Hickman & Chin (1976—Yarrie) and Hickman (1977—Port Hedland). These maps have formed the geological basis for this geochemical study. The regional setting of the batholith has been described by Hickman (1975; in press). The Moolyella tinfield has been described in detail by Blockley (1975; 1980). In summary, the batholith is roughly rhombic in outline, with a maximum north-south dimenSpec. Pubis geol. Soc. Aust., 7 (1981)
sion of 52 km and a maximum east-west dimension of 55 km (Fig. 1). Its shape is in part structurally controlled, along the southeast and southwest margins, and in part controlled by cross-cutting intrusive relationships. Geologically the batholith appears to present the relatively simple situation of a large mass of granodioritic material into which are intruded a number of smaller granodioritic and adamellitic bodies. The most notable later intrusion is the Moolyella Adamellite, which is associated spatially with tin mineralization. Belts of amphibolitic xenoliths within the batholith are thought to represent, in part, the keels of greenstone belts. Most discrete bodies within the batholith discussed here have been identified as plutons in the field, and have been previously reported by Hickman (in press). In the discussion which follows all such bodies of rock are referred to as plutons. Age determinations on rocks of the Mount Edgar Batholith have been reported by de Laeter & Blockley (1972) and Hickman (in press). Granitic rocks near Mount Edgar Homestead were dated at 3279 ± 169 Ma. Gneissic rocks west of Moolyella, and the Moolyella Adamellite were dated at 3125 ± 366 and 2670 ± 95 Ma respectively. Limited chemical analyses of Pilbara granitoids, including a small number from the Mount Edgar Batholith, have been reported by Glikson (1979), Blockley (1980) and Hickman (in press).
R. DAVY & J. D. LEWIS
374
Figs 1 and 2 —I— 119°45'
21°00'—
21°00'
MARBLE BAR O
21° 15-
- 21°15' Main Granodiorite Suite jHH
Adamellite
t^H
Granodiorite bodies
fHH
Tonalite bodies Quartz diorite
Inferred contact 59306 . , • - X - - Sample points
Figure 1
JU
I 120°15'
I
I
120°00'
119°45'
21°00'-
MARBLE BAR O
-21°15'
21°15'4 |
| Igneous Foliated rocks, including gneiss Highly sheared Micro granite
hypidiomorphic Boundary between ": : — > : ."" hypidiomorphic and xenomorphic xenomorphic textured rocks
15 km Figure 2
Inferred outline of buried dome/ stock of younger granitoid 2 1 o^Q,
Fig. 1.
Outline of the geology of the Mount Edgar Batholith. This diagram is modified from Hickman (in press). Sample sites are superimposed. Discrete bodies of rock referred to in the text and in later figures are identified by letters.
Fig. 2.
Plan showing distribution of various fabrics and petrographic textures in the Mount Edgar Batholith.
MOUNT EDGAR BATHOLITH 375 Bright red-brown sphene is distinctive in some Approximately 350 samples, each of 2-3 kg, samples where it forms large euhedral grains up were collected from the granitoid rocks, mostly to 3 mm across. Elsewhere the sphene is smaller, anhedral and of pale colour. near station tracks. In thin-section the rocks of the granodiorite Chemical analysis was carried out by the Western Australian Government Chemical suite show strained quartz, with some polygonLaboratories and by the Geological Survey. ization, but retain a generally igneous texture. Analyses for Si0 , A1 0 , F e 0 , CaO, T i 0 and Most are not markedly gneissose. A distinctive K 0 were carried out by X-ray fluorescence on feature of almost all samples is the presence of fused discs, using lithium tetraborate as the flux, large, well-shaped to sieved epidote. The univerand on Ba, Ce, Ga, La, Mn, Nb, Rb, Sr, Th, Sn, sality of this mineral, irrespective of whether the W, V, Y, and Zr by X-ray fluorescence on rock has gneissic or igneous texture, and the pressed powder pellets. Determinations of MgO, widespread nature of the chlorite-epidote alteraNa 0, Cu, Pb, Li, Mo, Ni, and Zn were made by tion suggests either a later metamorphic event or atomic absorption analysis after a perchloric- post-crystallization alteration. In a few outcrops rocks are green in hand specihydrofluoric acid digestion and extraction into 5% HC1. F was determined using a specific-ion men but contain no epidote and little plagioclase. electrode, and U by paper chromatography using They contain much sericite and muscovite, some of which appears to have been derived from bioPAN (1 -(2-pyridylazo)-2-naphthol). Full details of analytical and sampling methods tite. These rocks probably formed by local metaand results will be published elsewhere, and a full somatism, in many cases in shear zones. listing of analyses and computer-based statistics Minor Granodiorite-tonalite Plutons has been placed "on-file" at the Geological Small granodiorite and tonalite plutons occur Survey of Western Australia. or close to the margins of the batholith (Fig. A representative selection of about 130 samples on 1). Rocks marked A, B, and C in Figure 1 are was examined in thin section. Rocks were named clearly related chemically to the main granofrcm a visual inspection using the nomenclature diorite but distinctive textures. Pluton A has of Streckeisen (1973), with the additional separa- the overall have composition of granodiorite but intion of adamellite from the granite field. cludes quartz monzonite and adamellite varieties; it is coarse grained with euhedral oligoclase, up to 3 mm across, poikilitically enclosed in microcline. PETROGRAPHY In quartz-poor specimens both euhedral and Over 80% of the Mount Edgar Batholith con- rounded phenocrysts are present. Euhedsists of a large mass of granodiorite, within which ral green quartz several plutons, of composition ranging from prominent.hornblende and red-brown sphene are quartz-diorite to adamellite, have been identified The tonalite of pluton B has a chilled contact (Fig. 1). There are no true granites within the with its host rocks and is a coarse-grained equibatholith and the name "granitoid" has been granular rock composed of euhedral zoned oligoused as a general term for granitic rocks. clase, quartz and biotite with accessory microcline. The Main Granodiorite A quartz-diorite occurs southeast of Marble The main mass of rock has an average compo- Bar, D in Figure 1. This is the most mafic granitoid identified in the batholith. It consists of sition corresponding to granodiorite near the border dividing the granodiorite and tonalite andesine, hornblende and quartz and is slightly to fields. The rocks contain oligoclase, quartz, moderately foliated. South of Limestone Homestead, biotite granomicrocline and biotite, and a few contain green hornblende. Most of the biotite has been chlori- diorite (E in Fig. 1) with hypidiomorphic texture, is flanked by gneissic rocks of similar compositized. The proportion of microcline ranges from about 10% (approximately 20% of the total feld- tion. No contacts have been observed, and its spar) to 'not-visible". A small number of the relationship to the gneiss is unknown. Other previously recognized, and some inferrocks are therefore actually tonalites or trondhjemites, but as the variation appears continuous, red, granodiorite and tonalite plutons are identified in Figure 1. the rock-type is referred to simply as granodiorite. Accessory minerals include, in decreasing order of abundance, opaques, sphene, apatite and allanite. Monazite and zircon are very rare.
Methods of Study
2
2
2
4
2
3
2
3
2
%
• ' E 3
• EE] •B
<68 ]
68 - 7 2 72-75
<0.2 0.2-0.5 0.5-1.0 >1.0
o
> < fcp
r
• •
Marble Bar o
ffl
0
Figure 3c
p.p.m
p.p.m •
<20 20-50 50-100
Marble Bar o
>100
15 km
Figure 3d Li
Ce
Fig. 3a-d.
Patterns o f chemical distribution in the M o u n t Edgar Batholith: 3a, S i 0 2 ; 3b, M g O ; 3c, Ce; 3d, L i .
<50 , 50-100 f H J (50 - 200 at Moolyella)
E33
100 - 2 0 0
ffl
200 - 4 0 0
•
>400
15km
m 3
MOUNT EDGAR BATHOLITH
Adamellites
CHEMISTRY
There are several plutons of adamellite (Fig. 1), characterized by abundant microcline which is equal to, or slightly greater in amount than sodic oligoclase. These adamellite bodies show some polygonization of quartz and often contain minor chlorite and epidote. They are probably older than the Moolyella Adamellite. Adamellite east of Marble Bar (F in Fig. 1) is unusual, having a coarse granophyric texture. It is composed mainly of quartz and perthite. The Moolyella Adamellite (M on Fig. 1) contains a greater proportion of microcline than the other adamellites, and the only mafic mineral present is pale phengitic muscovite. It contains no metamorphic epidote or chlorite. Pluton H (Fig. J) is texturally and mineralogically similar to the Moolyella Adamellite but additionally contains small euhedral garnets (probably spessartine).
Regional Geochemistry
Regional Fabrics and Textures Regional fabrics and textures are indicated on Figure 2. The fabric of the main part of the batholith is not related to the position or shape of minor plutons or to the boundary of the batholith. However, a strong foliation or cataclastic fabric can be related to the sheared marginal phase of the batholith, and to shear zones within the batholith. Large areas of foliated or gneissic fabric and homogeneous igneous fabric can be distinguished but these have no relationship to rock composition. The younger Moolyella Adamellite is conspicuous by its igneous texture and intrusive contacts into the surrounding gneissic rocks. Thin section study has shown that for much of the batholith there is a marginal zone, up to 10km wide, in which the rock has a hypidiomorphic texture with euhedral to subhedral plagioclase and interstitial quartz and microcline. In the central portions of the batholith this texture is replaced by a xenomorphic granular texture, but without a sharp division between the two textural types. Most of the minor plutons which can be outlined on field or chemical evidence are intruded into the marginal hypidiomorphic zone, but those emplaced into the central xenomorphic granodiorite also have a distinct hypidiomorphic texture. This textural zonation of the granodiorites may reflect an original intrusion sequence and cooling process. The marginal phases, intruded late in the sequence and cooling more rapidly, retained an igneous texture, whereas the earlier central portions, cooling more slowly, were subject to late recrystallization and alteration.
377
Regional geochemical maps have been prepared for each analysed component and full details will be given later by Davy & Lewis elsewhere. For this paper representative maps for Si0 2 , MgO, Ce and Li are presented in Figure 3a-d. Contouring in the strict sense is difficult but for many components the rocks may be divided into zones with high, medium or low values. The patterns for any two components are never identical but for some components the patterns can be grouped; MgO, Ti0 2 , V and Sr, and to a lesser extent CaO, all have similar patterns, while Ce, La, Zr and Ba form another group. There is a general indication of a belt-like zone of rocks with distinctive chemistry running from the northwest to the northeast of the batholith, but excluding the northern-most 4 'plutons'' (Fig. 3). Rocks in this belt include adamellite stocks but overall the zone is granodiorite, although more potassic than rocks to the south and southeast. Other areas with distinctive chemistry include the southeast and southwest sheared marginal-zone of the batholith. However, this zone does not continue north past Limestone Homestead, even though rocks with similar petrography are present. Adamellite in the southeast (G on Fig. 1) also has a distinctive chemistry, being characterized by a relatively wide range of compositions different from nearby granodiorites (Table I).
Compositional Features of the Older Granodiorites Typical analyses, mean values, standard deviations and the range in values for rocks of the main granodiorite suite are given in Table I. The mean values of the analyses of all the older granitoids, of the main granodiorite mass, and of the various bodies within the older granitoids are given in Table II. Most plutons have a low scatter of values, as illustrated in Figures 4, 5 and 6. The compositions of plutons B and C (Fig. 1) are virtually identical with that of the main granodiorite mass, even though there is clear field evidence for their intrusive nature. Other plutons (A, J on Fig. 1) tend more to the granodiorite-adamellite boundary, while a further group (F, A, G, I, K on Fig. 1) are adamellites. However, each body shows slight differences from other bodies of similar general composition, particularly in the proportion of trace elements. A plot of principal component values (Davis, 1973) allows discrimination between the
378
R. D A V Y & J. D . L E W I S TABLE I
Geochemical parameters of main granodiorite mass and the Moolyella Adamellite M O O L Y E L L A ADAMELLITE No. of samples = 10
MAIN G R A N O D I O R I T E M A S S No. of samples = 249
Rock type Sample
56306
56386
56986
°mean
s.d.
range
26269
mean
s.d.
S i 02 AI2O3 Fe203* MgO CaO Na20 K20 Ti O2 L0I
66.6 16.1 4.2 1.2 3-7 4. 3 1.5 0.47 0.97
69-2 14.8 3.5 0.96 2.9 4.4 2.4 0.34 1.1
75.3 13.3 1.4 0.06 1.1 3.7 4.4 0.04 0.46
70.5 14.8 3.0 0.81 2.5 4.5 2.4 0.27 0.89
2.59 0.87 0.93 0.43 0.82 0.48 0.94 0.12 0.32
65,2 - 7 6 . 8 11.8 - 1 6 . 7 0.8 - 5 - 3 0.05- 2.0 0.45- 4.2 2.8 - 5.6 0.65- 6.0 0.03- 0 . 5 7 0.20- 2:i
75-0 13.7 1.4 0.15 0.70 3.7 4.9 0.07 0.62
74.4 13.3 1.4 0.13 0.75 3.9 4.6 0.05 0.66
0.59 0.32 0.11 0.01 0.09 0.16 0.28 0.03 0.11
450 30 30 400 30 25 25 325 <1 <5 <5 50 480 <10 4 <1 50 10 105 195
800 70 5 390 30 20 35 390 1 10 5 70 400 10 <2 <1 40 15 65 165
450 60 25 90 20 20 10 100 <1 <5 <5 80 150 30 <2 6 <10 10 25 90
635 67 26 481 42 16 67 327 <1 7 9 99 356 11 3
380 38 26 354 26 9 105 114
140- 3000 < 1 0 - 250 < 5 - 200 10-2850 < 1 0 - 160 < 1 0 - 40 < 5 - 720 3 5 - 775 < 1 - 50 < 5 - 30 < 5 - 30 15- 440 6 0 - 800 <10- 50 <2- 60 < 1 - 10 < 1 0 - 70 < 5 - 100 10- 120 30- 380
240 60 10 650 40 35 55 330 <1 <5 25 420 55 40 8 12 <10 40 75 90
226 64 10 700 36 35 165 332 <1 <5 25 444 60 34 10 12 <10 43 55 95
21 10 6 170 12 10 94 31
%
ppm Ba Ce Cu F La Pb Li Mn Mo Ni Nb Rb Sr Th Sn U V
Y
Zn Zr Total
Fe as
(0
24 14 61 156
5 6 50 132 9 5 2 16 9 20 53
5 34 8 5 4 4 5 8 16
Fe20 3
Means have slight high bias w h e r e range includes < values. Samples w i t h values < have been a s c r i b e d a value of half a limit defined as: Cu(5), Pb(lO), Ni(5), Th(5), Sn (2) , U(l), V (10) , La (10) , Ce(lO), Y (10), Li (5), Mo(l). Representative
individual
samples:
56306 = tonalite; 56386 = g r a n o d i o r i t e ; 56986 = a d a m e l l i t e .
bodies (Fig. 4). In general, granodiorite-tonalite intrusions can be separated from each other by their Ce, La, Cu, and Li contents, and to a lesser degree by Si0 , Fe 0 and MgO. Similarly, the adamellites can be separated from each other by their Ba, F, La, Li, Nb, Rb, Sr, Th and Zr contents, and their Rb/Sr ratios. The map of MgO content (Fig. 3b) indicates three areas of relatively high values. In general the values decline to the northwest. The plot for Ce (Fig. 3c) shows two zones with high values, the western zone being superimposed on three widely differing penological rock types. The group of elements of which Ce is representative (La, Zr, Ba and Ce) is the only one which suggests gradual variation from the centre to the margin of the batholith. Some components, e.g. AI2O3, Na 0 and Y, show little regional variation. The Moolyella Adamellite and Surrounding Areas The Moolyella stock is a fractionated adamellite, distinguishable from the older granitoids by high values for Rb, Th, Sn, U and Rb/Sr (>10), and by a very low K/Rb ratio (< 100). Chemical data are summarized in Table I. 2
2
3
2
The regional geochemical maps for Si0 and Li suggest the presence of a circular feature within the Mount Edgar Batholith, of which the Moolyella Adamellite represents the north-east margin (Figs 3a, 3d). High Si0 values form an annulus near the adamellite, and MgO (Fig. 3b) Ti0 , CaO and V, all display indistinct partial annular features of lower-than-average values. 2
2
2
DISCUSSION The main findings from this study are: (i) The Mount Edgar Batholith is composite and chemically heterogeneous, (ii) A high proportion of the older granitoids are granodiorites and not tonalites as originally thought (Glikson, 1979), (iii) The rocks, like most Archaean granitoids, are high in Na. This applies not only to the older granitoids, but also, to some degree, to the Moolyella Adamellite, (iv) There is a widespread Li halo centred on the known primary cassiterite mineralization which crops out southwest of the Moolyella Adamellite.
379
MOUNT EDGAR BATHOLITH TABLE II
Geochemistry of older granitoid rocks of the Mount Edgar Batholith """" *R©ck
All ~ M a l n 9 ran older . .°"
type
granl t 6 S
Number 324 samples % 70.8 Si O* 14.7 A1203 2.9 Fe203 + MgO 0.77 2.4 CaO Na20 4.5 2.6 K20 0.25 Ti0 2 0.93 LOI ppm Ba Ce Gu t f La f Pb t M Mn Mo Ni t Mb t Rb S-r Th t Sn .+ U f V + Y + Zn if- v
'605 65 25 456 40 16 58 320 • <1 • 7 9 101 330 11 3 2 24 15 58 156
' " " " "
~ ~ S e p a r a t e b o d i e s w i t h i n t h e Bathol ith
—
—-
-
a
B
C
D
E
F
G
H
.
J
K
L
N
0
249
8
6
6
4
2
7
6
3
7
7
2
2
2
2
70.5 14.8 3.0 0.81 2.5 4.5 2.4 0.27 0.89
70.6 14.8 3.1 0.78 2.3 4.4 2.8 0.30 0.95
72.4 15.2 2.1 0.64 3.0 5.0 1.6 0.19 0.70
70.6 14.8 2.9 0.75 2.2 4.8 1.9 0.29 1.2
67.1 14.7 4.5 1.5 3-3 4.3 2.0 0.49 1.9
71-7 14.7 2.3 0.84 2.6 4.4 1.9 0.25 1.97
74.5 12.9 2.3 0.23 0.50 4.5 4.1 0.14 0.92
75.3 13.5 1.3 0 . 12 2.0 4.5 3.5 ND 0.46
74.8 14.5 1.1 0.16 0.51 3-8 4.1 ND 1.1
74.6 13.5 1.6 0.19 0.94 ' 3.6 4.7 0.05 ND
72.9 13.9 2.2 0.37 1.6 4.0 3.5 ND 0.90
76.0 13.7 1.2 0 . 10 1.0 4.4 3.9 ND 0.39
68.5 15-3 3.6 1.0 2.9 4.7 2.0 0.45 1. 1
74.0 13.3 2.0 0.33 1.6 4.3 2.8 0.07 1.6
70.0 13-2 2.5 0.50 2.5 3-6 2.5 0.15 2.8
635 67 26 481 42 16 67 327 <1 7 9 99 356 11 3 (1) 24 14 61 156
463 55 22 490 34 20 44 326 <1 12 8 128 300 <10 <2 2 30 12 68 143
333 18 13 258 13 <20 38 187 <1 <10 <5 48 463 <10 4 <1 17 <10 41 120
417 52 37 363 30 <15 19 286 <1 <5 <10 54 390 <10 5 1 25 10 59 148
600 93 31 480 55 <10 26 471 <1 16 10 68 305 10 2 2 55 28 49 188
400 20 22 265 <20 <10 30 272 <1 7 5 62 400 <10 <2 <1 25 10 57 105
536 127 12 487 69 10 5 154 1 <5 21 106 45 23 <2 2 <10 41 22 288
348 <20 1 57 12 28 14 132 <1 <5 8 91 136 <10 3 1 <10 18 18 102
167 <20 58 160 <10 32 17 307 1 <10 20 183 22 <10 <2 2 <10 32 41 50
511 53 12 210 35 24 27 237 <1 <10 1 i 202 104 33 3 2 <10 24 41 113
985 76 32 196 50 10 14 209 1 <5 <5 84 216 17 4 <1 6 9 44 175
340 <10 10 50 <10 <15 10 205 <1 <5 5 140 65 <10 <2 <2 <10 15 21 30
550 40 22 340 <20 <15 22 337 <1 10 <10 62 470 <10 2 <1 40 12 69 150
400 45 20 100 25 10 25 303 <1 <5 17 125 110 <10 2 1 <10 27 42 112
375 35 20 215 30 10 100 365 <1 <5 10 130 130 <10 2 <1 <15 22 47 157
limit.
Limits are
(GD)
* Rock-bodies i n d i c a t e d by l e t t e r o n F i g u r e s . t Some bias i n v o l v e d . Samples with values below detection Pb(5), N i ( 5 ) , T h ( 5 ) , Sn ( 2 ) , U ( l ) , V ( 1 0 ) , L a ( 1 0 ) , Y (10). + Total iron as F e ^ O a . ND = not d e t e r m i n e d .
\
limit have been ascribed a value of half that
Cu(5),
The Moolyella Adamellite and Adjacent Areas The Composition of Older Granitoids of the The Moolyella Adamellite is distinctive in age, Batholith The average Na 0 content of 4.5%, CaO 2.5% petrography and chemistry. It represents a late and K 0 2.6% compare with world averages for intrusion into the granodiorite. There is a geochemical indication of a circular granodiorites of Na 0 3.75%, CaO 3.83% and feature within which the adamellite outcrops. K 0 2.73% (Le Maitre, 1976). Both the mean This feature has been confirmed from interpreta- value and the range in composition are close to tion of aerial photography and is also seen on those of granodiorites from other Archaean satellite imagery. Creeks crossing the batholith granitoid masses—for example the Johannesshow deflections around the features, the prob- burg-Pretoria dome in South Africa (Anhaeusser, able extent of which is shown on Figure 2. It is 1973). However, the Mount Edgar Batholith apinterpreted as a dome underlain by a stock of pears to be less potassic than granitoids from the younger granitoid, which reaches the surface as significantly younger Yilgarn Block (Archibald et the Moolyella Adamellite. Primary tin-bearing al., 1978; Davy, 1978). Results published by pegmatites of the Moolyella tin-field occur within Glikson (1979) for the older granites of the this dome, but only in the older granodiorites Mount Edgar Batholith are biased for some elewhich form the roof to the stock. The pegmatites ments, for example Li, because of the proximity possibly occur as a kind of stockwork in fractures of the sampling to the Moolyella Adamellite. Harker diagrams (Figs 5, 6) and triangular opened by doming. The large lithium halo (Fig. 3d) associated with plots of AFM, Na 0-Ca0-K 0 and Ba-Rb-Sr the dome may indicate the mineralization and (not shown) indicate that the main granodiorite may reflect either alteration of older granitoids mass and the smaller plutons of older granitoids by Li-rich fluids related to tin mineralization, or are not derived from greatly different magmatic fee related to the magma underlying the dome. In sources. All plots indicate that many of the either case it puts probable limits on the extent of plutons are essentially homogeneous bodies but that their composition is encompassed within the primary tin mineralization. 2
2
2
2
2
2
R. DAVY & J. D. LEWIS
380
Figs 4-7 Figure 5
Figure 4
P, = 0.36 Rb + 0.27 Y + 0.23 Nb - 0.42 Sr - 0.37 V - 0.31 Z n - 0.25 Li - 0.22 F - 0.21 Zr - 0.20 Ni
Figure 7
Figure 6
LTernary eutectic (isobaric minimum) for water-saturated melts at 1000 bar (10^ kPa)
confining pressure (after
Tuttle and Bowen (1958))
Envelope of 9 0 % of granodiorite suite.
%
Si02
Overall envelope of granodiorite suite
Fig. 4.
Plot of principal component values established by principal component analysis (Davis, 1973) for trace elements in discrete bodies within the older granitoids, based on standardized natural data.
Fig. 5.
Plot of Rb against Si0 2 to show distinctive nature of Moolyella Adamellite.
Fig. 6.
Plot of Fe (as Fe 2 0 3 ) against Si0 2 . The regression line suggests a non-minimum melt.
Fig. 7.
Plot of normative Ab-Or-Q to show scatter of rock compositions in relation to (theoretical) minimum melt composition.
MOUNT EDGAR BATHOLITH
envelope of scatter of the main granodiorite mass. The AFM diagrams indicate that the rocks are generally calc-alkaline. The field relationships and petrography of certain plutons intrusive into the batholith, notably pluton B on Figure 1, suggest that these may be post-tectonic relative to the granodiorite but older than the younger Moolyella Adamellite. However, pluton H (Fig. 1) has a different petrographic character from all other posttectonic older granitoid plutons. There is evidence of potash metasomatism along its sheared eastern margin and its composition is substantially more variable than most of the discrete plutons. This rock may be one of the younger granites but this cannot be established without isotopic age determination. Origin of Batholith The batholith is considered to be a composite dome of intruded granitic rock, probably a nested group of smaller intrusions. Hickman (1975, Fig. 3b) postulated a series of cells within the batholith, separated, at least in part, by septa of partly digested relict greenstone belts. The present work differs, in part, from this interpretation. The maps of chemical distributions show that the pattern of variation differs markedly for the several components (Fig. 3a-d). For the most part there is no break of composition at the position of the postulated greenstone keels, suggesting that the lines of greenstone xenoliths have neither acted as a barrier to the intrusion of the granitoids, nor indeed have affected their composition to any significant degree. The origins of these greenstone xenoliths remain obscure. The foliation pattern bears little resemblance to the geochemical pattern. The overwhelming impression is that foliation and intrusions took place together. Intrusion in one part of the batholith could have caused compressional foliation in earlier emplaced rocks. Identified plutons of the older granitoids are almost always marginal; most are outside the foliated granitoid area. Their similar composition to the main mass suggests that they are of the same general age as the main granodiorite. The different textures and clear field evidence of intrusion of some plutons indicates emplacement after most of the compressional activity had ceased. There is no consistent evidence of fractionation. For example, although pluton F (Fig. I) is adameliite, its trace-element signature is not that of a fractionated granite but is similar to the older granites. The adamellite I has a more fractionated character but its textures are not dissimilar from those of the tonalite B, which is also post-tectonic, but unfractionated.
381 The petrographical evidence suggests that the marginal intrusions and other marginal rocks may have cooled first. The remainder of the rocks have recrystallized during slow cooling after intrusion ceased. Thus, the intrusive sequence presented is of a succession of mushroom-like plumes jostling against each other. Inner plumes merged and were recrystallized, concealing the evidence for their origin, while outer plumes cooled relatively quickly, retaining their igneous textures. The size of the post-tectonic plutons may give an indication of the size of the primary plumes. Origins of the Batholith Magma The suggested mode of emplacement described throws little light on the origins of the magma. Plots of normative Q-Or-Ab indicate that many adamellites were likely to have been near-eutectic melts derived from anatexis of older material (Fig. 7). However, the displacement of the plots of the greater part of the granodiorite toward the quartz-albite margin of the diagram indicates that most rocks were not formed from minimum melts but were either intruded at a higher temperature, or as a crystal mush consisting of a minimum or non-minimum melt plus residual restite minerals or xenoliths. An examination has been made of the older granites according to the criteria suggested by Chappell & White (1974) for distinguishing between granitoids of igneous and metasedimentary origin, and by White & Chappell (1977) to determine whether they are the products of ultrametamorphism. On the basis of these criteria all rocks except pluton H (Fig. 1) have an igneous parentage. The latter rock, garnetiferous muscovite adamellite, possibly had a non-igneous parent. However, the garnets are similar to those found in the aplites and pegmatites of the batholith and are probably spessartine of igneous origin. In no rock have undoubted relict sedimentary minerals been identified. Most granitoids of the Mount Edgar Batholith probably crystallized from a non-minimum melt since there is little or no indication of xenolithic material except in the southeast and southwest marginal zones, and in a zone southwest of the outcrop of the Moolyella Adamellite. Though many components on Harker diagrams decrease with increasing Si0 , the apparent point of disappearance is nearer 78% Si02 than the 75% Si0 predicted by White & Chappell (1977) for minimum melts (Fig. 6). The scatter about the trend line for components such as CaO, MgO, total iron, and A1 0 , also suggests either that crystal fractionation from non-minimum melts 2
2
2
3
R. DAVY & J. D. LEWIS
382
has influenced formation of the granitoids, or that there are several superimposed trends. Some authors, for example O'Nions & Pankhurst (1978), and Glikson (1979), believe the source rock for Archaean tonalite magmas is basaltic. However, purely for spatial reasons this seems an unlikely source for a mass the size of the Mount Edgar Batholith. Moreover, Drummond (1979) on the basis of seismic profiles between Shay Gap, Mount Goldsworthy and Mount Newman has proposed a lower crust, below a depth of 13 km, of acid to intermediate granulites. It is easier, therefore, to conceive of magma formation by partial melting of an acid to intermediate lower crust, followed by diapiric uprise. This mechanism is similar to that envisaged by Condie & Hunter (1976) for the emplacement of granitoids in the Barberton region of South Africa. The initial Sr/86Sr ratio of 0.7016, found by de Laeter & Blockley (1972) for the Mount Edgar older granitoids, however, suggests an immediate mantle derivation. One explanation of this conflict is that the formation of the lower crustal granulites and their remelting to give rise to granitoid rocks followed so closely that the batholith retains an essentially primitive-initial S r / S r ratio. 87
87
86
Economic
Potential
The survey has suggested that major tin-fields, other than that known at Moolyella, are unlikely. The Moolyella field is characterized by high background Sn values of about lOppm, high Li (Fig. 3d), and high but patchy F and Ba. There is also an indication of high Si0 . Though some of these elements, especially Sn itself, are high in the exposed portion of the Moolyella Adamellite, it can be shown that the enrichment is related to 2
ACKNOWLEDGMENTS This paper is published with the permission of the Director of the Western Australian Geological Survey. The Western Australian Government Chemical Laboratories carried out most of the analyses and provided computing support for statistical studies.
REFERENCES
R., 1973: The geology and geochemistry of the Archaean granites and gneisses of the Johannesburg-Pretoria dome; in Lister, L. A. (Ed.) Symposium on Granites, Gneisses and Related Rocks, 361-385. Spec. Pubis geol. Soc. South Afr., 3.
ANHAEUSSER, C .
mineralizing fluids rather than being a primary feature of the adamellite since values decrease in the adamellite away from mineralized areas. High F values are localized in the tin-bearing area proper, whereas Li has a halo, with values exceeding lOOppm extending at least 5 km beyond the observed mineralization. The maximum Li value obtained was 720 ppm, against a regional background of 10-30 ppm. In the adamellite pluton I (Fig. 1) the maximum Sn value is 6 ppm and Li is up to 200 ppm. Cassiterite occurrences are known in this rock (Hickman, 1977) but no ore has been mined. Above-background values of Sn have been detected in older granitoids, particularly south of Mount Edgar Homestead. Further exploration of this area is warranted. Isolated samples with values exceeding lOppm Sn were found elsewhere. A broad belt of granodiorite, 2-5 km wide, across the north and northeastern portion of the batholith has Li values in the range 50-150 ppm (Fig. 3d). The high values occur in rocks which contain negligible Sn. Rocks with high Li values do not form a distinct petrographic type and the cause of this anomaly is unknown. Other elements of economic interest were sought but no truly anomalous values were found.
ARCHIBALD, N . J . , BETTENAY, L . F . , BINNS, R . A . , GROVES, D . I., & GUNTHORPE, R . J . , 1978: T h e
evolution of Archaean greenstone terrains, Eastern Goldfields Province, Western Australia. Precamb
Res., 6, 103-131. BLOCKLEY, J . G., 1975: Pilbara Block; in Geology of Western Australia, 81-93. Mem. West. Aust. geol. Surv., 2, 81-93.
, 1980: Tin deposits of Western Australia, with special reference to the associated granite. West. Aust. geol. Surv. Min. Res. Bull., 12.
CHAPPELL, B. W . , & WHITE, A . J . R . , 1974: T w o con-
trasting granite types. Pacific Geol., 8, 173-174.
CONDIE, K. C . , & HUNTER, D. R., 1976: Trace element
geochemistry of Archaean granitic rocks from the Barberton region, South Africa. Earth planet. Sci. Lett., 29, 389-400. DAVIS, J . C . , 1973: Statistics and Data Analysis in Geology. Wiley, New York. DAVY, R . , 1978: A comparative study of the geochemistry of Archaean bedrock in part of the northeast Yilgarn Block. Rep. West. Aust. geol. Surv., 4. DE LAETER, J . R . , & BLOCKLEY, J . G., 1972: Granite ages within the Archaean Pilbara Block, Western Australia. J. geol. Soc. Aust., 19, 3 6 3 - 3 7 0 .
MOUNT EDGAR BATHOLITH DRUMMOND, B. J., 1979: A crustal profile across the
Archaean Pilbara and Northern Yilgarn cratons, northwest Australia. BMR J. Aust. Geo I. Geophys., 4., 171-180. GLIKSON, A. Y., 1979: Early Precambrian tonalitetrondhjemite sialic nuclei. Earth-Sci. Rev., 15, 1-73. HICKMAN, A . H., 1975: Precambrian structural geology of part of the Pilbara region. Ann. Rep. West. Aust. geol. Surv. for 1974, 68-73. . 1977: Explanatory notes on the Precambrian part of the Port Hedland-Bedout Island 1:250 000 Geological Sheets, Western Australia. Rec. geol. Surv. West. Aust., 1977/2. . 1978: Nullagine, Western Australia—1:250 000 Geological Series. Explan. Notes geol. Surv. West. Aust., SF51-5. , in press: Geology of the Pilbara Block and its environs. Bull. geol. Surv. West. Aust., 127. HICKMAN, A. H., & CHIN, R. J., 1976: Explanatory notes on the Precambrian part of the Yarrie
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Geological Sheet, Western Australia. Rec. geol. Surv. West. Aust., 1976/16. HICKMAN, A. H . , & LIPPLE, S. L . , 1978: Marble Bar, Western Australia—1:250000 Geological Series. Explan. Notes geol. Surv. West. Aust., SF50-8. LE MAITRE, R . W . , 1976: The chemical variability of some common igneous rocks. J. Petrol., 17, 1:250000
589-637.
O'NIONS, R. K., & PANKHURST, R. J . , 1978: Early
Archaean rocks and geochemical evolution of the earth's crust. Earth planet. Sci. Lett., 38, 211-236. STRECKEISEN, A. L., 1973: Classification and nomenclature of plutonic rocks, recommendations by the IUGS Subcommission on the systematics of igneous rocks. Neues Jahrb. Min. Mh., 1973, 149164. WHITE, A . J. R . , & CHAPPELL, B. W . , 1977: Ultrametamorphism and granitoid genesis. Tectonophys., 43, 7-22.
i
ARCHAEAN EVOLUTION OF THE HIGH-GRADE GNEISS COMPLEX AT ERRABIDDY, NORTH WEST YILGARN BLOCK, WESTERN AUSTRALIA J. R. Muhling Department of Geology, University of Western Australia, Nedlands, Western Australia 6009 ABSTRACT The gneissic complex at Errabiddy lies in the transition zone between the Archaean West Yilgarn Gneiss Domain, and the Proterozoic Gascoyne Province. Within the transition zone Proterozoic deformation is confined to shear zones, between which Archaean structures and metamorphic assemblages are preserved. The intrusive 2.60 Ga Rocky Bore Granite at the southern margin of the transition zone provides a limit to the minimum age of the gneiss complex. The gneiss complex consists predominantly of adamellite gneisses with abundant pegmatites. The complex encloses lenses of mafic gneiss and supracrustal rocks, and has been intruded by pods of ultramafic and granitic rock. Foliations and lithological layering are parallel and relationships between rock types cannot be determined. All rocks have been recrystallised during an interval of static metamorphism under estimated conditions of 720-800°C and 4-6kb which imply a comparatively high geothermal gradient of 35-40 °G/km. The gneisses at Errabiddy differ from those of many other Archaean high-grade terrains in that adamellites are abundant whereas tonalites are rare.
INTRODUCTION High-grade Archaean gneiss terrains have been the subject of much recent geological research (see Windley, 1976, 1977; Barker, 1979 for reviews). They contain remnants of the oldest known terrestrial rocks and have been studied to enable modelling of the processes responsible for formation of the Earth's crust and its evolution through Archaean time. Study of high-grade gneiss terrains may also elucidate the change in Earth processes which took place in the Proterozoic as many Archaean gneiss complexes are the loci of Proterozoic mobile belts. High-grade Archaean gneiss terrains have been extensively studied in many parts of the world, particularly the North Atlantic Craton (e.g. McGregor, 1973; Sheraton, 1970; Bridgwater and Collerson, 1976). Until recently, little has been done on the high-grade Archaean terrains of Australia since the early work of Prider (1944), Davidson (1965) and Wilson (1958). Exposures of amphibolite- and granulite-facies gneisses are preserved in the northwest of the Yilgarn Block, and have recently been mapped at 1:250000 scale by the Geological Survey of Western Australia (e.g. Elias & Williams, 1977; Williams et al., 1979). This paper presents results from the first detailed investigation of part of the northwest Yilgarn high-grade terrain. The Archaean history
Spec. Pubis geol. Soc. Aust., 7 (1981)
and development of the gneisses are described, and comparisons are made with gneissic complexes from other regions. REGIONAL SETTING The gneiss complex at Errabiddy is correlated with the West Yilgarn Gneiss Domain—an arc of amphibolite- and granulite-facies gneisses which wraps around the western margin of the Yilgarn Block (Gee, 1979). The northern continuation of the gneiss domain formed basement for early Proterozoic sedimentary rocks which, with their basement, were deformed and metamorphosed to form the Gascoyne Province. Between the Yilgarn Block and the Gascoyne Province is a transition zone (Fig. 1), best exposed in the vicinity of Errabiddy, in which there are no Proterozoic metasediments but where the effects of Proterozoic deformation and metamorphism can be discerned within Archaean gneisses. South of the transition zone, gneisses trend approximately north-south and are intruded by Archaean granitoids. The Rocky Bore Granite on the southern margin of the transition zone has been dated by the Rb/Sr whole-rock method at 2603 ± 149 Ma. Williams et al. (1978) believe that this age records the intrusion of the granite, and therefore that it provides a younger limit for the age of the gneiss. They further state that 4 T h e quality of the iso-
386
J. R. M U H L I N G
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Low Grade Iron Fm. [ # ] Felsic Gneiss With Trend |
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Geology of part of the transition zone between the West Yilgarn Gneiss Domain and the Gascoyne Province. Proterozoic shear zones control trends of Archaean gneisses and distribution of Archaean metamorphic facies. The area outlined is shown in more detail in Figure 2.
chron precludes the possibility of major Proterozoic events having affected the rock . . North of the transition zone, trends are dominantly east-west, and Williams et al. (1978) report ages of 1561 ± 100 and 1596 ± 165 Ma for granitoids which intrude gneisses and folded metasediments in this region. Within the transition zone, Proterozoic deformation and metamorphism are confined to approximately east-west-trending shear zones between which Archaean structures and metamorphic assemblages are generally well preserved. Both Archaean and Proterozoic granitoids occur in the transition zone. Four major Proterozoic shear zones or faults, namely the Bullbadja, Dingo, Trinity and Errabiddy shear zones, cut the transition zone into three blocks and control the distribution of Archaean metamorphic facies (Fig. 1). Between the Dingo and Trinity Faults granulite-facies
assemblages are preserved in the gneisses, whereas elsewhere in the transition zone rocks are of amphibolite-facies grade. Outcrop in the granulite-facies block is generally continuous although deeply weathered, but in the amphibolite-facies blocks gneisses crop out only adjacent to the Bullbadja and Errabiddy shear zones. Study of the gneissic complex has therefore concentrated on the granulite-facies block, although reconnaissance work indicates a similar history for the amphibolite-facies area, the major difference being the grade attained during Archaean metamorphism. The amphibolite-facies blocks will not be considered further in this paper, and the following descriptions apply only to the block of granulite-facies gneisses. STRUCTURE The complex consists predominantly of layered quartzofeldspathic gneisses enclosing lenses of mafic and supracrustal rocks. Gneissic foliation and lithological layering are parallel, but lithological contacts are not well enough exposed to determine whether the gneiss and its enclaves have been tectonically interleaved as in other gneissic terrains (Bridgwater et al., 1974; Myers, 1976), or whether the contacts are intrusive or sedimentary. No major fold structures can be discerned in the area due to lack of suitable marker horizons and to disruption by later shear zones, but minor folds are present in both the gneiss and its enclaves. Isoclinal intrafolial folds within banded iron formation have been refolded by minor folds associated with larger scale (0.1 to 0.2 km) mesofolds. The larger-scale folds are depicted by iron formation and quartzite ridges, and show dextral vergence throughout the granulite-facies block. The gneisses, where undisturbed by later Proterozoic deformation, trend between northwest and northeast, and all structures are close to vertical. An early phase of horizontal tectonics cannot, therefore, be proved in this area. ROCK TYPES General Lithology The granulite-facies block is composed mainly (75%) of layered quartzofeldspathic gneisses, with lenses of mafic rock (0.5%), and metasedimentary banded iron formation (1%), quartzite (0.5%) and calc-silicate gneiss (<0.1%). The largest enclave (22%) is the 4 x 8 km Diorite Bore Supracrustal Belt (Fig. 2). Lenses of quartzofeldspathic orthogneiss (0.5%) and ultramafic rock (<0.1 %) intrude the layered complex, but have been metamorphosed with it under granulite-facies conditions. Small (100 x 300m)
GNEISS COMPLEX AT ERRABIDDY, YILGARN BLOCK pods of granitoid (0.5%) were intruded into the gneiss complex after it had been deformed and metamorphosed. Quartzofeldspathic Gneisses The dominant quartzofeldspathic rocks are medium- to coarse-grained, pegmatite-layered gneisses. Foliation is due mostly to alternation of pegmatitic leucoadamellite layers, 10-50 mm thick, with finer grained layers which range in composition from granite to tonalite, but with adamellite most abundant. The foliation in the finer grained layers is enhanced by bands of mafic minerals from 1 to 5 mm thick. The pegmatitic layers make up about 30% of the gneisses and apparently resulted from an early anatectic event. Their coarser grainsize and uniform modal composition suggest that they have formed from partial melting of the compositionally more variable gneisses in which they occur. The foliation has been folded by minor tight to isoclinal asymmetric folds. The short limbs of many folds have been sheared out and subsequently intruded by pegmatitic neosome. Similar shears, without associated folds, give rise to irregularities in the layering which in some places resemble cross-bedding. However, the association of similar structures with folds suggests that these have a tectonic rather than sedimentary origin. Medium- to coarse-grained layer-parallel and cross-cutting pegmatitic veins, as well as irregular masses of leucoadamellite, cut across the gneissic layering and foliation, and are believed to represent a second phase of partial melting. The pegmatitic veins which intrude along minor shears within the gneiss are probably part of this anatectic event. All structures have been recrystallised, and all phases have medium- to coarse-grained, equidimensional fabrics with mineral assemblages consistent with crystallisation under granulite-facies conditions. No relic sedimentary or igneous structures have been found within the quartzofeldspathic gneisses. The layering is due to metamorphic and migmatitic processes, but the lack of any original compositional layering suggests that the gneisses are meta-igneous, probably metamorphosed plutonic rocks. The quartzofeldspathic gneisses are mostly leucocratie with less than 10% mafic minerals. They are generally composed of quartz, perthitic microeline and antiperthitic plagioclase, although in some samples the feldspar is mesoperthite. Magnetite and biotite are the most abundant mafic minerals, although orthopyroxene, hornblende and clinopyroxene occur in some samples. Accessory minerals are apatite and zircon. The
387
felsic gneisses have a medium-grained, equigranular texture with smooth, curved grain boundaries, although most samples show evidence of minor late deformation and recrystallisation. Their texture gives them a characteristic granular appearance which contrasts with the platy fabric of younger unlayered orthogneisses and the flaser or mylonitic fabrics developed by Proterozoic deformation. Mafic Gneisses Mafic lenses from less than one metre to ten metres across, and up to 600 m long, are distributed throughout the gneiss complex at Errabiddy. They are particularly abundant in an area 4 to 5 km northwest of Dingo Bore (Fig. 2) where they make up 30 to 40% of the exposed rocks. Where contacts with layered felsic gneiss are exposed, lithological and metamorphic layering are parallel, but felsic neosome of the second anatectic event commonly surrounds and intrudes mafic lenses and obliterates contact relationships. Except where affected by later alteration, the mafic rocks consist of fine- to medium-grained, equigranular polygonal intergrowths of clinopyroxene, orthopyroxene and plagioclase, with accessory ilmenite and apatite. Weakly aligned hornblende is present in most samples, but quartz is absent. The only clue to the origin of the mafic rocks is the rare occurrence within them of elongate crystals of clinopyroxene and orthopyroxene up to 20mm long and 7.5 mm across. These crystals predate the granulite-facies metamorphism since they are deformed and have recrystallised along kink bands and around their margins to finer, equant grains typical of the groundmass in which they occur. Such large crystals are most likely to have formed in a mafic igneous body, and their presence indicates that at least some of the mafic rocks, particularly those northwest of Dingo Bore, were part of a gabbroic intrusive complex. Metasedimentary Rocks Rocks whose compositions show that they are of metasedimentary origin occur as discontinuous layers within the felsic gneiss. Banded iron formations are most common, particularly in the western two-thirds of the area, whereas quartzites are abundant in the remainder. Minor calcsilicate gneisses occur throughout the area, but metapelitic rocks are virtually absent. Banded iron formations form ridges up to 2 km long and 100m across and, where fresh, are composed of medium-grained equigranular quartz, magnetite, orthopyroxene and clinopyroxene. Quartzites also form prominent ridges. Many of them consist only of strongly elongate and
J. R. MUHLING
388
Trinity B
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Diorite
[
] Quaternary
Alluvium
Felsic Gneiss
Proterozoic
Granitoid
Mafic Gneiss
Archaean
Granitoid
Orthogneiss Fig. 2.
Banded Iron Fm. And Quartzite
v67
Dip And Strike Of Foliation
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Major Fault
Amphibolite
Geology of part of the granulite-facies block between the Trinity and Dingo Faults. No major fold structures are discernible, although mesofolds are depicted by banded iron formations and quartzites.
crystallographically aligned quartz grains, but some contain aluminosilicate, garnet, phyllosilicates and plagioclase. All the aluminosilicate is now kyanite, but it appears pseudomorphous after orthorhombic sillimanite which formed poorly terminated prisms, strongly aligned in the foliation. The kyanite is believed to have formed during Proterozoic metamorphism. Calc-silicate gneisses are mostly mediumgrained, equigranular quartz-diopside rocks with minor plagioclase, microcline, tremolite and orthopyroxene. There are also diopside-scapolite rocks with minor quartz and garnet.
The Diorite Bore Supracrustal Belt Within the centre of the granulite-facies zone is a belt 4 x 8 k m across, of amphibolite-facies
supracrustal rocks. This is the Diorite Bore Supracrustal Belt and is by far the largest of the enclaves within the predominant quartzofeldspathic gneisses. F r o m west to east the sequence across the belt is: banded iron formation (1.5-2 km), quartzofeldspathic gneiss (0.5-1 km), amphibolite (0.7 km), felsic gneiss (0.4 km) and finally m o r e amphibolite (0.7 km). The western margin of the belt is obscured, but on the eastern margin amphibolite is interlayered with quartzofeldspathic rocks typical of the surrounding gneiss complex. Banded iron f o r m a t i o n s of the Diorite Bore belt are similar to those within the adjacent felsic gneiss complex, but they contain amphibole rather than pyroxene. The quartzofeldspathic gneisses are deeply weathered, whereas the am-
GNEISS COMPLEX AT ERRABIDDY, YILGARN BLOCK
phibolites crop out strongly. The amphibolite bands themselves consist of a wide variety of rock types. Hornblende-plagioclase-quartz amphibolite and homblende-plagioclase-garnet-quartz amphibolite are the most abundant rocks, but there are also layers of marble, hornblende rock, gedrite-bearing rock, quartz-garnet rock, pelitic schist and quartzite. The diversity of rock types, including some of undoubted sedimentary composition, indicates that the Diorite Bore rocks formed from a supracrustal sequence, although no relic sedimentary or volcanic structures are preserved within it. The timing of incorporation of the Diorite Bore belt within the gneiss complex presents a problem, since the supracrustals display a lower metamorphic grade than the surrounding gneisses. The sequence of deformation, partial melting and metamorphism in the supracrustal rocks is the same as that in the gneisses, and therefore the supracrustal belt has been part of the gneissic complex since its earliest recorded history. Furthermore, amphibolites of the Diorite Bore belt are interlayered with felsic gneisses containing two-pyroxene-bearing mafic gneiss on the eastern margin of the belt, and the felsic gneiss between the two amphibolite bands also contains some thin two-pyroxene-bearing mafic layers. These data indicate that there was no temperature gradient between the supracrustal belt and the gneiss complex, and therefore the difference in metamorphic grade is attributed to higher water activity in the supracrustal rocks, which stabilised amphiboles over pyroxenes.
Orthogneiss Five km southeast of Trinity Bore, just south of the Trinity Fault (Fig. 2), lenses (0.3 x 1.4 km) of foliated but unlayered quartzofeldspathic gneiss occur within the typical pegmatite-layered, equigranular gneiss of the region. These rocks have a platy foliation due to alignment of lenticular aggregates of quartz and feldspar, and thin stringers of mafic minerals—notably hornblende, biotite and magnetite, but they lack pegmatite layering. Biotite-rich quartzofeldspathic dykes within the orthogneiss have been folded into isoclinal folds, and the foliation of the gneiss is axial planar to these folds. As in the layered felsic gneiss, irregular patches of nebulitic neosome cut both dykes and gneissic foliation. These rocks are therefore interpreted as orthogneisses which intruded the layered gneiss complex after formation of migmatitic layering during the first anatectic event but before subsequent deformation and high-grade metamorphism.
389 Ultramafics Pods of ultramafic rock up to 10 m across intruded both layered gneisses and orthogneiss pods before granulite-facies metamorphism. All ultramafics retain evidence of an original coarsegrained fabric, dominated by orthopyroxene with inclusions of olivine, chromite and magnetite. Coarse-grained olivine (serpentinised), clinopyroxene and plagioclase are present in some specimens. The coarse-grained relics are surrounded by a fine-grained granoblastic groundmass of orthopyroxene, clinopyroxene, pale green or brown hornblende, olivine and green spinel, with minor plagioclase, phlogopite and opaque oxides.
Granitoids The final phase of Archaean activity in the transition zone is represented by small (100 x 300 m) pods of granitoid which intrude the gneiss complex. The granitoid pods are mostly elongate and aligned in the regional foliation, but have themselves no tectonic fabrics parallel to those of the gneisses. Contacts between gneiss and granitoid are sharp. The granitoids are medium-grained hornblende adamellites and granodiorites, although in many samples hornblende has been replaced by intergrowths of biotite and epidote. Textures range from equigranular hypidiomorphic to porphyritic. The granitic intrusions are correlated with the 2603 ± 149 Ma Rocky Bore Granite (Williams et al., 1978), which has an initial Sr^/Sr ratio of 0.7095 and is interpreted as a product of anatexis of crustal material. Preliminary geochronological studies on the granitoids from the granulite-facies block confirm this correlation. The Archaean granitoids have been affected in places by the earliest movements on Proterozoic shear zones, and in these areas have also been recrystallised during Proterozoic metamorphism. In contrast, Proterozoic granitoids within the transition zone post-date the earliest Proterozoic deformation and metamorphism. The Proterozoic granitoids can also be distinguished from Archaean rocks by the presence of muscovite as a minor primary phase. Primary muscovite is absent in Archaean granitoids from this area. METAMORPHISM Introduction Physical conditions during high-grade Archaean metamorphism at Errabiddy are difficult to quantify because of the effects of Proterozoic metamorphism, and the lack of reliable geothermometers and geobarometers among the 86
J. R. MUHLING 390 observed mineral assemblages. A high-tempera- curves with the amphibolite solidi suggests a ture-high-pressure (750-800 °C; lOkb) Protero- minimum pressure of 3-4 kb, but the upper preszoic metamorphism is indicated by a number of sure limit is not well defined. The inferred occurfeatures within the gneiss complex. Where rence of sillimanite indicates that pressures were Archaean gneisses and granitoids have been de- certainly less than 9kb, and the lack of prograde formed by Proterozoic shear zones, coronas of garnets within mafic rocks suggests pressures less garnet are developed between plagioclase and than 6kb. Most of the mafic gneisses are olivine mafic minerals. Dolerite dykes intruded along the normative and similar in composition to the shear zones show no deformation, but they also olivine tholeiites of Green & Ringwood (1967) have well-developed garnet coronas. Garnet coro- and NM5 of Ito & Kennedy (1971). The experinas can form by increasing pressure, decreasing mental work indicates that in rocks of this comtemperature or both, but the presence in the position at 800 °C, garnet starts to form at presDiorite Bore Supracrustal Belt of rocks with the sures of 5-6kb. The lack of garnet in mafic assemblage gedrite-kyanite-quartz (± garnet) granulites at Errabiddy, even the most iron-rich indicates that pressures were of the order of 10 kb (Fe/Fe + Mg = 40), suggests that pressures durat temperatures of 750-800 °C (Green & Vernon, ing Archaean metamorphism were less than this. The P-T grid of Figure 3 therefore allows esti1974). This high-grade Proterozoic recrystallisation is of restricted occurrence within the gneiss mation of the physical conditions prevailing durcomplex, but later retrograde alteration of low- ing Archaean metamorphism. Temperatures were temperature, rehydration character is wide- within the range 720-800 °C at pressures of 4-6 kb. Water activities were low in the quartzospread. feldspathic gneiss and its enclaves, but were probP-T Conditions ably close to 1 in the Diorite Bore supracrustal Fe-Mg distribution between coexisting phases, rocks. and empirical calibrations of the pyroxene solvus have been widely used to estimate temperatures in CONCLUSIONS high-grade metamorphic terrains. Application of The gneiss complex at Errabiddy evolved two-pyroxene thermometry to Errabiddy gneisses through a number of deformation, partialyields temperatures of 815 ± 70 °C, but Bohlen & melting and metamorphic events. It consists preEssene (1979, 1980) have shown these methods to dominantly of layered adamellite gneisses with be both inaccurate and imprecise. Geothermo- abundant pegmatite. Within the felsic gneiss are metry based on coexisting feldspars and coexist- slivers of supracrustal rocks, and mafic gneiss (at ing oxide minerals at Errabiddy is unreliable due least partly metagabbroic), which could represent to the effects of Proterozoic metamorphism, but either tectonic slices or xenoliths within an igconsideration of the relationships between de- neous precursor of the gneisses. The layered hydration and melting curves allows a reasonable gneisses have been intruded by pods of quartzoestimation of metamorphic temperatures. feldspathic orthogneiss and ultramafic rock, and A P-T grid relevant to Archaean metamorph- all have been recrystallised under static condiism of the gneissic complex is shown in Figure 3. tions at temperatures of 720-800 C and pressures It is assumed in this figure that temperature and of 4-6 kb. In the gneisses water activity was low load pressure were the same in the Diorite Bore and pyroxenes formed, whereas in the Diorite Belt and in the gneissic complex. Gedrite, antho- Bore Supracrustal Belt water activity was close to phyllite and hornblende within the Diorite Bore 1 and amphiboles remained stable. The final rocks were stable throughout metamorphic re- Archaean event in the area was the intrusion of crystallisation, therefore the hydrous breakdown pods of hornblende adamellite and granodiorite, curves of these minerals provide a maximum tem- which were probably derived from partial melting perature for the metamorphism, i.e. 770-800 °C. of crustal material. The presence of coarse-grained, cross-cutting The gneiss complex at Errabiddy is similar to quartz-plagioclase and plagioclase-hornblende- those of other high-grade Archaean terrains in quartz veins within amphibolites of the Diorite that it consists predominantly of quartzofeldBore belt indicates that temperatures exceeded spathic gneisses with minor intercalations of the amphibolite solidus at P = PJOTAL- The supracrustal rocks, and mafic and ultramafic solidi for olivine tholeiite under hydrous condi- lenses which could be the remains of layered tions (Yoder & Tilley, 1962) and for quartz- igneous complexes. However the gneisses at bearing amphibolite (Binns, 1969) provide a Errabiddy also show significant differences from lower temperature estimate for the meta- those of other areas. Metamorphic conditions morphism of approximately 720 °C. differ from those of the best-studied high-grade Intersection of the amphibole breakdown terrain, i.e. the North Atlantic Craton. Studies by G
H 0
GNEISS COMPLEX AT ERRABIDDY, YILGARN BLOCK
391
Fig. 3. P-T grid relevant to high-grade gneisses at Errabiddy. The sources of the experimental data are as follows: 1. Aluminosilicates—Holdaway (1971); 2. Aluminosilicates—Richardson et al. (1969); 3. Granite-melt-^ Luth et al. (1964); 4. Qtz Amphibolite solidus—Binns (1969); 5. Amphibolite solidus—Yoder & Tilley (1962); 6. Anth breakdown—Greenwood (1963); 7. Hbd + qtz breakdown—Binns (1969); 8. Opx + plag = cpx -I- grn (ol. tholeiite)—Green & Ringwood (1967); 9. Opx + plag = cpx + grn (NM5)—Ito & Kennedy (1971). Shaded area defines conditions for Archaean metamorphism of the gneiss complex at Errabiddy. Temperatures are limited by the relationships of melting and dehydration curves in the Diorite Bore Supracrustal Belt where a H 0 is close to 1, and are believed to apply to the rest of the gneissic terrain. Maximum pressures are limited by the lack of garnet in mafic pyroxene-bearing gneisses. 2
Wells (1976), Wood (1975) and Dickinson & Watson (1976) on late (2.4-2.8 Ga) Archaean metamorphism indicate conditions of 600-650 °C and 6-7 kb for amphibolite-facies rocks and 750-850 °C and 9-12kb for granulites, with a geothermal gradient of 20-25 °C/km. At Errabiddy, physical conditions of 720-800 °C and 4-6 kb for the granulites indicate a gradient of 35-40 °C/km—substantially higher than that of the North Atlantic region. Another obvious difference between the gneisses at Errabiddy and those of many other terrains is the rarity of tonalites and trondhjemites. Other high-grade terrains are dominated by tonalitic rocks which are believed to have been derived by 15-30% partial melting of mafic rocks, either amphibolites (McGregor, 1979;
Tarney et al., 1979) or mafic granulites (Collerson & Bridgwater, 1979). The abundance of adamellites and paucity of tonalites indicates that at Errabiddy the gneisses were derived either by lower degrees of partial melting which produced minimum melt compositions, or that earlier tonalitic rocks have been extensively reworked and remelted to produce adamellitic magmas. Geochemical and isotopic data are needed to test these hypotheses and place constraints on the origins of the gneisses. Such data are currently being obtained. ACKNOWLEDGMENTS This study was conducted while the author held a Commonwealth Postgraduate Research Award.
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REFERENCES F, (Ed.), 1979: Trondhjemites, Dacites, and Related Rocks. Elsevier, Amsterdam. BINNS, R . A . , 1 9 6 9 : Hydrothermal investigations of the amphibolite-granulite fades boundary. Spec. Pubis geoi Soc. Aust., 2, 3 4 1 - 3 4 4 . BOHLEN, S. R . , & ESSENE, E . J . , 1 9 7 9 : A critical evaluation of two-pyroxene thermometry in Adirondack granulites. Lithos, 12, 335-345. , 1980: Evaluation of coexisting garnet-biotite, garnet-clinopyroxene, and other Mg-Fe exchange thermometers in Adirondack granulites: summary. Bull. geol. Soc. Am., 91, 107-109. BARKER,
BRIDGWATER,
D.,
&
COLLERSON,
K.
D.,
1976:
The
major petrological and geochemical characters of the 3600 m.y. Uivak Gneisses from Labrador. Contrib. Mineral. Petrol., 54, 43-59. BRIDGWATER, D . , M C G R E G O R , V . R . , & M Y E R S , J . S . ,
1974: A horizontal tectonic regime in the Archaean of Greenland and its implications for early crustal thickening. Precamb. Res., 1, 179-197. K. D., & BRIDGWATER, D., 1979: Metamorphic development of early Archaean tonalitic and trondhjemitic gneisses: Saglek area, Labrador; in Barker, F. (Ed.) Trondhjemites, Dacites and Related Rocks, 205-207. Elsevier, Amsterdam.
COLLERSON,
L. R., 1965: Chemistry and paragenesis of some metamorphic orthopyroxenes. Ph.D. Thesis, Univ. Cambridge [unpublished].
DAVIDSON,
Variations in crustal level and geothermal gradient during the evolution of the Lewisian complex of northwest Scotland. Precamb. Res., 3, 3 6 3 - 3 7 4 .
DICKINSON, B . B . , & W A T S O N , J . , 1 9 7 6 :
ELIAS, M., & WILLIAMS,
S. J., 1977: Explanatory notes on the Robinson Range 1:250 000 geological sheet, Western Australia. Rec. geol. Surv. West. Aust., 1977/6.
GEE, R. D., 1979: Structure and tectonic style of the Western Australian Shield. Tectonophys., 58, 327-369. A. E . , 1 9 6 7 : An experimental investigation of the gabbro to eclogite transformation and its petrological applications. Geochim. cosmochim. Acta, 31, 7 6 7 - 8 3 3 .
GREEN, D . H . , & RINGWOOD,
T. H., & VERNON R. H., 1974: Cordierite breakdown under high-pressure, hydrous conditions. Contrib. Mineral. Petrol., 46, 215-226.
GREEN,
G R E E N W O O D , H. J., 1963: The synthesis and stability of
anthophyllite. J. Petrol., 4, 317-351. H O L D A W A Y , M. J . ,
1971: Stability of andalusite and the aluminum silicate phase diagram. Am. J. Sci 271 97-131. C., 1 9 7 1 : An experimental study of the basalt-garnet granulite-eclogite transition; in Heacock, J. G. (Ed.) The Structure and Physical Properties of the Earth's Crust. Am. geophys. Union, geophys. Mon. 14, 3 0 3 - 3 1 4 .
ITO, K . , & K E N N E D Y , G .
LUTH, W . C . , JAHNS, R . H . , & TUTTLE, O . F . ,
1964:
The granite system at pressures of 4 to 10 kilobars. J. geophys. Res., 69, 759-773. M C G R E G O R , V . R., 1973: The early Precambrian gneisses of the Godthaab District, west Greenland. Phil. Trans. R. Soc. Lond., A273, 343-358. _, 1979: Archaean grey gneisses and the origin of the continental crust: evidence from the Godthaab Region, West Greenland; in Barker, F. (Ed.) Trondhjemites, Dacites and Related Rocks, 169-204. Elsevier, Amsterdam, M Y E R S , J . S., 1 9 7 6 : Granitoid sheets, thrusting, and Archaean crustal thickening in West Greenland. Geology, 4, 2 6 5 - 2 6 8 . P R I D E R , R. T., 1944: The geology and petrology of part of the Toodyay District, Western Australia. J. R. Soc. West. Aust., 28, 83-137. RICHARDSON, S . W . , GILBERT, M . C . , & BELL, P .
M.,
1969: Experimental determination of kyaniteandalusite and andalusite-sillimanite equilibria: the aluminum silicate triple point. Am. J. Sci., 267, 259-272. S H E R A T O N , J. W., 1970: The origin of the Lewisian gneisses of northwest Scotland, with particular reference to the Drumbeg area, Sutherland. Earth planet. Sci. Lett., 8, 301-310. TARNEY, J . , WEAVER, B . , & DRURY, S. A . , 1979: G e o -
chemistry of Archaean trondhjemitic and tonalitic gneisses from Scotland and East Greenland; in Barker, F. (Ed.) Trondhjemites, Dacites and Related Rocks, 2 7 5 - 2 9 9 . Elsevier, Amsterdam. WELLS, P. R. A., 1 9 7 6 : Late Archaean metamorphism in the Buksefjorden region, southwest Greenland. Contrib. Mineral. Petrol., 56, 2 2 9 - 2 4 2 . WILLIAMS, S . J . , ELIAS, M . , & DE L A E T E R , J . R . , 1978:
Geochronology and evolution of the eastern Gascoyne Province and the adjacent Yilgarn Block. Ann. Rep. geol. Surv. West. Aust. for 1977, 50-56. WILLIAMS, S . J . , WILLIAMS, I . R . , C H I N , R . J . , MUHLING, P. C . , & H O C K I N G , R . M., 1 9 7 9 : Explanatory notes on the Mount Phillips 1:250 000 geological sheet, Western Australia. Rec. geol. Surv. West. Aust., 1978/13. WILSON, A. F., 1958: Advances in the knowledge of the structure and petrology of the Precambrian rocks of south-western Australia. J. R. Soc. West. Aust., 41, 57-83. W I N D L E Y , B. F. (ED.), 1976: The Early History of the Earth. Wiley, London. W I N D L E Y , B. F., 1977: The Evolving Continents. Wiley, London. WOOD, B. J., 1975: The influence of pressure, temperature and bulk composition on the appearance of garnet in orthogneisses—an example from South Harris, Scotland. Earth planet. Sci. Lett., 26, 299-311.
S. J R , & TILLEY, C . E., 1 9 6 2 : Origin of basalt magmas: an experimental study of natural and synthetic rock systems. J. Petrol3, 3 4 2 - 5 3 2 .
YODER, H .
OXYGEN-ISOTOPE AND OTHER GEOCHEMICAL ATTRIBUTES OF SOME ARCHAEAN GRANITOIDS FROM SOUTHWESTERN AUSTRALIA Allan F. Wilson
Department of Geology and Mineralogy, University of Queensland, St Lucia, Queensland, Australia 4067 ABSTRACT 5 O values (%o) of whole-rock samples from a broad sample of the Archaean Yilgarn Block are:—strongly banded gneisses (10) 7.91 ± 0.52, synkinematic gneisses (70) 7.89 ± 0.87, all groups of gneisses (80) 7.89 ± 0.83, all granites (99) 7.89 ± 0.54. However, distinct values apply to several groups within both the gneisses and the granites, and some of these different groups have slightly different Rb-Sr ages or initial S r / S r ratios. Some of these variations in 5 O may have mineral-search relevance. The oldest gneisses of the non-granulite facies (Dale Bridge and Canning Dam) have the lowest S 0 values, both averaging 6.85%o. Rocks with high Rb/Sr ratios normally are enriched in O . A 330m drill hole into the granite pluton at Doodlakine gave average S 0 values of 8.0%o. Variations and reversals of <5 0 and Rb/Sr trends suggest some geochemical fractionation and subtle layering. A deeper drill hole (1155 m) into the granitic core of the Kambalda Dome gave an average 5 O value of 8.3 %o, but reversals of trends suggest layering as at Doodlakine. Gold-related quartz porphyries in the dome give an average <5 0 value of about 10.4%o. The close similarity of average 6 O values between Australian and Canadian rocks of similar types supports the view that many of the granitoids had protoliths of acidic igneous rocks. Many were probably lavas or tuffs, but some of the larger bodies of gneisses are possibly deformed granitic plutons. Some of the gneisses with high 5 O values greater than 8.5%o probably had protoliths with varying proportions of weathered material. The granulite-facies rocks have very low Rb/Sr ratios, and correspondingly low 5 0 values. In the Dumbleyung region, acidic granulites with values from 3.6 to about 7.4%o occur. Some of those in the Lake Grace and Dangin regions have suffered late K and F metasomatism which has resulted in somewhat higher S O values. Notwithstanding these modifications, the average 5 O value for the Archaean granulites is 6.95%o and is very similar to those from Canada and the Algerian Sahara. ls
86
87
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1 8
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ls
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INTRODUCTION The aim of the study was to document and try to explain the 0 / 0 ratios of the granitoids of the Yilgarn Block of southwestern Australia. The oxygen isotopic composition of the granitoids of the Canadian Shield is now well known (Shieh & Schwarcz, 1974; Longstaffe, 1979) and a few values are available from southern Africa and the Sahara (Fourcade & Javoy, 1 9 7 3 ) . Most of the Australian studies have hitherto been on granulites, both Proterozoic and Archaean (Wilson & Green, 1971; Wilson & Baksi, 1978, and unpublished data). Samples from various geologists have supplemented my own collections. This paper is based on data from about 200 samples that represent a suitable "reconnaissance" sample of a very large portion of the Yilgarn Block. An early study was made of two deep diamond Spec, Pubis geol. Soc. Aust., 7 (1981) 1 8
1 6
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drill holes in granite. These granitic bodies were sampled in this way to see the degree of variation in 5 O that may be expected when random sampling of surface outcrops is carried out. ls
ANALYTICAL PROCEDURES Oxygen was extracted by means of the BrF method (Clayton & Mayeda, 1 9 6 3 ) . The important difference in our procedure is that no conversion of oxygen to CO2 is made. An internal quartz standard has been calibrated, and the results are tied to SMOW through NBS Quartz 28 which we are currently accepting to have a 5-value of 9 . 5 % o (Matsuhisa, 1 9 7 4 ) . The <5 0 values were measured on wellprepared powders, and are quoted as, say, 8.16%o. Most of the samples were measured once; however, where samples were made in duplicate the value is found to be ± 0 . 0 5 % o . The 5
18
394
Fig. 1.
A. F. WILSON
Location of samples from southwestern Australia (Tables I, II and III). A = Armadale, AL = Albany, BC = Beneubbin, BO = Bowgada, BR = Bonnie Rock, C = Canning Dam, CU = Cue, G = Geraldton, KH = Koolanooka Hills, KL = Kalgoorlie, KS = Koolanooka Siding, KU = Kulin, LG - Lake Grace, MM = Mt Magnet, MO = Morawa, MU = Mullewa, MZ = Menzies, NS = Norseman, NT = Northam, P F = Paynes Find, PJ = Perenjori, SC = Southern Cross, T = Toodyay, WK = Wyalkatchem, WM = Williams.
yield of oxygen was monitored to ensure complete reaction. T H E DEEP DRILL HOLES IN GRANITE Two granite bodies were sampled by deep diamond drill holes. One is the large granite body at Doodlakine in the Wheat Belt (UMP No. 1, Fig. 1, Ref. 20) and the other is the intrusive granite dome at Kambalda in the Eastern Coolgardie Goldfields (Western Mining Corporation DDH hole KD6003 and related holes) (Fig. 6; Ref. 29).
The Doodlakine Granitic Body Eight of the 16 samples used by Arriens for RbSr geochronology (Arriens, 1971) were analysed for <5lsO. The <5lsO value of each sample is plotted against its R b / S r ratio in Figure 5, and the samples are arranged f r o m top to bottom of the 300 m drill hole. For the first four samples it is found that <5180 increases directly with increase in R b / S r ratio. This is presumably due to alkali fractionation within the granitoid. However, although Figure 5 shows an initial gradual
395 pillowed basalts (Hanging Wall Basalt) are encountered in DDH KD1029. Their average <5 O value (10 samples) is much enriched (7.5%o), and is consistent with reaction with heated sea-water as Muehlenbachs & Clayton (1976) and others have found in modern oceanic ridges.
GEOCHEMICAL ATTRIBUTES OF GRANITOIDS
decrease of Rb/Sr with depth, there are some reversals. The most obvious is near the bottom of the drill hole where <5 0 has risen sharply to 8.21 %o (compare the top sample of 5 O = 8.24%o). The Rb/Sr ratio has also risen sharply to a level similar to those of the top 100 m of the hole. This chemical and isotopic layering suggests that the Doodlakine batholith is probably structurally layered and may represent a series of tabular bodies (perhaps 300 m thick) of differentiated granitic magma. 18
ls
The Kambalda Dome Granitic Body
ls
ARRIENS'S SAMPLES Arriens (1971) sampled a large number of granites and gneisses (180 and 88, resp.) from the Yilgarn Block. By using R b / S r methods he recognized three distinct episodes of granitoid emplacement and metamorphism, namely, 31002900, 2700-2550 and 2300-2200m.y. Representative samples (113 powders as prepared by Arriens) were analysed for 0 / 0 . 87
86
Diamond drill hole KD6003 begins in metamorphosed basalt (Hanging Wall Basalt) and at 179m it passes abruptly through a sub-horizontal intrusive contact of felsic porphyry and then imperceptibly into coarser granitic rocks. The hole bottomed in these rocks at 1310m. Detailed TABLE I systematic geochemistry has not yet been done on 5 0 values of gneisses from the western Yilgarn Block the drill core, but some isolated data are availarranged in order of approximate Rb-Sr age, increasing from top to bottom. able. Location of samples—Figure 1. The Kambalda pluton is a sodic granite or trondhjemite (chemical analyses in Oversby, Figure 1975). The age of the pluton is given by Rb-Sr as ReferSample and 6 0 value, ence m.y. 2720 ± 370m.y. (Roddick, 1974) and by Pb iso- GNE I SSES ^ 2200 Lake Grace to Kul in topes as 2760 ± 70m.y. (Oversby, 1975). 4164 4165 4162 4163 3021 3022 25-24 8.17 7.32 7-77 7.49 7-53 8.61 The § 0 values show abrupt variations with Meekering Quarry 4149 4151 depth of sample in the bore hole. The uppermost 11 7.14 7.43 sample is highly enriched in O (8.8%o). <5 0 GNEISSES 2600 to 2800 m.^. Koolanooka Hills values of two other samples fall off rapidly to 3 4310 4311 8.02 6.89 about 8.2%o by 206 m and less abruptly to about East of Pinjarra 8.0%o at 309m. Other highly enriched zones 14 4194 4193 4191 4192 7.07 5.47 7-93 7.27 occur at about 550, 975 and 1060m. The average Dumbleyung 4200 4160 4161 4199 6 0 value for the vertical depth of 1133 m (13 24 8.04 7.88 7-91 8.56 samples) is 8.26 ± 0.55%o. However, within that Bi1leranga Hills 2410 2411 4 depth there are three highly enriched zones (up to 10.16 8.05 Corrigin to Brookton 8.8%o) and three zones depleted in O to as low 4174 4171 4172 as 7.3%o. Further sampling especially of mineral 19-17 7-61 8.18 7-60 York phases, is under way to seek an explanation for 12 4018 4016 4017 4019 4012 4013 4015 8.92 10.82 8.61 8.38 9.89 9-90 these variations, for variation in modal mineraBruce Rock logy seems to be insufficient to explain the range 22 4001 4003 4005 7.70 7.73 in 6 0 values. The data to hand suggest that the GNEISSES older 7-95 than 2800 m.y. Morawa to north of Mullewa Kambalda pluton, like that sampled by drill at 2 2218 2219 2456 2450 2451 2453 7-74 8.64 7.01 Doodlakine, is layered geochemically and iso7.06 9-02 7-27 Bowgada to Koolanooka Siding topically. The layering is not evident in hand- 3 3019 2445 2437 2439 8.38 8.23 7-71 7-74 specimen study, for the drill samples appear to be Northam 4524 well preserved, unaltered and fairly homo- 10 4236 4213 4215 8.10 7-76 7.66 8.51 geneous. Dang i n 4011 4010 4008 4009 4006 4007 Numerous porphyry dykes intrude the host 16 7.24 6.25 7.76 7.48 7.18 5-53 rocks of the dome. A porphyry suite of interToodyay 4136 4140 mediate chemical composition averages 9.5%o 10 8.89 7-57 Armada 1e (14 samples). A suite of quartzose porphyries are 13 4090 4092 found to be commonly related to gold mineraliza8.42 8.67 Canning Dam tion at Kambalda and have S 0 values of 13 4033 ^096 6.26 7-40 10.4%o (17 samples) (Wilson & Golding, unpubDale Bridge 4182 lished data). 4175 ^77 15 7.21 5.90 7.45 About 1330m of metamorphosed tholeiitic 1 8
,8
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A. F. WILSON
This was attempted to determine any time relationship between fractionation of oxygen isotopes, and to try to establish the general levels of 0 in the various granitoid bodies in a large portion of the Australian Archaean sialic crust. The S 0 values are set out in Tables I and II. They are arranged generally according to Arriens's groupings on Rb-Sr ages and whether the samples are gneisses or granites. Some of his groups of samples have been rearranged more precisely according to their Rb-Sr age. These are gneisses from Dangin and Koolanooka Hills, and granites from Cadoux-Bonnie Rock, BencubbinWyalkatchem and Perenjori-Paynes Find. An attempt is made to use available chemical data to detect any correlation between 5 0 values and chemical composition. The data most
i
GRANITES •
,
•
18
• , 5
• 6
7
" GNEISSES 2SOO-2BOOI2S) GNEISSES > 2 8 0 0 ( 2 9 )
. 8
9
2200-2500(6)
GRANITES 2 5 0 0 - 2 6 0 0 (4) GRANITES 2600-2800(39)
.
1 8
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11
& 180°/oo
Fig. 2. Range and mean of <5 0 values of gneisses and granites grouped according to Rb/Sr ages (Arriens's collection). 18
commonly available are by-products of the R b / S r geochronology project, namely L Rb and £ Sr (hereafter referred to as Rb and Sr). K and Na values would have been particularly helpful but are generally not available. Rb is normally related to the K-feldspar or mica content of rocks, and unusually high Rb values, especially if linked with low Sr values, TABLE II <5 0 values of granites from the western Yilgarn Block could indicate highly differentiated magmatic or metasomatic rocks. High Sr commonly indicates arranged in order of approximate Rb-Sr age, an abundance of plagioclase or certain mafic increasing from top to bottom. Location of samples—Figure I. minerals. However, many highly differentiated quartz porphyries of the Goldfields regions conFigure Refertain very high Sr and low Rb. Some rocks, especiSample and 6 0 value, Z ence ally those affected by granulite-facies metaGRANITIC INTRUSIONS 2200 to 2500 m.y. Late gran ite porphyries, Koolanooka H ills morphism (such as those at Dangin and Lake 3 2467 2433 7.30 Grace—Table I) are highly depleted in Rb and 7.53 Porphyrit ic granite, Wave Rock, Hyden enriched in Sr. 4170 27 4169 7.60 7.45 Hence in assessing the correlations proposed in Grani tes, Waddour i ng Hill 4221 7 4257 Figures 2, 3 and 4, the limitations of the data and 7.68 7.23 the small sample population must be borne in Mi crogran ites, Northam 10 A216 4249 mind. 7-14 8.22 GRANITIC INTRUSIONS 2500 - 2600 m.y. Moreover, we should remember the basis of Quartz porphyry, Koolariooka Arriens's separation of gneisses from granites. 3 2431 8.41 He says " 'Granite' is used in the broad field Gran i te, Moorine Rock 28 4437 4439 4441 sense which includes rocks ranging in composi7.25 7-39 7.55 tion from granite to granodiorite. [I] . . . label GRANITIC INTRUSIONS 2600 to 2800 m.y. Granite porphyry, Koolanooka Hills obviously banded rocks as gneisses, and other 3 4187 7.58 apparently massive granitic rocks as granite". Grani tes, Cadoux to Bonnie Rock This means that rocks showing flow-banding or 6 4248 4267 4247 7.42 7.01 7.59 syntectonic features could be called gneisses by Grani tes, Cue to Mt Magnet to Paynes Find 1 2461 2464 4189 4314 Arriens, whereas other geologists would classify 7.44 7-33 7.60 8.03 them as granites. Gran i tes, Bencubbin to Wyalkatchem 8 87
86
18
18
18
Q
28-30 20-21 20
5 9 23 15.
4210 4240 4264 4274 7.81 7-36 7.63 8.16 Grani tes, Wiluna to Leonora to Kookyn i e 4315 4323 4329 4332 4337 7.76 8.22 7.57 7.86 7.82 Gran i tes, Doodlakine to Merredin 4081 4085 4075 7.86 7-96 7.05 Grani te, Doodlakine Borehole, UMP 1 4102 4112 4116 4124 4125 4126 8.24 7.94 7-74 7.76 8.04 8.26 Gran i tes, Perenjori to Paynes Find 4293 4295 69-223 7-54 7.48 9.49 Gran i tes, Wongan Hills 4144 4147 8.18 7-77 Gran i tes, Kond i n i n 4166 4167 4168 420l' 7.70 7.65 6.98 7-35 Gran i tes, Williams to-Arthur River 4157 4159 4198 7.44 7.43 7.63
The Gneisses
Ranges of 6 0 values (together with their means) are set out in Figure 2. In Table I, the gneisses are arranged from the bottom in decreasing order of age and (or) initial S r / S r ratios. Where there is doubt about the model age to be used, average values are taken. Inspection of Table I suggests that the oldest gneisses probably have the lowest 0 / 0 values. The marked separation of the <5 0 values of Canning Dam and Armadale, moreover, suggests that Arriens's pooling of their Rb-Sr age data for a joint isochron may not have been 18
87
86
1 8
18
1 6
GEOCHEMICAL ATTRIBUTES OF GRANITOIDS appropriate. These two spatially related localities also display a large separation on a plot of their Rb/Sr ratio versus d™0 (Fig. 3a). The low 5™0 values for Canning Dam suggest that these rocks may be older than those sampled at Armadale. The S 1 8 0 data for the various groups of gneisses (Table I) suggest several distinct lineages of gneisses. This is further emphasized in Figure 3 where the gradients of the lines show differing relationships of Rb/Sr to 1 8 0 content of the gneisses. In some gneisses there is a regular increase in 5 1 8 0 with increase in Rb/Sr, as shown in portions of some differentiated granite bodies (e.g., UMP No. 1 Drill Hole, Doodlakine). This may indicate magmatic protoliths for these gneisses. Alternatively, these trends could be shown by suites of tuffaceous or clastic sedimentary protoliths containing increasing amounts of alkali feldspars.
Fig, 3.
Relationship of S 1 8 0 to Rb/Sr in gneisses. (a) older than 2800m.y. (top) (b) 2600 to 2800m.y. (bottom) A = Armadale, B = Bruce Rock, C = Canning Dam, D = Dangin, DB = Dale Bridge, DU = Dumbleyung, KH = Koolanooka Hills, M-NM = Morawa to North of Mullewa, P = East of Pinjarra, T = Toodyay, Y = York. Refer Table I.
Some of the gneisses, both of the older group (>2800m.y.: Morawa-Mullewa and Toodyay) and of the younger group (2800-2500 m.y.: Billeranga Hills and York) are enriched in l s O . The York gneisses (618G = 8.38 to 10.82%o) are particularly interesting in that they show an inverse relationship between Rb/Sr and <5180 (Fig. 3b). This could indicate essentially sedimentary protoliths with varying proportions of fairly unaltered
397
rock fragments and quartz grains. Longstaffe & Schwarcz (1977) show that S 1 8 0 values of most of their Canadian Archaean clastic sediments range from 8.0 to 13.3%o. They also point out, however, that at high metamorphic grades (such as apply to the York samples) Archaean metasedimentary gneisses may have <5180 values that are indistinguishable from those of Archaean orthogneisses. A small group of much younger gneisses (between 2000 and 2200m.y.) are recognized by Arriens. Those from the Meckering quarry have initial 8 7 Sr/ 8 6 Sr ratios and 5 1 8 0 values similar to those of the granites of similar age from Waddouring Hill. The gneisses at Meckering may represent a deformed post-Archaean granitic stock. The gneisses from the Lake Grace—Kulin region, however, are completely different. They are charnockitic gneisses and are remarkably like those of the type charnockite localities near Madras, India. They are highly depleted in Rb (average of seven = 57ppm) and the Rb/Sr ratios average only about 0.17. The only comparable rocks are some of the ancient granulites from Dangin. A discussion of the granulite 5 1 8 0 values appears below. The Granites The 6 1 8 0 values from samples from a 350 m drill hole at Doodlakine in a large Archaean batholith have already been considered. The range of 5 1 8 0 values from 8.26 to 7.74%o in these samples (Fig. 5) from an apparently homogeneous granite is disturbing. For this reason the S 1 8 0 value of a single specimen collected from a granitic pluton or a large outcrop of a granitoid crust can scarcely be expected to give its average 18 0 / 1 6 0 composition to an accuracy better than ±0.3%o, especially as 5 l s O values for the minerals are not available. Table II and Figure 4 show that the 5 l s O values of many granites of comparable age fall within the range of the Doodlakine suite of samples which are adamellites with Rb/Sr ratios between 1 and 2. Many other granites, however, have <5180 values that cluster between about 7.0 and 7.8%o. These generally have Rb/Sr ratios considerably less than 1. Notable exceptions are among the youngest granites of the region, such as those at Moorine Rock (2600m.y.: Rb/Sr - 3 . 4 ) , Waddouring Hill (2300m.y.; Rb/Sr - 1.3) and Hyden ( - 2 3 0 0 m . y . ; Rb/Sr - 1 . 6 ) whose 5 l s O values average 7.4, 7.5 and 7.5%o respectively. There is, then, no obvious correlation between age of emplacement and 5 1 8 0 value of these granites. Indeed, their 5 l s O values are mono-
A. F. WILSON
398
TABLE I I I
Average <5 0 values of granites and gneisses from the western Yi/garn Block arranged according to age. Samples collected by Arriens (1971) 18
Youngest
gneisses
Youngest
granites
G r a n i tes G r a n i tes Oldest
granites
G n e i sses Oldest All
gneisses
No. of Samples
Age, m.y.
8 8
^ 21 00 ^ 2300
7.68 ± 0.49 7.52 ± 0.3A
8 13 22
2560-2630 2670-2700 2700-3000
7-52 ± 0.41 1.1k ± 0.29 7-80 ± 0.52
23 31
2500-2800 > 2800
6
A1 1 gnei sses
0 ± SD
{%0)
8.15 ± 1.11 7.68 ± 0.96
51 62
granites
1 8
7-70 ± M 3 7.86 ± -0.99 TOP
i
BOTTOM 1 -8
-
Rb/Sr 1 - 6
/
\
r 1
1 • 4
-
1 - 2
1 - 0
A
4
Fig. 4. Relationship of 5 0 to Rb/Sr in granites. (a) 2600 to 2800m.y. (top) (b) younger than 2600 m.y. (bottom) B = Bencubbin to Wyalkatchem, C = Cadoux to Bonnie Rk, K = Kondinin, M = Moorine Rk, N = Northam, WD = Waddouring Hill, WH = Wongan Hills, WM = Williams. Refer Table II. 18
7
8
< )
1
8
l 9 0 ° / o o
Fig. 5. Relationship of <5 0 to Rb/Sr in samples from drill hole in Doodlakine Granite. ,8
GEOCHEMICAL ATTRIBUTES OF GRANITOIDS
tonously similar, and probably reflect those of the gneisses from which many have been derived. Average S 0 values of the rocks sampled by Arriens are set out in Table III, and are arranged according to their Rb/Sr age. It is clear that there is no significant difference in 5 O between groups of rocks of different ages nor between granites and gneisses. SAMPLES FROM THE EASTERN GOLDFIELDS PROVINCE Five samples from the Wiluna-LeonoraKookynie region are included in Arriens's collection. (Table II). Goldfields material from three other sources has been studied. The largest is a set of granitoids for which many chemical data are available, and covers a wide area of the Eastern Goldfields (Gee, 1979). Archibald & Bettenay (1977) have made detailed geochemical studies of granite-gneiss relationships in the Goldfields, and portions of the original powdered samples of many of these rocks were made available by Dr Bettenay. Another set of samples has come from Dr Gee, and resulted from his field mapping of the Southern Cross 1:250 000 sheet of the Goldfields (Gee, 1979). His classification of the granitoids is largely based on field and petrographic observations, and it is not always possible to relate a granitoid type in Bettenay's collection to those supplied by Gee. A third set of samples was collected by myself, and includes drill core from the Kambalda Dome (by courtesy of Western Mining Corporation), and samples of gneisses and granites from the Norseman region. Each set of samples is first treated separately in this paper. DR BETTENAY'S SAMPLES Figure 6 gives the location of the granitoid samples and the generalized geology of the Eastern Goldfields Province. Details of the geology are discussed by Archibald et al. (1978). The greenstones have not been included in this present study. The samples have been classified on majorelement chemistry, textural and field characteristics. Their 6 G values are set out in Table IV. The 6 0 values are arranged in presumed geological order in Figure 7. It should be noted, however, that the radiometric age control of these samples is not as complete as for the samples of Arriens. The Gneisses These rigidly banded granitoids are represented by five analyses from Bettenay's collection.
399
I8
ls
18
,8
50 Fig. 6. Map of Eastern Goldfields Province showing location of samples in Tables IV and V. Legend: black = greenstone belts; random dashes = migmatite complexes; single spots = synkinematic granitoids; double spots = porphyritic biotite adamellites; arrows = hornblende granitoids; blank spaces = postkinematic complexes and some gneisses. For details, see Archibald et al. (1978), and for the Southern Cross Sheet (rectangle in SW.) see Gee (1979).
Their Rb/Sr ratios are very low (average 0.13) and their range of 5 O values is from 7.61 to 8.32, with a mean of 7.88%o. Gneisses other than those from this collection were analysed, and discussed elsewhere in this paper. The Synkinematic Granitoids These are gneissose, dome-like diapiric bodies that normally intrude the greenstone belts concordantly. Six samples from this collection and one from Gee's collection (50863, Olga Rock) have been studied. Their S 0 values range from 7.45 (50863) to 8.14 (77686). An interesting feature of these rocks is that some have Rb/Sr ratios far higher than those of the gneisses (-0.12)—for example, 2.5 and 3.8 for Ref. 9 and 42, resp. This may indicate acquisition of metamorphic differentiates (rich in Rb) that could have facilitated diapiric uprising. ls
18
Post-Kinematic Granitoids Figure 6 shows that these comprise the bulk of the exposed granitoids of the region. Several granite types become progressively enriched in
A. F. WILSON
400 TABLE
IV
5lsO values of granitoids from the Eastern Goldfields Province of the Yilgarn Block arranged in approximate order of age, increasing from top to bottom. Location of samples—Figure 6. Fig. Ref.
Sample No.
Rock Type
6180 U J
selected to see if there are significant differences in 1 8 0 / 1 6 0 ratios between the various groups of granitoids of the region. Their 5 l s O values are set out in Table V in presumed order of geological age (pers. c o m m . Dr Gee, 2 5 / 2 / 8 0 ) , and their classification is that used by Gee (1979).
POST--KINEMATIC GRANITOIDS POSTKINEMATIC
GRANITOIDS
13 .4 1
Fract ionated Leuco-adamel1i tes 821 38 Mungari Ularring 81879 Kellys Rock 77271
8.65 8.42 7.80
27 -8
Biotite Adamel1ites 82116 Builabul1i ng 81929 Mt Correll
7.43 8.64
Equigranular
25 25 26 26
Equigranular Leuco-adamel1ites & Granodiorites 82090 Boorabin, leucocratic Boorabin, monzonitic enclave 82091 Wool gangie Res., leucogranodiorite 82103 82102 Wool gangie Res., tonalite enclave
7.65 7-37 7.60 7.39
Porph. Biot. A d a m e l l i t e (5)
6 10 21
Coarse Leuco-•adamel1i tes 82160 Barklay Rock East of Peninsula 82009 Mt Clara 82079
8.51 7.61 7.63
35 43 3 7 39
Porphyritic Etiotite Adamel1ites £ Granodiorites 82059 82041 Wheeler Rock 81864 Bui lock Holes 81926 Nierguine West 81955
7.86 7.80 7.86 7.35 7.72
5 40
Hornblende Granitoids 82162 20 km S of Menzies Lake Johnston, tonalite 82199
8.36 7.23
Fractionated Leucoadamellite(3)
«-
B i o t . A d a m e l l i t e (2)
Coarse
BANDED
<
•
•
1 I—®—I
GRANITOIDS(6)
G N E I S S E S (5)
1
#-H i—
•
i• 7
, 6
Fig. 7.
#
i
L e u c o a d a m e l l i t e (3)
Hornblende G r a n i t o i d i 2 ) SYNKINEMATIC
—
)
Leucoadamellite(4)t
1 8
9
1 8 0 ° / o o
Range and mean of <5180 values of granitoids from Goldfields Region (Bettenay collection—Fig. 6 and Table IV).
SYN-KINEMATIC GRANITOIDS 12 36 20 42 9 2 "
82130 77686 81988 82032 81994 77258
The Banded Gneisses
ites Coolgardie pluton. Widgiemooltha W of Yellowdine Lake Johnston Peninsula Pigeon Rock
7.90 8.14 7.9A 7.8I 8.03 7.87
Lake Johnston West Lake Johnston West Koolyanobbing East Depot Rocks, enclave in gn. gran. Split Rock West
7.5] 7.6I 7.00 8.46 8.32
BANDED GNEISSES 4] 41 11 28 33
82023 82025 81884 82129 81974
Two samples of the presumed most-ancient granitoids show a range of < 5 1 8 0 values from 7 . 6 8 to 8 . 8 3 % o . Some of the rocks placed in this group show relics of metasediments, and the high 5 1 8 0 value of 8 . 8 3 % o probably reflects a metasedimentary component.
S 1 8 0 as the R b / S r ratio increases. The earliest granitoids have 5**0 ~ 7 . 5 % o and R b / S r ratio — 0.4, whereas the most fractionated leucoadamellites have 8™0 ~8.6%o and R b / S r ratio — 5.1. The data in Table IV suggest that the biotite adamellite f r o m Mt Correll (Ref. 8) should be included in the " f r a c t i o n a t e d leucoadamellites , , for it has high R b / S r (3.0) and very high d ^ O = 8.64%o. The Menzies-type hornblende-bearing granitoids are represented by only one analysis (5). It has an anomalously high < 5 l s O ( 8 . 3 6 % o ) for its exceedingly low R b / S r ratio (0.10). Analyses of Fig. 8. more samples of this granitoid are required. DR G E E ' S S A M P L E S Figure 8 gives the location of the granitoid samples and the main geological features of the Southern Cross 1:250000 sheet. The relationship of the sheet to the general Goldfields region is shown in Figures 1 and 6. The 17 samples were
Map of the Southern Cross region (for regional setting, see rectangle inset in Fig. 6) showing the location of samples included in Table V. Geology simplified from 1:250000 map (Gee, 1979). Legend: black = greenstone belts + metasediments + felsic metavolcanics; single dots = gneisses and synkinematic granitoids; double dots = porphyritic adamellite (post-kinematic); biotite adamellite (post-kinematic).
GEOCHEMIGAL
ATTRIBUTES OF
401
GRANITOIDS
TABLE V
W
values of granitoids from the Southern Cross 1:250,000 Sheet arranged in approximate order of age, increasing from top to bottom. Location of sample>—Figures 6 and 8.
Figure Reference
6
Samp 1e No.
0
Rock Type
POST-KINEMATIC G RAN I TO I OS Even-grained Bioti te Adame11 i t es (Agv)
If
FP 568 50877
32
5 km ENE of Keokanie Rk Roe Dam, near Gnamma Hill
8.17 8.15
3 km SSE of Bacon Hill Caroling Rock
7.83 7.91 7.19
Porphyri tic Biotite Adamellites (Agl) 17 22 24
FP 574 FP 172 50864
2 km E of No. 7 Pumping Station (age relationshi ps uncertain) Agmatitie Gneisses (Am)
SVN-KINEMATIC GRANITOIDS
18 16
7-55 7.90 7.55 7.62 8.01 7-38 7.59 7.35
Biotite adamellite, Biljahuie Rk
7.42
50863 50875
Biotite adamel1ite gneiss", Olga Rk, W side of Parker Dome Tonalitic gneissic granite, 5 km W of Muntadgin
7.45 7.30
50824 50865
Adamellitic migmatitic gneiss, SE of Parker Dome Biot-garn adamellitic gneiss, Split Rk
7.68 8.83
50874
ik
23
Porph. biot. adamellitic gneiss, Nurdungurra Rk leucosome, Moorine Rk Quarry . (leucocratic pa 1aeosome , , r (melanocratic leucosome, 12 km S of Kerman Rk Agmati te most mafic phase leucosome, Police Rk Tank Agmat i te palaeosome
50868 50870 50869 B 50869 A 50872 A J 50871
19
50873
J
Foliated Granites (Agg) FP 167
Agmat i te
Gneissic Granites (Ang) 31 30 BANDED GNEISSES (AnI)
34 35
The Synkinematic Granitoids Dr Gee makes three groupings of these rocks whereas the Bettenay collection contains only samples from the gneissic granite domes. The Foliated Granites and Gneissic Granites These two groups are discussed together for their S 18 0 values are very similar, i.e., ~ 7.4%o. One of the samples of gneissic granite comes from Olga Rock in the western side of the Parker Dome (Ref. 31). We should note, however, that the gneissic granites from the domal structures sampled in Bettenay's collection form a tight cluster of d 1 8 0 values about 7.95%o, and thus on average are enriched in l s O by about 0.5%ocompared with Olga Rock. The Kambalda Dome (29) which is not fully unroofed is even more enriched (8.3%o). This suggests that within the gneissic granite domes there could be somewhat higher 6 values, especially near the domal apices where silica/K-feldspar metamorphic differentiates could be anticipated. For example, some of Bettenay's domal rocks have unusually high Rb/Sr ratios, e.g., 3.8 for 82032 (Ref. 42). However, in contrast, a sample of gneissic granite from Olga Rock in Parker Dome has a much lower Rb/Sr ratio of 0.23 and could explain the lower 6 l s O value of 7.45%o from that locality.
The Agmatitie Gneisses Although these make up the largest group of the collection, their 5 l s O values are not as varied as one might expect. The leucosomes of the agmatites show only a small enrichment in S 1 8 0 over that found in the palaeosome, indicating that the latter have probably attained isotopic equilibrium with the more siliceous host rock, at least on the scale of centimetres. The Post-Kinematic Granitoids Vast volumes of granitoids (largely biotite adamellites) were intruded into gneissic and greenstone terrain. In Gee's collection two major types are sampled. The porphyritic adamellites fall into the <5180 range of many of the granitoids, whereas the much larger expanses of even-grained biotite adamellite have virtually identical 6 l s O values. However, only two samples, collected 110km apart, have been measured. Their high S 1 8 0 values ( ~ 8.2%o) probably reflect their fairly high silica content. OTHER SAMPLES These were collected by myself from various parts of the Yilgarn Block. The locations of those from the Goldfields are shown in Figure 6, and
A. F. WILSON
402
the <5 0 data appear in Table VI. Those from the Granitoids from the Norseman Region Wheat Belt region are shown in Figure 1, and the Banded Gneisses. The three representative <5 0 data appear in Table VII. samples (Ref. 37, 38) from the eastern flank of TABLE VI the Pioneer Dome north of Norseman have S 0 values of various granitoids from the virtually identical <5 0 values of 7.9%o. This Norseman region. implies a remarkable degree of oxygen-isotope Location of samples—Figure 6. homogenization of rocks which are rigidly Fig. Sample 6 0 banded and which are 9 km apart along tectonic Ref. No. Rock Type (U strike. A similar gneiss from the same locality POST-KINEMATIC GRANITOIDS (Bettenay collection, 77684) has a low Rb/Sr Fractionated Granitoids ratio of 0.13. Thus the fairly high <5 0 value of 37 AFW83 Fine even-grained granite, Fifty-Mile Rks 8.80 7.9%o in these banded gneisses may be due to 45 AFW111 Even-grained granite, 4 km S of Dundas Rks 8.74 high quartz content of the original ?felsic vol49 AFW 124 Fluorine-rich granite, Bedonia Rk 8.14 canic protolith. Coarse Even-grained Granites 18
18
18
18
8
18
44
AFW108
Dundas Rks
8.17
48
AFW123
Bulandia Rk Hole
7.72
Synkinematic Granitoids.
Porphyritic Granitoids 46
AFW114
1.5 km SE of McPherson Rk
7.14
AFW113
McPherson Rk
<5 0 values of granulites from the Wheat Belt region. Location of samples—Figure 1. 18
7.52
Granitic Gneisses 50
AFW100
1.5 km N of Salmon Gums
18
TABLE V I I
Gneissic Granites 47
The <5 0 values of
the three samples fall within the range of the other synkinematic granitoids already studied.
SYN-KINEMATIC GRANITOIDS
7.36
Figure Reference
37
AFW68
Fifty-mile Rks, summit
7.88
37
AFW106
Ditto, E side
7.92
2k
38
AFW97
Twenty-five-mile Rks
7.86
BANDED GNEISSES
Kambalda Dome
19
A detailed stable isotope study of many rocks from underground workings and drill core from Western Mining Corporation has been completed (Wilson & Golding, unpublished data). Some of the averaged granitoid data are referred to above and are incorporated in Figure 9. The trondhjemitic core of the dome has an average 5 O value of8.3%o (13 samples, 1133 m).
26
2k 18
ls
Sample S i1i ceous granuli tes Dumbleyung region 18898 ^ 39690 39691 Corrigin region 39689
7.4
Charnockitic gneisses Dingo Rock 37887 15622'
7,3 7.3
Cordierite gneisses Dumb]eyung region 18983 S. Quairading 41-322 41318 41331 41313
6.0 6.4 7.2 7.2
6.8 7.4 3.6
5.0
W.A. A R C H A E A N G R A N I T E S ( 9 9 ) • W . A . A R C H A E A N G N E I S S E S (BO) A R C H A E A N GRANITIC ROCKS (several l O O ' s ) A ARCHAEAN SEDIMENTS(several lOO's) STRANGWAYS — W A . \
— ^
R G . C E N T . A U S T . F E L S I C G R A N U L I T E S (21)
A R C H A E A N F E L S I C G R A N U L I T E S (25) GRENVILLE
HIGH G R A D E PARAGNEISS ( 4 4 )
F O O T P R I N T G N E I S S (11) IN O U Z Z A L
— l
GRANULITES(16)
C E D A R L . - C L A Y L . - T W I L I G H T G N E I S S (17) •
1
PUCKWASH
*
—1
10
11
12
13
P A R A G N E I S S (12)
F E L S I C P Y R O C L A S T I C S (21) •* C L A S T I C M E T A S E D S . ( 4 4 ) 14
6 O °/oo 18/-
Fig. 9. A
u
S
groups
d
w
i
r
S (
average j ^ O values of Archaean granites, gneisses and granulites from
"
d s e w
C d
ere
t n a n g l C S
-
F o r e i g n d a t a m a i n l
i n d i G a t e
r a n g e s
o f
y after Longstaffe & Schwarcz (1977)
a v e r a g e s
'
D o t s
are averages of specific
GEOCHEMICAL ATTRIBUTES OF GRANITOIDS 403 Post-kinematic Granitoids. These are all en- sediments, acid pyroclastics and tectonically riched in O , especially those where magmatic modified igneous intrusives. The granites, on the differentiation has been involved. The even- other hand, appear to have largely developed by grained granite south of Dundas (Ref. 45—Table reconstitution of the protoliths of the gneisses or VI) is highly enriched in 0 and is a freshly of the gneisses themselves. For this reason the blasted sample from the spoil dump of a cassiter- average 5 O of the granites is virtually identical ite prospect. to that of the gneisses ( 7 . 8 % o ) , but the spread of values is much less broad ( 7 . 0 to 9 . 5 % o ) . The Wheat Belt The average <5 O of the Australian Archaean Several acid and basic pyroxene granulites and granites and gneisses falls within the range of some granites and gneisses from the Wheat Belt averages compiled from several hundred other region are included in another ms. (Wilson & Archaean granitic rocks. Baksi, unpublished data). A summary of 6 0 values of the acidic rocks is included in this paper ECONOMIC SIGNIFICANCE in Figure 9, and Table VII. Although the average 5 O values of the The granites and gneissic granites generally granites is about 7 . 8 % o , there are significant have S 0 values similar to those already noted in differences between granite plutons. Moreover, this paper. Exceptions apply to the gneissic differences can be recognized within plutons granites and siliceous pyroxene granulites and themselves. An example of the possible economic charnockites which are somewhat depleted in relevance of this comes from the Dundas region 0. Removal of the products of partial melting south of Norseman. There the coarse "Dundas during granulite-facies metamorphism (probably Granite'' has a characteristic 6 0 value of recognizable by unusually low Rb/Sr ratios) 8 . 1 7 % o (Ref. 4 4 , Table VI). However, a somecould be a major cause of the depletion in the what finer even-grained granite phase occurs acidic rocks. about 4 km to the south associated with a tin prospect. Recognition of the 5 O signature of this granite could be useful in following extenCOMPARISONS sions of the tin phase (Ref. 45, Table VI). Not In Figure 9 several groups of Archaean rocks only may fresh rocks be used, but the source of are compared with the Australian Archaean weathered quartz grains may also be followed by granitoids. The lowest metamorphic grades are this technique. shown nearest the base of the figure and metaSome unusual differentiates of granites have morphic rocks of increasing grades are above been recognized during this study, and explorathem. It is clear that as grade increases, 5 0 tion for several commodities in these and related decreases, as Longstaffe & Schwarcz (1977) have granite phases commends itself. already pointed out. The Western Australian The close correlation between <5 0 of goldArchaean granulites (Table VII; Table I (16) and bearing quartz and certain veins and hydro(25-24)) have an average <5 0 of about 7.0%o; thermally altered country rocks has now been this is almost identical to the averages of the used in several Australian gold prospects Footprint Gneiss (granulites from Ontario; Long- (Golding& Wilson, 1 9 8 0 ) . staffe, 1979) and the In Ouzzal granulites from Not only is it possible to recognize a goldthe Sahara (Fourcade & Javoy, 1973). However, bearing quartz vein, but in favourable situations the intensely metamorphosed Proterozoic felsic one can even estimate the depth in a mine where a granulites from the Strangways Range in central sample was collected. This is because variations Australia are very much more depleted in 0 in 5 0 in a suite of mineral samples may be due with an average 6 0 of about 4.6%o. This to change of temperature with depth as well as to extreme O depletion is thought to be due to a intrinsic differences between veins or wall-rock combination of two effects: (a) pre-granulite alteration. Active research is in progress to lowering of O by reaction with heated waters, evaluate the economic potential of these techand (b) removal of much of the products of par- niques. tial melting which would be enriched in O CONCLUSIONS (Wilson & Baksi, unpublished data). Reference has already been made in this paper The normal (that is, non-granulite) Archaean gneisses from Western Australia are plotted near to some of the mechanisms for the production of a wide range of 5 0 values in Archaean granithe top of Figure 9. It is important to note that the spread of the 6 0 values for the gneisses is toids. We have already seen that once granitic bodies great (5.5 to 10.8%o). It appears that the gneisses represent a wide variety of Archaean clastic have formed by reconstitution of the gneisses and ls
1 8
ls
ls
18
ls
18
18
18
ls
18
18
18
1 8
18
18
ls
ls
ls
18
18
404
A. F. WILSON
granitoids shows that water rather than C 0 was the dominant fluid. However, in the granulite terrains such as those near Dumbleyung, DanginS. Quairading and Lake Grace (Fig. 1, Ref. 24; 16 and 18; 25-26, resp.) Rb is highly depleted, and many rocks are composed of anhydrous minerals, except where subsequent metasomatism has introduced some fluor-rich phlogopite. The fluid with which the anhydrous granulites reacted was probably C02-rich, as has been found in other granulite terrains (Touret, 1971). Deepseated (?) mantle-derived fluids may have been effective in flushing out the water from the protoliths of the granulites, thus enabling them to take on remarkably similar 0 / 0 ratios that are commonly 1.5 to 2%o lower than those of normal Archaean granites. The upward spread of <5 0 values in the granulites is almost always found to be due to the activity of fluids (especially water).
sediments, many of them underwent magmatic differentiation. In general it appears that an increase in <5 0 correlates with an increase in Rb/Sr ratio, as was pointed out in Figures 3, 4 and 5. The observation of this feature in gneisses as well as granites raises the question of whether the Rb/Sr ratios are inherited from their protoliths, many of which are presumed to be composed largely of fairly fresh acid igneous detritus. However, the degree of chemical alteration from weathering will also raise the <5 0 value of the protoliths of the granitoids. The. effect has been well documented in Canada where Shieh & Schwarcz (1978) show that the "granitic rocks from the Superior, Slave and Churchill Provinces vary only slightly from region to region (<5 0 = 6.9 to 8.4%o) and are significantly lower in O than similar rock types from the younger Grenville Province(<5 0= 9.2 to 10.0%o)". They estimate an average of <5 0 of 8.0%o for the surface crystalline rocks of the Canadian Shield which represents an enrichment with respect to probable mantle-derived starting materials of about 2%o. The average of the Australian Archaean granitoids described in this paper is almost the same (7.9%o) and may be similarly explained. A remarkable feature of the Australian data is that the <5 O values are so similar over vast areas, which must imply interaction and probable isotopic equilibrium with enormous bodies of wellbuffered fluids. The abundance of biotite in these
ACKNOWLEDGMENTS Samples were donated by Dr P. A. Arriens (Canberra), Dr L. F. Bettenay (University of Western Australia), Dr R. D. Gee (Geological Survey of Western Australia), and Western Mining Corporation. Oxygen extraction and massspectrometry was done by D. Barnbaum and B. Wood, and E. Thompson maintained the technical facilities. The research was funded by ARGC grant E78150361.
ARCHIBALD, N. J., & BETTENAY, L. F., 1977: Indirect
LONGSTAFFE,
ARCHIBALD, N . J . , BETTENAY, L . F . , BINNS, R . A . , GROVES, D . I . , & GUNTHORPE, R . J . , 1 9 7 8 : T h e
LONGSTAFFE, F . J . , & SCHWARCZ, H . P . , 1 9 7 7 :
18
18
18
l s
18
18
ls
evidence for tectonic reactivation of a pre-greenstone sialic basement in Western Australia. Earth planet. Sci. Lett., 33, 370-378.
evolution of Archaean greenstone terrains, Eastern Goldfields Province, Western Australia. Precamb. Res., 6, 103-131. ARRIENS, P. A . , 1971: The Archaean geochronology of Australia. Spec. Pubisgeol. Soc. Aust., 3, 11-23.
CLAYTON, R . N . , & MAYEDA, T . K . , 1 9 6 3 : T h e u s e o f
bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. cosmochim. Acta, 27, 43-52. FOURCADE, S., & JAVOY, M., 1973: Rapports 0 / 0 dans les roches du vieux socle catazonal d'ln Ouzzal (Sahara algerien). Contrib. Mineral Petrol., 42, 235-244. GEE, R. D., 1979: Explanatory notes on the Southern Cross 1:250000 Geological Sheet, Western Australia. Rec. geol. Surv. West. Aust., 1979/5. GOLDING, S. D . , & WILSON, A. F . , 1 9 8 0 : The application of oxygen isotope studies to the occurrence of gold mineralization in eastern Australia. Australas. l.M.M. Conf., New Zealand, May 1980, 6 7 - 8 0 . , 8
, 6
2
I 8
I 6
18
F. J., 1979: The oxygen-isotope geochemistry of Archaean granitoids; in Barker, F. (Ed.) Trondhjemites, Dacites and Related Rocks, 363-399. Elsevier Press, Amsterdam. , 8
O/ 0 1 6
of Archaean clastic metasedimentary rocks: a petrogenetic indicator for Archaean gneisses? Geochim. cosmochim. Acta, 41, 1 3 0 3 - 1 3 1 2 . MATSUHISA, Y . , 1 9 7 4 : 0 / 0 ratio of N B S - 2 8 and some silicate reference samples. Geochem. J I., 8, 1 8
, 6
103-107. MUEHLENBACHS, K., & CLAYTON, R. N., 1976: Oxygen
isotope composition of the oceanic crust and its bearing on seawater. J. geophys. Res., 81, 4365-4369. OVERSBY, V . M . , 1 9 7 5 : Lead isotopic systematies and ages of Archaean acid intrusives in the KalgoorlieNorseman area, Western Australia. Geochim. cosmochim. Acta, 39, 1 1 0 7 - 1 1 2 5 . RODDICK, J. C. M., 1974: Responses of Strontium Isotopes to some Crustal Processes. Ph.D. Thesis, Aust. Nat. Univ, [unpublished]. SHIEH, Y. N., & S C H W A R C Z , H. P., 1974: Oxygen isotope studies of granite and migmatite, Grenville Province of Ontario, Canada. Geochim. cosmochim. Acta, 38, 21-45.
G E O C H E M I C A L ATTRIBUTES OF GRANITOIDS
405
1978: The oxygen isotope composition of the surface crystalline rocks of the Canadian Shield. Can. J. Earth Sci., 15, 1773-1782.
WILSON, A . F . , & BAKSI, A . K . , 1978: S i g n i f i c a n c e o f
TOURET, J:, 1971: Le facies granulite en Norvege meridionale. II: Les inclusions fluides. Lithos, 4, 423-436.
WILSON, A . F . , & GREEN, D . C . , 1971: T h e u s e o f o x y -
oxygen isotope studies on granulite facies rocks; in Windley, B. F., & Naqvi, S. M. (Eds) Archaean Geochemistry, 289-301. Elsevier, Amsterdam. gen isotopes for geothermometry of Proterozoic and Archaean granulites. Spec. Pubis geol. Soc. Aust.,3,
389-400.
ARCHAEAN MINERALIZATION
THE METALLOGENESIS OF ARCHAEAN BASE-METAL DEPOSITS IN WESTERN AUSTRALIA R. J. Marston & D. I. Groves 1
1
2
Geological Survey of Western Australia, Perth, Western Australia 6000 Present Address: Aurex Pty. Ltd., Perth, Western Australia 6000 Department of Geology, University of Western Australia, Nedlands, Western Australia 6009 2
ABSTRACT Lithological, stratigraphic and tectonic controls determine the distribution of base-metal deposits in the Pilbara and Yilgarn Blocks. Similar deposits occur in each block, but there may be little or no age overlap of either mineralization or host rocks between the two blocks. Magmatic activity was of prime importance in ore genesis and consequently the nature of the evolution of the Archaean mantle was a fundamental control on metallogenic evolution. The Pilbara Block formed mainly in early Archaean times when low thermal activity in the mantle resulted in relative crustal stability and small volcanic basins. A low level of chemical differentiation in the mantle, a lack of crustal lineaments, and the presence of shallow-water, oxidized environments all inhibited base-metal mineralization. Most of the Yilgarn Block evolved in late Archaean times when high thermal activity and chemical differentiation in the mantle resulted in the genesis of S-rich magmas and in larger, linear, well-mineralized volcanic basins. Crustal lineaments allowed access of gaseous and high-temperature magmas into reducing basins. The importance of a late Archaean metallogenesis is evident in Western Australia and Archaean terrains elsewhere in the world.
exposed areas of the Pilbara and Yilgarn Blocks are respectively about 60000 and 650000 km ; both figures would increase substantially if extensions buried under Proterozoic and Phanerozoic platform sequences were included. The Pilbara Block (Fig. 1) consists of small, arcuate volcanosedimentary belts of diverse strike and two larger east-to-northeast-striking belts of turbiditic sedimentary rocks east of Nullagine and east of Whim Creek. All the supracrustal rocks of the Pilbara were probably deposited more than 3.2b.y. ago. Major strike faults of regional extent are absent except for one example in the west. Regional metamorphism is commonly greenschist facies or below but locally attains amphibolite facies. Ultramafic volcanic and associated intrusive rocks are poorly developed whereas mafic volcanic rocks abound. Layered or multiple gabbroid intrusions are common in the western two-thirds of the block. Felsic volcanic complexes interfinger with mafic volcanics (e.g. Barley, 1980) and occur at low and high stratigraphic levels. Sedimentary rocks GEOLOGICAL SETTING within and above the main volcanic sequence inThe Archaean shield of the State consists of the clude cherty rocks of various origins and fluvial Pilbara and Yilgarn Blocks of which geological clastic rocks all suggestive of shallow-water accounts are given in this volume by Hickman depositional environments (Barley et al., 1979; (1981) and Gee et al. (1981) respectively. The Eriksson, 1981). However, some fluvial clastics Spec. Pubis geol. Soc. Aust., 7 (1981) INTRODUCTION This paper represents the first modern, albeit brief, metallogenic synthesis of the Archaean shield of the State, although there are summaries of individual regions (e.g. Geological Survey of Western Australia, 1975; Gemuts & Theron, 1975; Hickman, in press) and reference to the metallogeny of this shield has been made in global reviews (e.g. Watson, 1973; Lambert & Groves, 1981). Descriptive summaries of nickel, copper, vanadium, tungsten, molybdenum and tin deposits are given by Marston (in press), Marston (1979), Baxter (1978), and Blockley (1980). This paper attempts to synthesize these and other observations (Figs 1, 2, 4 and Table I) and to view the mineralization in the context of crustal evolution of the Pilbara and Yilgarn Blocks (Figs 3, 5). Studies of most deposits and of their tectonic setting and geochronology are at an early stage and therefore many conclusions are tentative.
2
117 00' INDIAN
YOUNGER METAMORPHIC
A N D PLUTONIC
ROCKS
_ Em _ P r o t e r o z o i c , includes r e w o r k e d A r c h a e a n rocks
R E F E R E N C E FOR Y I L G A R N A N D P I L B A R A BLOCKS
T Y P E S OF BASE M E T A L DEPOSITS 1 A Fe-Ni-Cu i n t r u s i v e d u n i t e - a s s o c i a t e d
M E T A M O R P H O S E D A R C H A E A N S U P R A C R U S T A L ROCKS 9 u p r a c r u s t a l belts
j
A R C H A E A N G R A N I T O I D A N D GNEISSIC ROCKS + + + + + + + +
it t t +
2 a Fe-Ni-Cu volcanic p e r i d o t i t e - a s s o c i a t e d
Leucogranite to adamellite, post-tectonic, fractionated G r a n i t o i d , m i g m a t i t e and g r a n i t o i d gneiss
Felsic v o l c a n i c c o m p l e x e s
Gneiss, schist and a m p h i b o l i t e of h i g h m e t a m o r p h i c grade, i n c l u d e s m e t a s e d i m e n t a r y r e m n a n t s
j Layered gabbro- peridotite intrusions
3 A Fe-Ni-Cu gabbroid-associated 4 ® Fe-Ni-Cu layered v o l c a n o s e d i m e n t a r y
W h i m Creek G r o u p : m a f i c t o felsic m e t a v o l c a n i c r o c k s of c a l c - a l k a l i n e a f f i n i t y (Pilbara B l o c k o n l y )
5 o Fe-Ni-Cu + A u v e i n - t y p e arsenical 6 d Fe-Ti-V stratabound gabbroid
M a j o r t e c t o n i c l i n e a m e n t , f a u l t or shear zone
7•
F e - C u ± A u vein or shear (strike i n d i c a t e d ) 104
j i f ^ j Olivine - r i c h u l t r a m a f i c r o c k s
Fig. 1.
E a r l y Proterozoic g a b b r o i d d y k e ( Y i l g a r n B l o c k o n l y )
Fe-Cu-Zn+Pb stratabound volcanosedimentary Fe-Cu±Mo stockwork
,
Sn-Ta v e i n , p e g m a t i t e W ± M o greisen, vein
Solid geology of the Pilbara Block showing distribution of base-metal deposits (geology after H i c k m a n , in press).
METALLOGENESIS IN WESTERN AUSTRALIA 411 also pass laterally into deep-water turbiditic mineralization. The deposits strongly reflect the fades (Eriksson, 1981). Sulphidic sedimentary predominance of magmatic activity in the rocks appear to be rare. Hickman (1981) argues Archaean and the bulk of the economically imthat the volcanosedimentary belts exhibit strati- portant mineralization occurs as stratabound graphic consistency across the entire block. layers or lenses of disseminated or massive ore The Yilgarn Block (Fig. 2) contains linear, minerals in volcanic environments or within north-northwest-striking, volcanosedimentary related intrusives. The distribution of most belts in the Norseman-Wiluna belt and the deposits dominated by iron sulphides and oxides Southern Cross-Forrestania region. In the Cen- is related to stratigraphic and/or tectonic contral Yilgarn and Murchison regions, arcuate trols in addition to lithological factors. The key volcano-sedimentary belts of variable strike (i.e. relationships are summarised in Table I and illussimitar to the Pilbara) prevail. The rocks in all trated diagrammatically in Figure 3. these belts were probably deposited less than 3.0 The Pilbara Block is notable for an abundance b.y. ago. The remaining western areas of the of small, low- to high-grade deposits and a block are dominated by granitoids and gneisses, paucity of large deposits irrespective of grade. and metasedimentary remnants of high meta- Mineralization is irregularly distributed in the morphic grade. The volcanosedimentary belts Yilgarn Block: the Central Yilgarn region is parcommonly exhibit greenschist or low amphibolite ticularly impoverished, whereas the Norsemanfacies of metamorphism. Olivine-rich ultramafics Wiluna belt is rich in base-metal deposits. (>38% MgO, volatile-free) of volcanic and Nickel deposits are concentrated in the linear, intrusive types abound in the linear belts but are faulted belts of the east and south Yilgarn Block, rare elsewhere. Gabbroidal intrusive complexes and the few present in the Pilbara occur mainly are best developed in the Murchison region. near a major strike fault. Most nickel deposits Felsic volcanics of subaerial and submarine (1-3 in Table I) show correlation between magorigin are most common in the Norseman-Wiluna nesia content of the host ultramafic and the bulk belt. Oxide-type BIF is absent from the Norse- Ni/Cu ratios of contained sulphides (Fig. 4E). man-Wiluna belt (except in the northeast), Pyrrhotite, pentlandite, pyrite, chalcopyrite, whereas bedded sulphidic, cherty sedimentary magnetite and zoned ferrochromite are the comrocks abound in the volcanic sequences. The mon phases, with the sulphides having a resame belt also contains most developments of stricted bulk composition (Fig. 4C). The host stratigraphically high, conglomeratic sedimen- dunites and peridotites are of komatiitic affinity tary rocks. Major strike faults are a key feature with a compositional continuum between volof the belt as they influence the distribution of canic ultramafics and basalts (Fig. 4D). Type 1 ultramafics, felsic volcanics and conglomerates and 2 deposits are found typically at high and low (e.g. Williams, 1974), and also affect tectonic and stratigraphic levels respectively, at the base of metamorphic styles ( e.g. Archibald et al., 1978). thick sills or flows near major strike faults (Fig. Only in the Murchison region is stratigraphic 3). Type 3 nickel deposits are few and occur in consistency evident (Gee et al., 1981), where, as small gabbroid intrusions which appear to be of in the Pilbara, the accompaniments are: (a) tholeiitic affinity in the few examples studied poorly developed ultramafics, (b) oxide- rather (e.g. Purvis etal1972). than sulphide-type sediments in volcanic Only three layered volcanosedimentary nickel sequences, (c) numerous gabbroid intrusions, and deposits (4 in Table I) are recognized (Sherlock (d) a lack of major strike faults. Stocks of post- Bay, Cruickshank, F shoot Windarra): all are in tectonic fractionated leucogranitoid are common metamorphically reconstituted, bedded Ca-Fein the east Pilbara, but are small and rare in the Mg-rich rocks of probable volcanic-exhalative Magnetite may be abundant, importantly Yilgarn with the notable exception of the fluorite- origin. ferrochromite and chromite are absent, and rich Mount Boreas Adamellite, north of Laverton pyrrhotite/pentlandite ratios are higher than for (Bunting & Chin, 1979). other nickel deposits. Rare arsenical vein-type nickel deposits (e.g. Mt Martin, Bamboo) are DISTRIBUTION AND PETROLOGY OF located in or near fractured and carbonated ultramafics. Quartz-carbonate veinlets and disseminaDEPOSITS The deposits are classified into eleven types in tions contain Fe-Ni-Cu sulphides and arsenides Table I and their distribution is shown in Figures and some erratically distributed gold. 1 and 2. The criteria used in classification are: (i) Stratiform deposits of Fe-Ti-V oxides are the nature of the host rock and enclosing found in layered intrusions of tholeiitic affinity sequence (or intrusive complex), and (ii) the (Fig. 4B) in the northwest Yilgarn, west Pilbara geometry and contact relationships of the and in a dyke at Coates, near Perth. The oxides,
412
R . J . M A R S T O N & D . I. G R O V E S TABLE I
Metallogenic summary of Archaean base-metal deposits in Western Australia Ore Genesis
Ore Minerals
Host or Associated Rocks
(pre-metamorphism)
Possible Controls
1) Fe-Ni-Cu su1ph ides
"intrusive dunite" =dunite-olivinite, M g O , in subconcordant lenses mainly: deposits have bulk Ni/Cu >19 and form lens-like shoots
thickest part of lens, stratigraphy, strike faults
separation during early crystallization of sulphur-saturated komatiitic magma and settling of sulphide liquid; contamination with crustal sulphides and metamorphic modification may be important
"volcanic peridotite" = olivine peridotitedunite 38-^5% MgO, in thick flows; deposits have bulk Ni/Cu = 7~19 and form lenses at base of flows
seafloor topography, stratigraphy, strike faults
as for (1)
3)
"gabbroid" = gabbronori te-pyroxenite-peridotite in layered intrusions; deposits have bulk Ni/Cu <7 and form layers and pipes
tholeiitic intrusions, linear tectonic belts
as for (2) but pa rent
M
bedded ca1cium-iron-magnesium rich sediment in ultramafic, mafic or mafic-felsic volcanics; deposits have variable bulk Ni/Cu and form layers
linear eruptive zones with explosive vulcanism?
volcanic exhalative?; possible decarbonatio during metamorphism (Sherlock Bay)
5) Fe-Ni-Cu sulphides arsenides + Au
veins in ultramafics, or other supracrustals near ultramafics
faults, strongly deformed and fractured zones
metamorphic-hyd.rothernial, concentration by carbonate-rich fluids
6) Fe-Ti-V oxides
magnetite-gabbronorite in gabbroid-anorthosite layered/multiple intrusions or dykes; deposits form extensive layers and lenses
tholeiitic intrusions, more stable areas
intermittent pressure-temperature fluctuations in a crystallizing basaltic magma
7) Fe-Cu-Zn ± Pb sulphides
heterogeneous felsic volcanics and chert, less commonly carbonaceous pelite, mafic volcanic rocks commonly associated; deposits form layers, lenses and stockworks
end of volcanic cycle, explosive vulcanism, linear eruptive zones, highlevel plutons, depth of seawater (deep?)
hydrothermal leaching of fractured volcanic piles in areas of steep thermal gradient, precipitation on contact with seawater; metamorphic modification may be important
ith tholeiitic magma as
8) Fe-Cu sulphides ± Au
vein quartz commonly in mafic to ultramafic rocks, less commonly in sedimentary and felsic volcanic rocks
faults and fractures, moderate-grade regional 'metamorphism
leaching of syngenetic metals; metamorphic remobi1ization of (7)
9) Fe-Cu ±Mo
quartz/feldspar phyric hypabyssal granodiorite and enveloping volcanic rocks; deposits form stockworks
flanks of major granitoid domes, synvolcanic granitoids, depth of seawater (shallow?)
release of ore-bearing juvenile and meteoric fluids into fracture system during degassing of crystallizing felsic magma at shallow depths
10) Sn-Ta oxides
quartz-feldspar-mica pegmatite spatially associated with post-tectonic granitoid, miarolitic granitoid and/or mafic remnants; deposits may concentrate at metasomatic reaction zone contacts
fractionated granitoids of various ages in upper parts of domal granitoid complexes
fractional crystallization of granitoid magma gives a water-saturated melt enriched in incompatible elements, largely expelled into fractures in country roeks
11) W ± Mo oxi desulph ide
quartz-rich pegmatite or greisen in or near leucogranitoids, as sheets or veins; deposits may concentrate at metasomatic reaction zone contacts or form stockworks in country rocks
as for (10)
as for (10)
N.B.:
Deposits 1, 2 , 3 , A , 6 and 7 can be called stratabound 1979)
as a whol e , and may be i n part stratiform
chiefly vanadiferous magnetite with exsolution lamellae of ilmenite and subordinate, separate ilmenite (Hudson, 1977) commonly occur as extensive, disseminated to massive layers in rhythmically layered gabbronorite and anorthositic gabbroid. The distribution of Fe-Cu-Zn ± P b deposits in the Norseman-Wiluna belt may relate to major faults (in which case more remain to be found), but such an association is not apparent in the Murchison region where these deposits seem more numerous despite a lower proportion of felsic volcanic complexes. Some mineralized felsic volcanic formations appear to have overlain mafic formations which are at the base of the sequence (e.g.Golden Grove), but deposits are so widely dispersed that a regional stratigraphic control is equivocal. In the Pilbara, deposits are found at low (e.g. Big Stubby) and high stratigraphic levels (e.g. Whim Creek).
(terms used as defined by Marston,
The host sequences for Fe-Cu-Zn ± Pb deposits are predominantly felsic (e.g. Golden Grove), but may also be mafic (e.g. Mt Mulcahy) or mixed (e.g. Anaconda-Nangeroo), with the mineralization confined to fragmental felsic rocks, chert or pelitic sediment (Fig. 3). A calc-alkaline affinity is indicated for some host sequences (Fig. 4B) but data are sparse. In the Yilgarn Block, identified calc-alkaline complexes are conspicuous for their lack of base-metal mineralization, which may partly relate to a dominantly subaerial character (e.g. Giles, 1981). Cobaltiferous pyrite and pyrrhotite are the dominant opaque phases (magnetite may be abundant, as at Golden Grove), accompanied by chalcopyrite, Fe-sphalerite and minor galena, arsenopyrite and cubanite. Bulk Cu/Zn ratios show a wide range with some Pbrich deposits in the Pilbara Block (Fig. 4A). The stratigraphic top of many deposits is Zn-rich and contains most of any lead present. Wall-rock
METALLOGENESIS IN WESTERN AUSTRALIA
2.
413
Solid geology of the Yilgarn Block showing distribution of base-metal deposits (geology after Geological Map of Western Australia, 1979 by Geological Survey of Western Australia). Reference on Figure 1.
R. J. M A R S T O N
T-XTT^l ' 1 M . I • g ^ M M M ^ ^ ^ ^
& D.
I.
GROVES
•
Intrusive dunite - olivinite
, ° o o Coarse clastic polymictic sediments Wacke and shale
n
Mafic-felsic volcaniclastie sediments
^
Chert, BIF, sulphidic sediment
| . A | Fine-coarse felsic fragmentals
Post-tectonic leucogranite <
|ujjj
. 3.
Pegmatite (mineralized)
v
Komatiitic
basalt
| * ^ x j Layered gabbroid complex Felsic flow rocks
^ Tholeiitic basalt
Miarolitic granite - pegmatite
| ~1~ | Post-tectonic mafic granitoid
| Ore deposit
• •ZJLI
Sub-volcanic granitoid
® ® Pre- or syntectonic granitoid
S c h e m a t i c sections d e p i c t i n g t h e g e o l o g i c a l setting o f d e p o s i t s w i t h e x a m p l e s i n d i c a t e d . N u m b e r s refer to d e p o s i t types in T a b l e I . D o m a l g r a n i t o i d - g n e i s s c o m p l e x i n l o w e r c a r t o o n is b a s e d o n t h e S h a w B a t h o l i t h , east P i l b a r a B l o c k .
415 elements, the Pilbara granitoids that have been related to cassiterite-tantalite-columbite-albite pegmatites are comparable with "tin-mineralizing granites" in New South Wales (Juniper & Kleeman, 1979) and Tischendorf's (1977) "specialized granite", although the tin and fluorine contents of these Pilbara granitoids (Blockley, 1980) are relatively low. The same Pilbara granitoids contain few pegmatites and lack miarolitic cavities. However, adjacent small bodies of fluorite-bearing miarolitic granite that are petrologically similar to the cassiterite pegmatites are found in the Shaw Batholith near the Cooglegong Adamellite and may have a closer genetic link with mineralization than the larger granitoid bodies. Small pegmatitic deposits of tungsten and molybdenum are known from the eastern Pilbara, but the single important deposit is related to the syntectonic Mulgine stock in the Murchison region of the Yilgarn Block (Baxter, 1978). The mineralization at Mulgine is in a metamorphosed quartz-microcline-muscovite greisen sheet and a related contact metasomatic reaction zone, and in adjacent vein quartz stockworks, all marginal to a small intrusion of muscovite ± biotite granite. Scheelite, molybdenite, fluorite, pyrite and chalcopyrite are conspicuous minerals in the greisen (particularly at metasomatic contacts with ultramafic rocks) and veins.
METALLOGENESIS IN WESTERN AUSTRALIA
alteration involving Fe-Mg-Al addition is conspicuous at some deposits (e.g. Mons Cupri). Vein deposits of Fe-Cu sulphides ±Au are numerous but small, most having been mined because of supergene enrichment. Most deposits are in mafic to ultramafic rocks and appear to relate to local stratigraphic, tectonic and metamorphic control. The deposits at Ravensthorpe are exceptional in their abundance, spatial restriction to narrow felsic to mafic volcanic units, and association with small stratabound FeCu-Zn deposits. These features and petrological characters (Marston, 1979) suggest that the vein deposits may represent metamorphically remobilized stratabound deposits. Little is known about the petrology of most other Fe-Cu sulphide ± Au vein deposits except that copper contents are commonly low (<2°7o) in the primary zone, pyrite or pyrrhotite being the dominant sulphide phase. Stockwork Fe-Cu ± Mo sulphide deposits are so far known only from the eastern Pilbara Block, with one large deposit (Coppin Gap) and several small ones (Fig. 1). The deposits are in and around porphyritic hypabyssal granodioritic intrusions which contain rounded and embayed quartz and subhedral feldspar phenocrysts. Quartz-pyrite-chalcopyrite-molybdenite ± carbonate veins form the stockworks which are associated with chloritic, sericitic and silicic wallrock alteration. Small stratabound Fe-Cu-Zn deposits may be associated, and a primitive "porphyry-style" of synvolcanic emplacement and mineralization is indicated. The eastern Pilbara Block also contains numerous occurrences and small deposits of Sn and Ta oxides in pegmatites (largely exploited in derived eluvial-alluvial concentrations), many of which are peripheral to, but not within, fractionated post-tectonic leucogranitoid stocks emplaced at high structural levels. Some Ta-rich deposits (e.g. Wodgina) are distant from these stocks. In contrast similar pegmatites in the Yilgarn Block are widely scattered, but are generally found in mafic rocks without such stocks nearby. The Yilgarn pegmatites have been metamorphosed and tin may concentrate in marginal metasomatic reaction zones. Greenbushes is the only Sn-Ta deposit of note, but production from this pegmatite body exceeds that from the whole Pilbara Block and recent company reports suggest very large Ta reserves. The Pilbara granitoids that are spatially associated with Sn-Ta deposits are more siliceous, alkali-rich and have high Na/Ca ratios compared with the Gobbos granodiorite, a typical representative of an intrusion related to Fe-Cu ± Mo mineralization (Fig. 4F and G). In terms of major
ORE GENESIS AND CRUSTAL EVOLUTION These topics are summarized in Table I and Figure 5. It must be emphasized that considerable uncertainty currently attaches to details of the genesis of most deposit types, even for nickel deposits (e.g. Groves et al., 1979) which as a type are by far the best documented. The major factors influencing Archaean metallogenesis are the generation of specific types of metalliferous acid or basic or ultrabasic magmas in the upper mantle or lower crust, their movement into the upper crust where fractionation may take place, and their emplacement in appropriate stratigraphic or tectonic environments. Hydrothermal leaching of more permeable (mainly felsic) volcanic piles is probably important in the genesis of Fe-Cu-Zn and Fe-Cu ± Mo deposits although it is not clear to what extent metals were derived from fractionating magmas. There is also the possibility of remobilization and further extraction of syngenetic metals during later regional metamorphism, for example in the formation of FeNi-Cu ± Au and Fe-Cu ± Au vein deposits. A strong genetic link between "aluminiumundepleted" (Nesbitt et al., 1979) komatiitic
R. J . M A R S T O N & D . I. G R O V E S
Fe-Cu-Zn
Canadian Archaean deposits (Sangster, 1972)
DEPOSITS
(Marston, 1979)
Fe-Ti-V
Canadian Phanerozoic deposits (Sangster, 1972)
o Halls Creek Province • Pilbara Block
>
x Yilgarn Block
/
New South Wales Phanerozoic deposits (Felton etal., 1974) THOLEIITIC
F e - C u - Z n HOSTS • Mons Cupri (Sylvester, pers. comm.).
Fe-Ni-Cu
CALC ALKALINE
HOSTS OF INTRUSIVE DUNITE ASSOCIATED ORES
HOSTS OF VOLCANIC PERIDOTITE ASSOCIATED ORES B ZONES OF KOMATIITIC " PERIDOTITE FLOWS
D E P O S I T S AND HOSTS (Groves and Hudson,1981)
Bravoite
A ZONES OF KOMATIITIC PERIDOTITE FLOWS
Vaesite^v Violarite
Pyrrhotitei
HOSTS
A B Trend of Skaergaard rocks CD Trend of v Windimurra rocks \ A Windimurra (Ahmat, \ pers. comm.) Coates (Hudson, * V 1967)
Olivine fractionation trend
Millerite
KOMATIITE FIELD
Pentlandite
KOMATIITIC BASALT FLOWS/
Heazlewoodite C O M P O S I T I O N A L F I E L D OF MOST SULPHIDE ORES FROM DEPOSITS WITH B U L K Ni>1%
(R. A . Binns^pers. comm.) I I
Olivine + clinopyroxene fractionation trend
THOLEIITIC BASALT FLOWS
GENERAL THOLEIITE FIELD T H O L E I I T E S IN NORSEMAN—WI LUNA BELT
Olivinite Field Mount Keith Goliath ° Mount Hope ° ©Six Mile Q Cosmic B o y Dunjte p oBlack Swan 2 44
GBGabbroid asstd field 42 VP Volcanic peridotite asstd field (•= deposits) 40 ID Intrusive dunite asstd field 38 (o= deposits)
ODigger R o c k s • S o u t h Windarra ## Nepean
am bald a • Wannaway Mount E d w a r d s •Mount Windarra
Olivine Peridotite Field
Post-tectonic granitoids, Pilbara Block (Blockley, 1980) Gobbos granodiorite, Pilbara Block (Barley, 1980) Mulgine granitoids, Yilgarn Block (Collins, 1975) Marda granite-granophyre, Yilgarn Block (Hallberg et al. 1976) Mt. Boreas adamellite, Yilgarn Block (Bunting & Chin, 1979) "Specialized granite" of Tischendorff (1977) Field of "tin-mineralizing granites" New England, N.S.W. (Juniper & Kleeman, 1979)
GRANITOIDS ASSOCIATED WITH S n - T a , W ±Mo, F e - C u ±Mo 1oy DEPOSITS
Fig. 4.
Summary geochemical features of some ores and host or associated rocks.
METALLOGENESIS IN WESTERN AUSTRALIA 417 magmas and volcanic-associated nickel ores is sive oxide- and carbonate-type banded iron forindicated. The host rocks of the dunite-associated mations, (iii) some stratigraphic consistency bedeposits have a composition and spatial relation- tween supracrustal belts and (iv) large, layered, ship (Fig. 3) which is consistent with their being gabbroid intrusions containing Fe-Ti-V deposits. residual accumulations of fractionating komatiA possible explanation for the progressive itic peridotite melts. Nickel deposits and some Fe- decrease in stability of Archaean volcanic basins Cu-Zn deposits show a marked control in their and for an increase in abundance of nickel-associdistribution by linear tectonic features which are ated "aluminium-undepleted" komatijtes in the interpreted as reflecting linear eruptive zones that late Archaean is given by Weaver & Tarney in turn relate to fundamental, long-lived crustal (1979). They propose that the size of rising lineaments. mantle diapirs increased during the Archaean Although geochronological data are sparse and because of increased mantle viscosity (in turn some conflict with geological deductions, we pro- caused by a decline in geothermal gradient as a pose that crustal evolution in Western Australia, result of crustal growth), and that resultant as in Archaean terrains elsewhere, did not favour supracrustal basins became larger, less stable and an important metallogenesis until the late more linear with time. They also suggest that the Archaean (Fig. 5). Comparison of the tectonic abundant late Archaean "aluminium-undevelopment of the older and younger Archaean depleted" ultramafics, which are depleted in terrains of the Pilbara-western Yilgarn Block and Ti0 and light REE, were derived by partial the central-eastern Yilgarn Block respectively, melting of a similarly depleted mantle source. reveals some important differences that may Such a mantle probably developed in the late reflect significant controls on metallogenesis. Archaean due to growth of an upper mantle Similar types of base-metal deposits are found depleted by formation of dominantly basaltic in the Yilgarn and Pilbara Blocks, but there is melts in the early Archaean. A depleted mantle probably no overlap in the age of either the may also allow hotter liquids, richer in MgO, to mineralization or the host rocks in each block. form, thus perhaps explaining the abundance of The blocks also became cratonic units at different olivine-rich ultramafics in the late Archaean. times. The late Archaean komatiitic magmas may Stable crustal conditions in the Pilbara are have been S-rich because of melting of mantle indicated by the apparent broad stratigraphic sulphides in zones of high geothermal gradient continuity (Hickman, in press), the development (e.g. Naldrett, 1973), although contamination by of widespread horizons of shallow-water chemi- crustal sulphides in high-level magma chambers cal sediments or mature clastic rocks (Barley et cannot be discounted (e.g. Groves et al., 1979). all, 1979; Eriksson, 1981), and the poor develop- The weak development of Archaean Sn-Ta and ment of crustal lineaments. Local oxidizing con- W-Mo mineralization may indicate that the ditions in the shallow basins resulted in sulphate associated granitoids were volatile-poor, insufprecipitation (e.g. Lambert et al., 1978) and ficiently fractionated, and derived from a mantle together with limited water depths may have in- or crustal source impoverished in large-ion lithohibited the formation of large volcanogenic Fe- phile elements. A sialic source is most likely and Cu-Zn deposits. The early Archaean supracrustal repeated crustal working may be needed to prorocks of the west Yilgarn Block are mainly clastic duce sufficiently high initial Sn or W contents for and lack volcanic elements or mineralization of further concentration by fractional crystallization (e.g. Lambert & Groves, 1981). Preservation note. In the late Archaean, by contrast, high thermal of the roof zones of granitoids is also important activity in the upper mantle, particularly under if Phanerozoic deposits are a guide, and is less the east Yilgarn Block, generated vigorous and likely in Archaean terrains. Alternatively, tin widespread volcanism. Associated mineralization may not have concentrated into residual phases in probably concentrated in unstable, linear troughs Archaean granitoids where sphene (Sn substitucharacterized by reducing conditions and rapid tion for Ti) and iron oxides (high f0 ?) abound in facies changes in volcanic and interfingering early crystallized granitoids. Few early Archaean terrains show much volcanoclastic sedimentary piles. The crustal lineaments probably promoted the access to the evidence for an important metallogenesis in terms upper crust of both high-temperature komatiitic of base-metal ores (excluding iron). The immagmas and explosive, volatile-rich felsic portance of late Archaean metallogenesis in Ausmagmas. Nevertheless, at the same time, relative tralia and elsewhere (e.g. Canada, India) is stability in the central Yilgarn and Murchison perhaps primarily a function of major geochemiregions is indicated by (i) the lack of olivine-rich cal differentiation in the upper mantle at that ultramafics, (ii) abundant and regionally exten- time (cf. Windley, 1977). 2
2
YILGARN
EVENTS
YILGARN CRATON STABILIZED
BLOCK
TECTONOMETA MORPHIC
00
vWWV
N/W
+ ISOTOPIC HOMOGENIZATION
.>J=L gabbroids
PLUTONIC
mafic dykes
Central Yilgarn gneiss
West Yilgarn gneiss
granitoids pegmatite
AAI
DEPOSITIONAL West Yilgarn supracrustals
volcano sedimentary polymictic supracrustals conglomerate
73
2>
TECTONOMETAMORPHIC
w
PILBARA
BLOCK
•
*
leucogranitoids pegmatite HAMERSLEY
Gorge Creek Group
3.4
3.5
rn m
• ?
pegmatite
*
Warrawoona Group
o 73 O <
mafic dykes
domal granitoids
B A
Rp
SOME ISOTOPIC HOMOGENIZATION
gabbroids
gneisses
DEPOSITIONAL
M / W
+ ISOTOPIC HOMOGENIZATION
PLUTONIC
73 cn H O Z
PILBARA CRATON STABILIZED
Whim Creek Group
3.3
Fortescue Group
3.2
3.1
3.0
2.9
BASIN 1 Hamersley Group
2.6
2.7
b.y.
Fig.
5.
Interpreted
crustal
evolution
of
the
Pilbara
and
Yilgarn
Blocks
with
the
timing
of
mineralization
i n d i c a t e d . D e p o s i t s y m b o l s a r e as f o r F i g u r e 1, a n d s i z e o f s y m b o l s is r e l a t e d t o t h e a b u n d a n c e o f
mineral-
i z a t i o n . G e o c h r o n o l o g i c a l i n t e r p r e t a t i o n is b a s e d o n d a t a f r o m m a n y s o u r c e s ( s u m m a r i z e d b y d e L a e t e r et a/.,
1981).
2.5
2.4
METALLOGENESIS
IN W E S T E R N
ACKNOWLEDGMENTS We wish to thank the following for access to unpublished data: A. L. A h m a t , M. E. Barley, L. F. Bettenay, R. A. Binns, J. G. Blockley, A. H. Hickman, I. B. Lambert and G. C. Sylvester. R. D. Gee is thanked for observations on
ARCHIBALD, N .
J.,
BETTENAY,
L.
F.,
BINNS, R .
GROVES, D . I . , & G U N T H O R P E , R . J . ,
1978:
A., The
evolution of Archaean greenstone terrains, Eastern Goldfields Province, Western Australia. Precamb Res., 6, 103-131. BARLEY, M. E., 1980: The Evolution of Archaean Calcalkaline Volcanics: a Study of the Kelly Greenstone Belt and the McPhee Dome, Eastern Pilbara Block, Western Australia. Ph.D. Thesis, Univ. West. Aust. [unpublished]. BARLEY, M . E . , D U N L O P , J . S . R . , GLOVER, J . E . ,
&
I., 1 9 7 9 : Sedimentary evidence for an Archaean shallow-water volcanic-sedimentary facies, eastern Pilbara Block, Western Australia. Earth planet. Sci. Lett., 43, 7 4 - 8 4 . BAXTER, J. L., 1978: Molybdenum, tungsten, vanadium and chromium in Western Australia. Min. Res. Bull., geol. Surv. W. Aust., 11. BLOCKLEY, J. G., 1980: The tin deposits of Western Australia with special reference to the associated granites. Geol. Surv. W. Aust. min. Res. Bull., 12. BUNTING, J . A., & C H I N , R. J . , 1 9 7 9 : Duketon, Western Australia—1:250,000 Geological Series. Explan. Nates geol. Surv. West. Aust., SG51-14. COLLINS, V. P. M., 1 9 7 5 : The Lithostructural Setting and Controls of Wolfram-Bearing Mineralization at Mt Mulgine, N.W. Yilgarn Craton W.A. B.Sc. Hons Thesis, Univ. West. Aust. [unpublished]. GROVES, D .
M LAETER, J . R . , LIBBY, W . G . , & TRENDALL, A .
F.,
1981: The older Precambrian geochronology of Western Australia. Spec. Pubis geol. Soc. West. Aust., 7, 145-157. ERIKSSON, K. A., 1981: Archaean platform-to-trough sedimentation, east Pilbara Block, Australia. Spec. Pubis geol. Soc. Aust., 7, 235-244. FELTON, E . A . , GILLIGAN, L .
B . , MATSON, C .
R.,
&
STEVENS, B. P. J., 1974: Base-metal mineralization associated with Silurian acid volcanics in the eastern part of the Lachlan fold belt of New South Wales. Rec. geol. Surv. N.S. W., 16, 139-57. GEE, R . D . , BAXTER, J . L . , W I L D E , S . A . , & WILLIAMS,
I. R., 1981: Crustal development in the Archaean Yilgarn Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 43-56. GEMUTS, I., & T H E R O N , A., 1 9 7 5 : The Archaean between Coolgardie and Norseman—stratigraphy and mineralization; in Knight, C. L. (Ed.) Economic Geology of Australia and Papua New Guinea: 1, Metals, 66-74. Australas. Inst. Min. Metall., Melbourne. GEOLOGICAL SURVEY
OF W E S T E R N
AUSTRALIA,
1975:
Geology of Western Australia. Mem. geol. Surv. W. Aust., 2.
AUSTRALIA
419
an early draft. The paper was improved by the comments of D. R. Hudson. It is published with the permission of the Director, Geological Survey of Western Australia, and represents a contribution to I G C P Project 91 "Metallogeny of the Precambrian".
C. W., 1981: Archaean calc-alkaline volcanism in the Eastern Goldfields Province, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 275-286.
GILES,
GROVES, D . I . , BARRETT, F . M . , & M C Q U E E N , K .
G.,
1979: The relative roles of magmatic segregation, volcanic exhalation and regional metamorphism in the generation of volcanic-associated nickel ores of Western Australia. Can. Mineral., 17, 319-336. GROVES, D . I . , & H U D S O N , D . R . , 1 9 8 1 : The nature and origin of Archaean stratabound volcanic-associated nickel-iron-copper sulphide deposits; in Wolf, K. H. (Ed.) Handbook of Strata-Bound and Stratiform Ore Deposits, 9, 3 0 5 - 4 1 0 . Elsevier, Amsterdam. HALLBERG, J . A . , JOHNSTON, C . , & BYE, S . M . ,
1976:
The Archaean Marda igneous complex, Western Australia. Precamb. Res., 3, 1 1 1 - 1 3 6 . H I C K M A N , A. H . , 1981: Crustal evolution of the Pilbara Block, Western Australia. Spec. Pubis geol. Soc. Aust., 7, 57-69. , in press: Geology of the Pilbara Block and its environs. Bull. geol. Surv. W. Aust. H U D S O N , D . R . , 1 9 6 7 : The vanadium-bearing magnetite gabbro at Coates, Western Australia. J. R. Soc. West. Aust., 50, 6 0 - 6 4 . , 1977: Vanadium-bearing titaniferous magnetites of Western Australia; in Symposium on Current Commodity Trends in Mineral Exploration and Evaluation. Geol. Soc. Aust. Specialist Group in Econ. Geol. & Ass. Expl. Geochem., Adelaide. J U N I P E R , D . N., & KLEEMAN, J . D . , 1 9 7 9 : Geochemical characterization of some tin-mineralizing granites of New South Wales. J. geochem. Explor., 11, 321-333. LAMBERT, I . B . , DONNELLY, T . H . , DUNLOP, J . S . R . , & GROVES, D . I . , 1978: Stable isotopic compositions of early Archaean sulphate deposits of probable evaporitic and volcanogenic origins. Nature, Lond.,276, 808-811. LAMBERT, I . B . , & GROVES, D . I . , 1981: Early earth evolution and metallogeny; in Wolf, K. H. (Ed.) Handbook of Strata-bound and Stratiform Ore Deposits, 8, 339-447. Elsevien, Amsterdam. MARSTON, R . J . , 1 9 7 9 : Copper mineralization in Western Australia. Min. Res. Bull., geol. Surv. West. Aust., 13. , in press: Nickel mineralization in Western Australia. Min. Res. Bull., geol. Surv. West. Aust., 14. NALDRETT, A. J., 1973: Nickel sulphide deposits—their classification and genesis, with special emphasis on deposits of volcanic association. Trans. Can. Inst. Min. Metall., 76, 183-201.
420
R. J. MARSTON & D. I. GROVES
NESBITT, R . W . , S U N , S . - S . , & PURVIS, A . C . , 1 9 7 9 :
Komatiites: geochemistry and genesis. Can. Mineral17, 1 6 5 - 1 8 6 .
PURVIS, A . C . , NESBITT, R . W . , & HALLBERG, J . A . ,
1972: The geology of part of the Carr Boyd Rocks complex and its associated nickel mineralization, Western Australia. Econ. Geol., 67, 1093-1113. SANGSTER, D. F., 1972: Precambrian volcanogenic massive sulphide deposits in Canada: a review. Pap. geol. Surv. Can., 72-22. TESCHENDORF, G., 1 9 7 7 : Geochemical and petrographic characteristics of silicic magmatic rocks associated with rare-element mineralization; in Stemprock, M., Burnol, L., & Tischendorf, G. (Eds) Metallization Associated with Acid Magmatism, 2, 4 1 - 9 6 . Geol. Surv., Prague.
WATSON, J . , 1 9 7 3 : Influence of crustal evolution on ore
deposition. Trans. Inst. Min. Metall. Lond., 82, B107-113.
WEAVER, B. L., & T A R N E Y , J., 1979: Thermal aspects of
komatiite generation and greenstone belt models. Nature, Lond., 279, 689-692.
I. R., 1974: Structural subdivision of the Eastern Goldfields Province, Yilgarn Block. Ann. Rep. geol. Surv. West. Aust.for 1973, 53-59.
WILLIAMS,
WINDLEY, B. F., 1977: The crust-mantle boundary in
space and time. J. geol. Soc., 134, 99-102.
A PRELIMINARY STUDY OF SULPHUR ISOTOPES AND ORE GENESIS AT THE GOLDEN GROVE COPPER DEPOSIT, WESTERN AUSTRALIA P. K. Seccombe & K. M. Frater Department of Geology, University of Newcastle, New South Wales 2308 School of Applied Geology, South Australian Institute of Technology, Box 1, P.O., Ingle Farm, South Australia 5098 1
1 2
2
ABSTRACT £ SCDT values for massive chalcopyrite-magnetite ore lenses comprising the Archaean strata-bound Golden Grove deposit range from - 1.4 per mil to + 4.4 per mil (mean + 1.0 per mil). Similar distributions are obtained for disseminated sulphides from the Mineralized Horizon (range - 2 . 5 to +4.0, mean +0.9) and the zinc-rich Hanging-wall Chert (range - 0 . 1 to +4.3, mean +1.6). Isotopic disequilibrium amongst co-existing sulphides stems from depositional or diagenetic effects and is preserved through greenschist-facies metamorphism. A two per mil depletion in the mean 6 S value for the stratigraphically highest copper lens represents a temporary change to high f 0 in the ore fluid during sulphide deposition. The data support a model of sea-floor sulphide accumulation resulting from intermittent fumarolic activity in an active tectonic environment. Magmatic (juvenile) sulphur is indicated to be the dominant sulphur source in the ore fluid. 34
34
2
INTRODUCTION The Golden Grove volcanogenic Cu-Zn prospect is situated 500 km NNE of Perth in the Archaean Yilgarn block of the Western Australian Shield (Fig. 1). This Cu-Zn sulphide deposit is unique in the Yilgarn Block and one of the few potentially economic deposits of Cu-Zn sulphides known from the Archaean of Western Australia. Reserve potential of the main mineralization is 13.5 million metric tons of 3.59% Cu, 0.12%Zn and 13.4ppm Ag. Because of its Cu-Zn metal content and stratabound acid volcanic association, the Golden Grove mineralization has been compared (Frater, 1978) with other massive-sulphide deposits in similar volcanic terrains elsewhere, in particular, those of the Superior Province, Canada (e.g. Sangster & Scott, 1976) and the Kuroko ores of Japan (e.g. Lambert & Sato, 1974). Many similarities and a few differences, the latter relating to either mineralogy or sulphide distribution, emerge in the comparison. The present study tests a model of ore deposition suggested by Frater (1978) based on structural, penological and geochemical studies at Golden Grove. In this preliminary report sulphur-isotope data for 70 sulphide mineral separates are used to resolve more closely the environment of ore deposition and the nature of the Spec. Pubis geol. Soc. Aust., 7 (1981)
ore fluid. The data are discussed in terms of the following aspects of ore formation or geological setting at Golden Grove: (1) The source of sulphur and the degree of seawater-hydrothermal ore fluid mixing for each of the two main types of mineralization (pyritemagnetite-chalcopyrite; pyrite-sphalerite). (2) The degree of isotopic equilibration amongst co-existing sulphides and the effects of diagenetic and metamorphic processes on the sulphur isotopic distribution. (3) Trends, if any, in the physico-chemical environment during ore formation. (4) The spatial separation of copper and zinc mineralization. REGIONAL GEOLOGY The greenstone belt that includes the Golden Grove prospect is composed of two volcanosedimentary systems separated by a regional unconformity and flanked by granitic diapirs. The belt is intruded by Archaean ultrabasic, basic and acid igneous rocks as well as post-Archaean dolerite dykes. Dominant regional structure around Golden Grove is a shallow-plunging, very tight SSE-trending Fj syncline with a steeply inclined axial surface. Most rocks have been subjected to lower greenschist-facies regional metamorphism. The regional stratigraphic succession
P. K. SECCOMBE & K. M. FRATER
422
Igneous Intrusive Rocks Greisen -f-|
A
Granitoid
| Mafic layered
|
|
sill
Ultramafic
Volcanic and Sedimentary Rocki j
Unit 6 : l u t i t e s . minor
lo0o°ol
^ : rudites,
arenites
arenites
Unconformity p""""]
Unit 3 : sedimentary rocks,
basalts
Banded-iron formation
4" -b
Unit 2 : v o l c a n o g e n i c
sedimentary
rocks
4Unit 1 : Golden G r o v e v o l c a n o c l a s t i c
rocks
-h -h
R o c k - u n i t boundary Unconformity S t r i k e and dip of b e d d i n g
X
S t r i k e of
vertical
beds
Fault
Inferred
fault
S y n c l i n e . axial t r a c e and plunge d i r e c t i o n
MAP
Fig. 1.
LOCATION
Interpretative regional geology, south-east section of the Yalgoo Greenstone Belt.
SULPHUR ISOTOPES AND ORE GENESIS
in the vicinity of Golden Grove is summarized in the legend of Figure 1. The Golden Grove mineralization occurs in the lowermost unit (Unit 1). GOLDEN GROVE STRATIGRAPHY AND STRUCTURE Stratigraphy in the Golden Grove volcanoclastics (Unit 1, Fig. 1), in the vicinity of mineralization, is shown in Figure 2. The Mineralized Horizon consists of bedded to massive, fine- to coarse-grained felsic tuffs, lapilli tuffs, cherts, undifferentiated volcanogenic rocks and minor flows. Four high-level Archaean rock types intrude the volcanoclastics: granitic dykes and veins (probably related to the large batholith on the margin of the greenstone belt), dacite, diabase md porphyritic acid dykes. ROCK UNIT VoVoVoVoVo\toV oVoVoVoVoVoVo Vo VoVoVoVoVoV V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V
V V V V V V V V V V V V V
FELDSPATHIC VOLCORUDITE volcorudite or lahar
Maximum Thickness (metres) 50
UPPER QUARTZ-CRYSTAL TUFF tuff, quartz-sericite schist, volcorudite
150
^ " H A N G I N G - W A L L CHERT lutite, chert, tuff, volcarenite, breccia. massive sulphide
50
¥ v v v_v_w
Iff
SrJjpf?^ VV V V V V V V V V V V V V V V V V vv V V V V V V
y v v v v v v V T f V f Y l
v V V V V v v
MINERALIZED HORIZON tuff, lapilli tuff. chert, minor volcanic flows, massive magnetite and sulphide
FOOTWALL LAPILLI TUFF tuff, lapilli tuff "" LOWER BEDDED VOLCANOGENIC ROCKS tuff, lapilli tuff, arenite
150
100
60
w w w
W V W V V VV VV V V
v V vv V vv V VV VV V V V V V V V V
LOWER QUARTZ-CRYSTAL TUFF lithic tuff
200
w w w
V V V VV V V V V V V V V \j V v v
vv v
,y,y v v V v..
w'W'v''v'"V vv vv v v
+++++++++ Fig. 2.
BASAL TUFF volcarenite tuff
INN >1UU
(net •ncounttrad in drill eorm)
GRANITOID
Stratigraphic subdivision of Archaean rocks in the area of the Golden Grove Prospect.
Bedding (S0) is constant where preserved, and generally trends 127° and dips sub-vertically. Mesoscopic folds of three generations occur at Golden Grove, each becoming progressively more open in style. Steeply-inclined axial-plane cleavages accompany the two later fold phases but are conspicuously absent from the tight phase of folding of an earlier generation. These earlygeneration folds may be soft-sediment structures or, alternatively, associated with the regional Fj syncline. Intraformational brecciation, bedding
423
scour, truncated bedding, angular unconformities, soft-sediment transposition on fault surfaces and intraformational faulting and folding are soft-sediment deformation structures common to the Mineralized Horizon and Hanging-wall Chert. MINERALIZATION Zoning The mineralization at Golden Grove is stratabound, being confined essentially to the Mineralized Horizon and Hanging-wall Chert. In the Main Mineralized Zone, magnetite and sulphides (pyrite with subordinate chalcopyrite, pyrrhotite and sphalerite) occur in approximately equal abundance. An alteration assemblage of chlorite, quartz, ankerite and talc is associated with mineralization. There are two major ore-mineral associations, one Cu-rich and one Zn-rich. The Zn-rich mineralization can be further subdivided into ZnCu and Zn-Pb groups. A rough zoning can be recognized, consisting of a Cu-rich zone near the footwall, passing laterally outward and vertically upward into a Zn-Cu and ultimately to a Zn-Pb zone. Morphology The Main-Zone mineralization consists of a number of tabular or lens-shaped pods that coalesce and inter finger. These lenses of massive mineralization contain magnetite and pyrite, with chalcopyrite, sphalerite and pyrrhotite forming only a subordinate proportion. Figure 3 outlines schematically the location of massive mineralization in the Main Zone and Hanging-wall Chert and Figure 4 illustrates a typical geological section through the central portion of the Main Zone. Chalcopyrite occurs mainly at the base of the Main Zone; its distribution bears little relationship to the massive lenses shown in Figure 3. Mineralization in the Hanging-wall Chert is principally pyrite and sphalerite, either massive, bedded or veined. Ore Minerals Major ore minerals are magnetite, pyrite, chalcopyrite, pyrrhotite, sphalerite and galena. Minor ore minerals include tetrahedrite, arsenopyrite, pentlandite, valleriite, tellurobismuthite and cassiterite. Magnetite and pyrite show primary growth and overgrowth textures and have responded to deformation by brittle failure. Most primary growth textures have been destroyed in the more ductile minerals chalcopyrite, pyrrhotite, sphalerite and galena, which show varying degrees of recrystallization resulting from the metamorphic event.
P. K. SECCOMBE & K. M. FRATER
424
, >> I—, P/Py.'Sp.Gn
K
I
INTRUSIVE [ Q
_
Py.Sp
ROCKS
Dacite
MINERALIZATION |
Massive
sulphide-magnetite
VOLCANOCLASTIC fX 7 ] f^El
Upper
Hanging-wall
[y«j
Mineralized
( 2
Footwall
17800N I
lens
ROCKS
ciuartz-crystal
tuff
Mt
magnetite
Py
pyrite
Cpy
chert
ABBREVIATIONS
chalcopyrite pyrrhotite
horizon
lapilli
ORE-MINERAL
sphalerite
tuff
galena
1
18200 •
18000
I
18600
I
L
18800
19000N
Fig. 3. Schematic diagram of the distribution of massive ore mineralization - 200 m R.L., Golden Grove.
Overall, the <5 S distribution yields a mean of SAMPLE PREPARATION AND ANALYSIS 1.1 per mil with individual values ranging from Sampling for the present study was based on a -+ 2.5 per mil for a total spread of nearly group of 22 sulphide-rich drill core samples pre- 7 per tomil.+ 4.4 Individual massive-ore lenses show a pared for trace-element analysis (Frater, 1978), for which purities, estimated from polished 1 —r mineral grain mounts, were generally better than 99%. This suite of samples was supplemented by drill core to provide additional representation of the several copper-ore lenses, zinc-rich hangingwall mineralization and lateral and vertical coverage of the stratigraphic interval enclosing the mineralization from the base of the Mineralized Horizon to the top of the Hanging-wall Chert. Sample localities are indicated on Figure 5. Sulphides were prepared for isotopic analysis by crushing and subsequent heavy liquid and magnetic separation, or by direct extraction from a polished section using a dental drill. Quantitative conversion to SO2 was achieved under vacuum, using an excess of Cu 0 at 900 °C, following the method of Robinson & Kusakabe (1975). Sulphur-isotope ratios were determined on a Micromass 602D mass spectrometer and are reported as mil values, through direct comparison with Canyon Diablo troilite or other sub-standards. Eleven samples were determined in duplicate and the method yielded an average reproducibility of 0.2 per mil. 34
INTRUSIVE ROCKS
2
f g T ]
Dolerite
IaT-M
Dac,le
MINERALIZATION Massive sulphide magnetite lens
VOLCANOCLASTIC ROCKS
6 3 4 S C d t
p e r
—
j '• •' •'' •' j Mineralized horizon j Chlorite schist
| V y J / j Footwall lapilli tuff
E 81
RESULTS The results of 70 sulphur-isotope analyses are reported in this paper (see Table I and Figs 5 and
6).
Fig. 4. Geological cross-section showing the distribution of massive ore lenses in the central portion of the Main Mineralized Zone, Golden Grove. Mine section 18400 metres North.
'
SULPHUR
ISOTOPES A N D ORE
GENESIS
425
TABLE I
Sulphur-isotope data for the Golden Grove massive ore lenses and host rocks
Lens. Central A
South A
Central
B
South B
Samp1e Number
R.L.
39107 391H 39116 53413 61622 61623 71456 80339 80340A 80340B
-200 -400 -100 -300 -200 -200 -100 -200 -200 -200
39110 39115 53477 80324
-400 -100 -400 -400
39104 39112 53417 53674 80327 80331 80332
-200 -400 -300 -400 -500 -500 -200
39109
-300
80326
-300
Central 0
80336A 80336B 80337
-400 -400 -400
North
39101 39103 39113 39114 53677 71151 80345 80346
-400 -300 -300 -400 -200 -400 -300 -200
Central
C
Py = p y r i t e ,
Py
6 3< *S Cp + 1.6 +1.7
Sp H.5
+4.4
S t r a t igraph i c Unit
Sample Number
R.L.
Mineralized Horizon
39108 53415 53428 53466 53470 53643 63670A 6367OB 63670C 63670D 63670E 63671A 63671B 65410 70493 71462
-300 -100 -500 -100 -200 -100 -100 -100 -100 -100 -100 -100 -100 -400 -300 -100
39117 39118 39119 39120 39121 39122 39123 39124 63668A 63668B 63669 63672A 63672B 63672C 65259A 65259B
-300 -300 -200 -200 -300 -300 -200 -300 -400 -400 -400 -200 -200 -200 -300 -300
0.0 +1.4 +0.6 -0.3 +0.9 0.0 + 1.2 +1.9 +1.8 +0.8
+ 1.7
+2.4 + 1.5 +0.2
Hang i n g - w a 1 ] Ghert
+0.6 +0.8 +0.9
+ 1-3 + 1.6 -0.5 -0.6
-0.7 -1.4
-1 . 1
+ 1.9
63"S Py
(% e ) Po
Sp
+0.9 +0.7 +4.0 + 1 .5 +2. • 5 + 1 .7 , + 1 •. 3 - 2 .• 3 -1 ..6 - 2 . .5 + 1 •. 5 + 1..6 +0. .8 + 1 .0 . +0.2 +3-.6 + 1 .3 + 1 ,5. +0.7 -0.1 +0.8 + 1.1 • +2.3 +4. 3 +0. .8 + 1.,0 + 1 .8 . +2. .2 + 1 .9 + 1 •. 7 +2.2 + 1.9
+ 1.6 + 1.1 + 1.4 + 1.8 +2.7 +0.7 +2.0
Cp = e h a l c o p y r i t e ,
Po = p y r r h o t i t e ,
more limited range in values (approximately 2 per mil). A lens (mean + 1.3 per mil, range - 0 . 3 to + 4.4; one anomalously 'heavy' value), B lens ( + 2.0, +0.2 to +2.4), Northern lens (+1.7, + 0.7 to +2.7) and C lens (+1.6, one value) all show similar distributions. The magnetite-rich D lens, the highest lens stratigraphically in the copper mineralization, contains relatively 'light' sulphur (mean - 0.9 per mil, range - 1.4 to - 0 . 5 per mil). Sulphides (principally pyrite) from the Mineralized Horizon yield the greatest range in S34S values ( - 2 . 5 to +4.0 per mil). Three pyrite samples from the base of the Mineralized Horizon and to the north of the ore zone give the lightest S34S values (range - 1.6 to - 2 . 5 per mil) and suggest a bi-modal distribution on Figure 6. These three samples represent individual (1 mm thick) pyrite microbands from a 2 cm section of drill core. However, a fourth pyrite microband yields a <534S value of +1.3 per mil (sample 63670A) and pyrite from a discordant quartz vein from the same section of core (sample 63670E) + 1.5 per mil—values which are consistent with those from the massive-ore lenses. Pyrite and sphalerite from the Hanging-wall
Sp =
sphalerite.
Chert (mean +1.6 per mil, range - 0 . 1 to +4.3 per mil) yield a similar sulphur-isotope distribution to the ore lenses. With the exception of massive-ore lens D as noted above, differences are not noted in isotopic distribution amongst the various ore lenses or within individual ore lenses intersected at different R.L.'s (see Fig. 5). Rare 'heavy' values in pyrite occur in the Central A lens and stratigraphically above Southern A lens at - 1 0 0 m R.L. and in the Hanging-wall Chert to the south of the main ore zone at - 3 0 0 m R.L. A single sphalerite sample yields 'heavy' sulphur from vein mineralization close to the top of the Mineralized Horizon on the - 5 0 0 m R.L., also to the south. The limited data available to date fail to show evidence of isotopic equilibrium amongst coexisting sulphide minerals. Of the four coexisting mineral pairs sampled from A and D lenses listed in Table I, only one shows the correct direction of isotopic fractionation (sample 53477). This sample, however, yields geologically unreasonable temperatures applying the suggested fractionation factors of Ohmoto & Rye (1979). Even using overall averages for the
P. K. SECCOMBE & K. M. FRATER
426
-i
Py-1 3,*1 5t -2 5,-2 3, -1 6 \ -100m
RL
—
R L
Fvffl 2 0 0
m
SCJLTHERN A
CENTRAL
Sp+1-1,+1-9,+2 2
Cp-0-3 ^ / Py 0 0.+0-9J v v Py+1-5 Po-09
3 0 0
SOUTHERNS,
—
m
R L
p
Npo.r^Cp.l
l
Py+1 8 Cp*1 6
Py20
Py1 1
CpOO / Sp+4-0 Sp-11,-0-6,-0-51 (-500m R.L) Ry.r^-0 7 ^ ^ Cp+1-5- 4 0 0 m RL PO+V3'
Y
CENTRAL
NORTHERN
A
CENTRAL
SOUTHERN,
W
r
Sp«-2 3
/T
v
1
Py+0 6.-0 6 \
Cp*0-6"| cnnmRI 1 / io - P y + 1 0 / ~ P y + 0 ^
D
D
A
V rCp*12,-1-7_ lPy+0-8,+1 -8 I I
p^j
Hanging-wall chert
[
] Mineralized horizon
V MJ
I
CP+1-7
N
I
CENTRAL
A
V IPy*08, J10,
v
L.
I
Massive sulphide—magnetite lens
Footwall lapilli tuff
Fig. 5.
S S (per mil) values and sample locations for sulphides from massive ore lenses, Mineralized Horizon 34
cdt
and Hanging-wall Chert, Golden Grove. Cp = chalcopyrite; Po = pyrrhotite; Py = pyrite; Sp = sphalerite.
various sulphide species, a lack of consistency in DISCUSSION the direction and degree of inter-mineral isotopic The sulphur-isotope distribution for sulphides fractionation is apparent. Such isotopic disequi- from the massive-ore lenses and host rocks at librium is a common characteristic of sea-floor Golden Grove compares closely with reported exhalative ore deposits (Ayres ET AL., 1979; values for strata-bound massive sulphide deposits Solomon ET AL., 1969). in Archaean rocks elsewhere (Coomer & At Golden Grove, diagenetic or metamorphic Schwarcz, 1974; Sangster, 1976; Seccombe, effects on the sulphide mineral assemblage may 1977). be a factor in altering the sulphur-isotope distriThat the present sulphur isotopic distribution bution amongst different sulphide species. Pyrite at Golden Grove must correspond closely with is commonly in obvious textural disequilibrium the pre-metamorphic distribution may be argued with the more ductile sulphides and displays later from the dispersion of values (particularly pyrite) overgrowths. Secondary overgrowth pyrite, how- within ore lenses and rock units, differences ever, appears to be isotopically identical to early amongst ore lenses and evidence of isotopic dispyrite. Two generations of sphalerite are recog- equilibria. nized during sulphide recrystallization and Using typical compositional data and properpyrrhotite is commonly replaced by pyrite and ties of a Kuroko-type hydrothermal ore fluid magnetite. Reactions such as the formation of nearing the sea floor (Solomon & Walshe, 1979) pyrite at the expense of pyrrhotite may give rise and the presumed primary mineralogy in the to isotopic fractionation (Ohmoto & Rye, 1979) iron-sulphur-oxygen system for the Golden whereas simple recrystallization under green- Grove mineralization, an estimate can be made of schist metamorphism is unlikely to alter the pre- the relationship between the sulphur-isotope existing isotopic distribution. composition of the precipitated sulphide phases
S U L P H U R ISOTOPES AND ORE GENESIS -4
+2
0 Chalcopyrite • Pyrite A Pyrrhotite 0 Sphalerite
+4
+6
A A Lens
O O • AO
omon
ODBDD
•
O
B-C Lenses
AD O o n ADO
D Lens
0
00 ••
• o•
Northern Lens
•
Mineralized horizon • • •
A • ADD
• • • A
•
ADDD
•
DO
0•
Hanging wall Mineralization
0
•
DO
oono • ••• -4
-2
_L 0
_L
+2
+4
+6
834SCDT (PER MIL) Fig. 6.
Sulphur-isotope distribution (S 34 S CDT per mil) for sulphides from massive ore lenses, Mineralized Horizon and Hanging-wall Chert, Golden Grove.
and that of the total sulphur in the ore fluid. For the copper mineralization, pyrite and pyrrhotite occur mainly at the stratigraphic base and become less abundant higher in the pile, where magnetite becomes dominant. Toward the top of the Main Zone, textures in magnetite indicate that it is pseudomorphous after an original hematite-goethite assemblage. For low-salinity, mildly acid fluids, a temperature of 350 °C is required to transport sufficient Cu in solution for ore deposition (Walshe, pers. comm.). At this temperature, 5 34 S values for co-existing pyritepyrrhotite assemblages should approximate the & 4 S l s value of the ore fluid (Ohmoto & Rye, 1979). Thus a value near zero per mil is indicated for sulphur in the ore fluid responsible for the bulk of copper precipitation. Toward the top of the Main Zone where magnetite (or hematite) becomes the dominant primary iron-bearing phase (i.e. Central D lens) a shift toward more negative values for sulphides within this ore
427
lens may be explained readily by a higher oxidation state of the ore fluid (higher L S O ^ E ^ S in solution). From the data of Ohmoto & Rye (1979), however, it can be shown that equilibrium could not have been achieved between the aqueous sulphur species in the hydrothermal fluid, even at T = 350 °C, as a shift of nearly - 20 per mil would be required in 5 34 S values for sulphide precipitation in the hematite-magnetite field. With the return to a pyrite-dominated assemblage in the Hanging-wall Chert, sulphide 5 34 S values are near zero per mil to slightly positive, reflecting again a more reducing character in the ore fluid. The reason for such a fluctuation in oxidation state of the ore fluid at a particular stratigraphic level in the deposit is not clear. An active tectonic environment during ore deposition is inferred from a palaeogeographic reconstruction and the dilution by volcanoclastic material and spatial separation of mineralization (Frater, 1978). Tectonic effects on the marine basin leading to increased circulation and penetration of seawater and a temporary change in the redox conditions of the hydrothermal fluid (increase in f02) may be one explanation for the observed trend. A mean value for the 6 34 S of the ore fluid of near zero per mil suggests a sulphur source of magmatic origin, either as a component of a juvenile hydrothermal fluid or in the more likely process of ore formation involving deep circulation of seawater, through leaching of silicate and sulphide minerals (e.g. Solomon, 1976; Ohmoto & Rye, 1979). However, a sulphur contribution from Archaean seawater cannot be ruled out, particularly if the mean value for dissolved sulphur species in Archaean seawater was close to the meteoritic value (see Lambert, 1978). Isotopic disequilibrium noted amongst coexisting sulphide minerals implies quench precipitation of sulphides and lack of subsequent isotopic annealing within the volcanic-sedimentary pile which might have been expected as a result of succeeding hydrothermal activity and growth of the exhalative vent (Solomon & Walshe, 1979) or later metamorphism. The recognition of metal zoning within individual sulphide-rich pods of massive mineralization, whereby copper is concentrated at the base of pods and decreases in content upward and zinc occurs preferentially above the highest copper concentrations, suggests that much of the massive mineralization has formed on the sea floor. The significant spatial separation of ore lenses and the copper and zinc mineralization may mean that fumarolic activity was intermittent. Thus a prolonged annealing episode within the pile may not have been achieved.
P. K. SECCOMBE & K. M. FRATER
428
CONCLUSION
ACKNOWLEDGMENTS
Sulphur isotopic evidence supports the contention that copper-zinc mineralization at Golden Grove formed from sea-floor hydrothermal exhalations. Thus the formation and nature of the deposit may be compared with strata-bound copper-zinc deposits in acid volcanic rocks elsewhere, particularly those of Archaean age in Canada. Further sulphur isotopic measurements at Golden Grove are being undertaken to clarify stratigraphic variations within the deposit and in individual ore lenses and the contribution, if any, of bacteriogenic sulphur.
The authors thank the partners in the Golden Grove Joint Venture, Electrolytic Zinc Company of Australia, Amax Exploration and Esso Minerals, for release of data and support of this project. We are grateful to Mr T. H. Donnelly of the Baas Becking Geobiological Laboratory, Canberra and Dr R. Summons, Research School of Biological Sciences, Australian National University, Canberra, for use of the mass spectrometer. We also thank Mr W. Crebert for drafting the diagrams, Mrs J. Walker and Ms K. Jackson for photography and Mrs D. Winterborn for typing the manuscript. This study is funded by the Australian Research Grants Committee.
REFERENCES AYRES, D . E . , B U R N S , M . S . , & SMITH, J . W . ,
1979:
Sulphur-isotope study of the massive sulphide orebody at Woodlawn, New South Wales. J. geol.
Anal. Chern., 47, 1 1 7 9 - 1 1 8 1 .
Soc. Aust., 26, 197-201. COOMER,
P.
G.,
& SCHWARCZ,
H.
P.,
1974:
Sulfur
isotopic study of a massive, volcanogenic Archaean sulfide ore deposit. Geol. Soc. Am., Ann. Mtg.,
Abstracts with Programs, 693-694. FRATER, K. ML, 1978: The Golden Grove Copper-zinc Deposit, Western Australia—an Archaean, Exhalative, Volcanogenic Occurrence. Ph.D. Thesis, Univ. of Newcastle [unpublished]. LAMBERT, I . B., 1 9 7 8 : Sulphur-isotope investigations of
Archaean mineralization and some implications concerning geobiochemical evolution; in Glover, J. E., & Groves, D. I. (Eds) Archaean Cherty Metasediments: Their Sedimentolo-y, Micropalaeontology, Biogeochemistry and Significance
to Mineralization. Pubis geol. Dep. & Extension Service, Univ. West. Aust., 2, 45-56. LAMBERT, I . B . , & SATO, T . ,
1974: T h e K u r o k o
and
associated ore deposits of Japan: a review of their
features and metallogenesis. Econ.
Geol.,
69,
1215-1236. OHMOTO, H., & RYE, R. O., 1979: Isotopes of sulfur and carbon; in Barnes, H. L. (Ed.) Geochemistry
of Hydrothermal Ore Deposits (2nd Ed.), 509-567 Wiley, N.Y.
Quantitative preparation of sulfur dioxide, for 3 4 S / 3 2 S analyses, from sulfides by combustion with cuprous oxide.
ROBINSON, B . W . , & KUSAKABE, M . , 1 9 7 5 :
SANGSTER, D. F., 1976: Sulphur and lead isotopes in strata-bound deposits; in Wolf, K. H. (Ed.) Hand-
book
of
Strata-bound
and
Stratiform
Ore
Deposits, 2, 219-266. Elsevier, Amsterdam. D. F., & S C O T T , S . D., 1976: Precambrian, strata-bound, massive Cu-Zn-Pb sulfide ores of North America; in Wolf, K. H. (Ed.) Handbook
SANGSTER,
of Strata-bound and Stratiform Ore Deposits, 6, 129-222. Elsevier, Amsterdam. SECCOMBE, P. K., 1977: Sulphur isotope and trace metal composition of stratiform sulphides as an ore guide
in the Canadian Shield. J. geochem. Explor., 8, 117-137. SOLOMON, M., 1976: "Volcanic" massive sulphide deposits and their host rocks—a review and an explanation; in Wolf, K. H. (Ed.) Handbook of
Strata-bound and Stratiform
Ore Deposits, 6,
21-54. Elsevier, Amsterdam. SOLOMON, M . , & WALSHE, J . L . , 197.9: T h e f o r m a t i o n of
massive sulfide deposits on the sea floor. Econ.
Geol., 74, 797-813. SOLOMON, M . , RAFTER, T . A . , & J E N S E N , M . L . ,
1969:
Isotope studies on the Rosebery, Mount Farrell and Mount Lyell ores, Tasmania. Mineralium
Deposit a, 4, 1 7 2 - 1 9 9 .
STRUCTURE AND IRON-ORE DEPOSITION IN THE ARCHAEAN KOOLYANOBBING GREENSTONE BELT, WESTERN AUSTRALIA Adrian C. Griffin
School of Earth Sciences, University of Melbourne, Parkville, Australia 3052 ABSTRACT Banded iron formations, of Archaean age, intercalated within a sequence of intermediate to ultramafic lavas with minor dykes, sills and pelitic rocks, outcrop in the Koolyanobbing Range. The southwest margin of the Koolyanobbing greenstone belt is truncated by a major fault juxtaposing the metavolcanics with gneissic granites and mylonites. The northwest margin of the belt is poorly defined, consisting of a zone of lit-par-lit granitic intrusions and pendant structures before passing into granitic country rock. Three main deformation events are exhibited in the Koolyanobbing greenstone belt. D] is characterized by upright isoclinal folding and widespread axial-planar foliations developed during prograde metamorphism. The D 2 event probably accompanied the culmination of the main metamorphism, F 2 folds being asymmetric and associated with axial-surface faulting, brecciation and contemporaneous granite intrusion. The D3 deformation is characterized by open folding about horizontal axes. Three distinct types of iron-ore deposit occur within the banded iron formation host. The oldest type is associated with brecciated F2 fold cores and hydrothermally leached, banded iron formation. After leaching, the iron-rich leaching solutions cooled and began to precipitate hematite ''muds". Coarse-grained specular hematite and/or elongate magnetite nucleated on the hematite substrate and grew from the solutions. Discordant specularite veins resulted from the infilling of fractures in the banded iron formation and adjacent metatholeiites. The timing of hydrothermal deposition at specific localities, with respect to the D 2 event, can be inferred by the amount of deformation within the given deposit. Younger supergene mineralization envelopes the hydrothermal deposits and also occurs in deposits of entirely supergene origin. The supergene ores contain the two remaining genetic ore types. The first type developed as a response to exposure of unstable iron-bearing minerals, e.g. carbonates and sulphides, to weathering, and the second was structurally enhanced with oxidation of the host BIF in zones of intense F 3 folding and brecciation. Both types of supergene deposit are dominantly hematite, martite, goethite and limonite.
INTRODUCTION In this paper it is p r o p o s e d to discuss the stratigraphy and the structure observed within the Koolyanobbing greenstone belt, the types of ironore deposit developed, a n d the influence of d e f o r mation on localizing sites of mineralization. T h e Koolyanobbing greenstone belt lies approximately 360 km east of P e r t h (Fig. 1) a n d outcrops as a sequence of d e f o r m e d amphibolites and metamorphosed u l t r a m a f i c rocks with intercalated b a n d e d iron f o r m a t i o n s (BIF) and minor psammitic a n d pelitic assemblages (Fig. 2). Mineralogy indicates that grades of amphibolitefacies m e t a m o r p h i s m were o b t a i n e d , with subordinate greenschist-facies assemblages suggesting m e t a m o r p h i s m peaked over a geologically short time (Blais et al., 1977). T h e amphibolites show tholeiitic c o m p o s i t i o n s whereas the ultramafics were derived f r o m peridotitic komatiites
Spec. Pubis geol. Soc. Aust., 7 (1981)
and minor dunites. Rare metapelitic assemblages of andalusite and fuchsite indicate fairly lowpressure m e t a m o r p h i s m (Binns et al., 1976). T h e sequence contains less acid and intermediate lithologies t h a n most other Yilgarn Block greenstone belts (Horwitz & Sofoulis, 1965; Williams, 1975; Gee, 1979) and contains the only economic iron ore deposits. T h e main tectonic events affecting the Yilgarn Block (Archibald et al., 1978) are recognized in the Koolyanobbing belt, with tectonism commencing a b o u t 2 7 0 0 m . y . (Turek & C o m p s t o n , 1971) a n d culminating a b o u t 2 6 0 0 m . y . (Arriens, 1971). This event is believed to be widespread, covering most of the Eastern Goldfields Province (Archibald et al., 1978) a n d is c o n t e m p o r a n e o u s with structurally controlled, h y d r o t h e r m a l , iron-ore deposition. Supergene enrichment, although o f t e n associated with specific structural d o m a i n s ,
A. C. GRIFFIN
430
TABLE I +
: : jackson{4 : A^\ l\3° S + + fj
to +
+
Major rock types and mineral assemblages
MAJOR PHASES +
+ ^s, +
KOOLiJ^OBBING^
+
ACCESSORIES
APPROXIMATE TOTAL STRATIGRAPHIC THICKNESS 6,000 m
Amphibolite Hornblende Ep idote (Metamorphosed Act i noli te CIinozoisite Tholei i tes) P1ag ioclase Quartz (An^s-eo) + + + Biot ite Chlorite Garnet IImen i te S&U*HERt(\ CROSS Sphene Chlorite Dolomi te 800 i Tremoli te Calcite A1terat ion Greenstones + + ^ & + Antigorite Magnes i teProducts CIinozoisite Granite Talc Magnet i te s Chrysot ile ri t |p> + 0 25 50 75 1 1 i i Metamorphosed Talc Hornblende Kilometres i Tuffs Chlorite Quartz Ant igori te Plagioclase Act i noli te Fig. 1. Regional geology of the Eastern Goldfields Tremoli te Province, Yilgarn Block, Western Australia, Andalus i te after Binns etal. (1976). Pyri te 50 i Fuchs i te Plagioclase Biotite Quartz is a comparatively recent feature and is discussed only briefly in this paper. Banded Iron Magnet i te Hemat i te Gruneri te Garnet The styles of deformation, strong foliation Format i on Quartz Arsenopyr i te Cummi ngton i te development in some lithologies, and metaTalc morphic textures (relict greenschist-facies assemChlorite Dolomi te blages, showing cleavage, preserved in amphiboPyri te lites) suggest that the first tectonic event (Dj) occurred during prograde metamorphism. The second deformation (D2) did not produce wide- Sediments spread dynamic recrystallization and foliation Both chemical and minor detrital sediments development was limited. Some D2 anticlinal crop out at Koolyanobbing. The detrital sedicores were intruded by granite thought to be conments include quartzite (often containing temporaneous with the folding and hydrothermal fuchsite, biotite and altered feldspars) and pelite. activity which commenced the mineralization of The sediments of dominantly chemical origin are the main ore bodies. The granite intrudes the quartzite (being entirely quartz, derived by areas of highest metamorphic grade and probably recrystallization of chert) and banded iron forrepresents the base of the greenstone belt. mations, which are considered to be a single depositional facies. MAJOR ROCK TYPES The greenstone belt can be divided broadly into Banded Iron Formation Facies. Three main a greenstone sequence and a sedimentary iron formation units can be traced for the sequence. A chronological order of rock deposi- length of the Koolyanobbing belt and for contion has not been determined because of a lack of venience are called the Lower, Middle and diagnostic facing criteria. Within the sedimentary Upper BIFs. They are well layered with "chert" sequence banded iron formations are the host for (now entirely recrystallized to quartz), and carbonate laminations having thicknesses of 1 to iron-ore mineralization. 20mm and averaging 2 to 3 mm. Iron-rich The Greenstones laminations are; generally thinner and consist of The greenstone sequence is dominated by basic magnetite + amphiboles (grunerite or cumamphibolite, thin waterlaid basic tuff and ultra- mingtonite) ± pyrite with accessory talc, mafic rocks. The amphibolite is derived from chlorite, arsenopyrite, chalcopyrite, hematite pillow lavas, flows and intrusives, whereas the and garnet. Applicability of facies models (e.g. ultramafic rocks show fabrics characteristic of James, 1954) to the Koolyanobbing BIFs is peridotitic komatiite flows (Nesbitt, 1971), and poor as the mineralogy changes rapidly along less commonly, pyroxenites and dunites.' The strike. observed mineral assemblages are summarized in Generally, the iron-rich laminations of the Table I. BIFs contain grunerite as the main amphibole, +
+
BULLFINCfrfA \\
\A
\
\
+
+
* 31 S~
uM
73 O
z o
7* tn ^ O O
r >
z o
DO
Dd O O 70 m
tn
*
IRON ORE DEPOSIT
ULTRAMAFIC
MYLONITIZED
BANDED IRON FORMATION
GRANITE
GRANITE
PELITIC
PORPHYRITIC ACID D Y K E S
BASIC
UN MINERALIZED
MAJOR
z
ROCKS
en H O
z
m w m r H
SEDIMENTS
VOLCANICS,TUFFS
KILOMETRES
AND AMPHIBOLITES
F 2 FOLD CLOSURE Fig. 2.
Simplified geological m a p of the K o o l y a n o b b i n g greenstone belt, Western Australia.
u>
432
A. C. GRIFFIN TABLE II
Composition of amphibole in BIFs Si0 2
T i O2
AI2O3
FeO
MnO
MgO
CaO
Na 2 0
K20
Cr203
NiO
BaO
TOTAL
1.
54.59
0.00
0.14
26.73
1.10
15.48
0.43
0.03
0.00
0.04
0.00
0.09
98.51
2.
51.21
0.06
0.09
40.93
0.68
5-99
0.37
0.01
0.00
0.06
0.00
0.00
99.41
Amphiboles from the Middle BIF are iron rich at the north-west end of the greenstone belt and in the vicinity of iron orebodies (2) but more magnesium-rich towards the south-east extremity of the belt (l) where the BIFs contain abundant volcanogenic contamination.
but toward the southeast end of the belt the MgO/FejoT of the amphiboles increases (Table II). The width of the laminations in this region increases to an average of 5 mm, and the greenstone sequence lying between the Middle and Upper BIFs grades laterally into layered cummingtonite and quartz rocks (Fig. 3). The cummingtonite/ quartz rocks are similar in appearance to BIFs but contain iron-rich laminations consisting of Fe-rich cummingtonite, and magnetite, commonly up to 30cm thick. Cummingtonite rocks are also observed on Lake Seabrook Peninsula (Fig. 2), and occur as discordant lenses of intrusive origin, and consequently the cummingtonite within the sedimentary rocks is considered to result from volcanogenic contamination of banded iron formation. Further evidence for contemporaneous BIF deposition and volcanism can be found at Dowds Hill quarry and " A " Deposit, where talc + chlorite + quartz + amphibole rocks occur within the Middle BIF. These horizons are mostly tholeiitic in composition and are interpreted as lava flows. Carbonate-rich iron formation is localized within the middle BIF and occurs near the footwall at Dowds Hill and at " A " Deposit. The carbonate is dominantly dolomite and pistomesite with minor calcite, sideroplesite and ankerite. The footwall of both deposits also contains pyrite-rich lenses with accessory arsenopyrite and chalcopyrite. The lenses are discontinuous along strike and some contain serpentine, talc and graphite (Ellis, 1958). Detrital Quartzite. Detrital quartzite contains plagioclase, K-feldspar, sericite, fuchsite, garnet, hornblende and biotite, unlike the chemically precipitated quartzite which contains only quartz. The grains average 0.25 mm in diameter with a few grains attaining 2mm. Foliation development is poor, with only micas and amphiboles aligning in the regional foliation (SO orientation, and quartz remaining sub-equigranular. The rocks were probably derived by rapid erosion of sialic basement near the basin margin.
Pelite. Pelitic assemblages contain knotted andalusite porphyroblasts, up to 2 cm across, with envelopes of fuchsite and/or green biotite. The andalusite is poikiloblastic with inclusions indicating contemporaneous andalusite growth and Si cleavage development. Pelitic rocks outcrop as lenses usually less than 3 m thick and with strike lengths less than 200 m. STRUCTURE AND METAMORPHISM Facing evidence within the Koolyanobbing belt viz. pillows, spinifex-textured flows and graded bedding in tuff, is sparse and often inconsistent over small distances, particularly in areas of early isoclinal folding. In the North Range (Fig. 2), the Middle and Lower BIFs converge but there is evidence for neither fold closure nor stratigraphic continuity. The regional (Si) foliation is close to layer-parallel is most localities, apart from those showing cleavage refraction. The sense of bedding/Sj intersections is the same throughout the Koolyanobbing greenstone belt suggesting that the belt is a single limb of a regional isocline, truncated on the southwest side by a major fault (Fig. 2). The First Deformation (Dj) Unlike many other greenstone belts within the Eastern Goldfields, the Koolyanobbing belt shows no evidence of early nappe tectonics (Archibald et al. 1978) but has an early deformation event characterized by folds with upright axial surfaces. The folds are isoclinal, have an axial-surface foliation, and fold axes which trend NNW (parallel to the regional strike). Cleavage/ bedding intersections suggest that the major fold axes were probably steeply plunging. However, minor folds plunge from vertical through to horizontal, suggesting that strain was inhomogeneous at least on a local scale. The existence of greenschist-facies assemblages as Si cleavage-forming elements in tuffaceous lithologies and amphibolite-facies assemblages in the amphibolites, suggests that the deformation Di occurred during the prograde metamorphic event. Foliations are most strongly developed in y
IRON ORE IN KOOLYANOBBING GREENSTONE BELT
M
00
3D
3.
Diagrammatic geological cross section showing the stratigraphic relationship of Middle and Upper BIFs, facies relationships of basic tuffs and amphibolites to cummingtonite/chert rock and the location of orebodies.
433
fine-grained clastic lithologies such as basic tuff, in which chlorite grows with (001) parallel to schistosity whereas antigorite and amphiboles tend to form flat, sheaf-like aggregates that define a schistosity plane. Quartz and plagioclase are usually interstitial grains between amphiboles and chlorite, and tend to be fine grained (less than 0.1 mm) and maintain random dimensional orientations. A less penetrative schistosity is developed in amphibolites. Amphibole grains grow with long axes in the schistosity plane whereas the quartz and plagioclase show only slight dimensional preferred orientation. Pillow lavas and komatiitic flows rarely develop any foliation and consequently preserve many igneous textures. Using the criteria of Binns et al. (1976), the first deformation appears to be transitional from static to dynamic tectonism. Penetrative foliations do not occur in all rocks and the metamorphic assemblages forming the foliation belong to both greenschist and amphibolite facies. Although some rock types show only greenschist-facies assemblages, most rocks with penetrative foliations show greenschist minerals as a relict phase being replaced by higher-grade, amphibolite-facies assemblages, thus indicating the Sj surface was formed during prograde metamorphism. Because the metamorphic assemblages are compositionally controlled in some rock types, the variation of metamorphic grade within the sequence cannot be readily shown. If the grain size of the greenstone sequence and the major BIFs is considered, coarsening indicates the metamorphic grades were probably highest along the northeastern margin of the North Range and near Lake Seabrook, in the South Range (Fig. 2). The model proposed by Binns et al. (1976) suggests that dynamic domains are usually associated with amphibolite-facies grades of metamorphism and static domains with greenschist-facies grades, and therefore the Koolyanobbing belt represents a transition in both metamorphic grade and structural style during the prograde metamorphic event. The Second Deformation (D2) Most folds observed throughout the Koolyanobbing greenstone belt deform the S\ foliation. Development of new foliations during the D2 event appears to be independent of lithological controls and confined to areas of high strain. New foliations (S2) crenulate the pre-existing Sj surface and produce a crenulation lineation coaxial to the F 2 fold axes. Recrystallization of amphiboles in domains of less deformation results in mineral elongation lineations coaxial with F 2 . Few areas show the development of S2
A. C. GRIFFIN
434
(a)
(b)
Fig. 4. (a) Schmidt net of poles to bedding in the Dowds Hill orebody. The distribution of poles is a result of F folding. (b) Schmidt net showing F fold axes in the Dowds Hill orebody.
2
2
foliations in the absence of Si and those which do generally have poorly developed non-penetrative foliations defined by moderate dimensional preferred orientation of amphiboles. The F folds are asymmetric 'S' shaped folds with an average plunge of 50° and axial trends varying from northeast to southeast (Fig. 4). The axial surfaces dip moderately with directions of maximum dip varying from north through to southeast. Much of the orientation variation resulted from progressive strain in which most F folds commenced as buckles, with axial surfaces perpendicular to the layering. As strain increased, the fold amplitudes increased and axial surfaces rotated, and new buckles commenced in adjacent areas. The folds preserved after the deformation ceased show a range of development and strain, and consequently a range of orientations of their structural elements. The grade of metamorphism increases toward Lake Seabrook resulting in coarser grain size in the banded iron formations and development of garnet from chlorite in the amphibolites. These rocks probably represent deeper crustal levels which have been subjected to metamorphism by a heat source within the granite below the greenstone belt (Archibald et al., 1978). Five major F closures occur within the Koolyanobbing Range. Three have extensively brecciated cores, consisting of angular fragments of BIF ranging from sand-sized to metres across. The breccias are pipe-like bodies which resulted from brittle failure of competent horizons 2
2
2
following tight folding. The breccia matrix consists of specular hematite with minor pyrite, quartz and magnetite. This type of mineralization forms the core of orebodies at Dowds Hill, "C" Deposit and " D " Deposit. The two remaining major F structures are not brecciated or mineralized (Fig. 2). Block faulting along the belt is associated with F folding and where exposed (Dowds Hill and " C " Deposit) is coplanar with the axial surface of major F structures. The remaining faults are not exposed but inferred from airborne magnetics. These faults generally cross the Middle BIF at locations where its true thickness is small and are also assumed to be contemporaneous with D . Although specular hematite mineralization in the F fold cores is associated with brecciation, many occurrences are deformed, suggesting that ore emplacement had commenced before culmination of the D event. Undeformed specular hematite further suggests that mineralization continued beyond the time span of D which can also be verified by the specularite veins crosscutting small F folds. Brecciation contemporaneous with hydrothermal hematite precipitation, with clasts of leached BIF dropping to the bottom of the cavities in which hematite was forming, explains the existence of subhorizontal breccia sheets which formed in the ore at " C " Deposit, and verifies the contemporaneous tectonism which deformed the ores. As most breccia sheets at "C" 2
2
2
2
2
2
2
2
IRON ORE IN KOOLYANOBBING GREENSTONE BELT
Deposit must have been close to the culmination of the D2 event they have suffered little largescale deformation since that time. In contrast, the Dowds Hill deposit shows only small amounts of undeformed specularite and no horizontal breccias, indicating that most of the ore was in place well before the culmination of D 2 while deformation was dominantly ductile rather than brittle. The last stages of mineralization persisted until well after D2 with small veins of hematite crosscutting F2 folds. This indicates that the hydrothermal event at Dowds Hill spanned a greater time than at " C " Deposit. " D " Deposit contains mainly post-tectonic undeformed specularite representing the last stages of hydrothermal enrichment. This is further verified by the lower temperature of fluid inclusions associated with gangue minerals of this deposit (Sullivan, 1973). The Third Deformation (D3) Deformation post-dating the D 2 event, folds the F2 fold axes. The F 3 folds are generally smallamplitude open folds with subhorizontal axial surfaces and subhorizontal fold axes trending 210°. The folds are mostly symmetrical, but at "F" Deposit the axial surfaces vary from subhorizontal to subvertical (northeasterly dipping) and the folds are asymmetric. In localities exhibiting steep S3 surfaces, brecciation in F 3 fold cores is common. D3 breccias are not associated with specular hematite mineralization, but do result in greater permeability to meteoric water and consequently enhance supergene enrichment. THE IRON OREBODIES Koolyanobbing Range contains five ore deposits with reserves totalling 1.2 x 108 tonnes. Three are structurally controlled, resulting in hydrothermal mineralization commencing during the D2 event. The hydrothermal ore forms in the core of F2 structures and is now surrounded by an envelope of supergene ore. The remaining deposits contain ore of supergene origin only. Dowds Hill Type Orebodies Dowds Hill, " C " Deposit and " D " Deposit contain hydrothermal ore and are characterized by intense D2 deformation and pipe-like breccias developed along the axes of major F 2 closures. The brecciated zones contain most of the hydrothermal mineralization. The extent of brecciation is unknown because of the irregular cross section of the breccia pipes and the effects of hydrothermal replacement around the pipes. It is however unlikely that any breccias exceeded 120 m across, which is the maximum width of structureless hydrothermal ore typical of the breccia
435
zones. BIF material remaining in or adjacent to the pipes was leached of iron and now appears as friable banded quartzite. The ore zones directly associated with such material are often layered and contain planar accumulations of leached BIF clasts, cryptocrystalline red hematite, and euhedral quartz. Specular hematite occurs above the red cryptocrystalline hematite, and is the main ore mineral, but is not layered itself. The ores are coarse grained and consist of specularite up to 30 cm across, euhedral quartz crystals up to 8 cm long, bladed magnetite up to 2 cm long and pyrite up to 2 cm across. BIF horizons adjacent to hydrothermal pipes, in particular those that were rich in carbonate, show extensive hydrothermal replacement. Replacement textures are finer grained than ore in the breccia pipes because replacement took place grain by grain, and textures reflect the grain size of the host BIF. BIF clasts within the breccia pipes are extensively leached of iron, as are some of the surrounding silicate-rich BIFs. The leaching is generally restricted to zones of intense deformation. The specularite forming the matrix of the breccia shows strain variations from intense to undeformed. Dowds Hill contains both deformed and undeformed specularite, whereas all " C " Deposit specularite is deformed and all " D " Deposit specularite undeformed. This suggests that " D " Deposit ore was emplaced near the end of the D2 event, " C " Deposit ore before the end of D 2 , and Dowds Hill mineralization probably commenced early in the D2 event and continued until after the culmination of D 2 . Heterogeneous strain cannot account for the deformation observed within the Dowds Hill specularite pipes as they contain undeformed ore irregularly distributed and commonly adjacent to deformed ore. All hydrothermal orebodies in Koolyanobbing Range are enveloped by martite, hematite, goethite, limonite ores. Such ores occur only in the oxidized zone and consequently do not extend to the depths of hydrothermal mineralization. Supergene Enrichment
Orebodies
"A" Deposit. " A " Deposit is the simplest type of orebody developed by supergene enrichment. It is a planar orebody dipping steeply northeast and is the only orebody in which the entire width of the Middle BIF has been mineralized. The ore contains goethite, limonite, hematite, pyrite, siderite, magnetite and minor chalcopyrite. The ore along the footwall is gossanous being derived from oxidation of pyrite, and the remainder of the orebody is hematite-rich with minor siderite, suggesting a carbonate/magnetite BIF host. The nature of the fresh BIF was confirmed by Ellis (1958).
436
A. C. GRIFFIN
The southeastern end of the orebody is bounded by a fault bringing much thicker BIF into juxtaposition with the " A " Deposit host horizon. The thicker BIF shows no evidence of anomalous sulphide, carbonate or magnetite concentrations.* "F" Deposit. " F " Deposit contains tight, overturned F3 folds with extensively brecciated fold cores. Mineralization occurs only in the zone of intense F 3 folding and shows many features in common with " A " Deposit and the zones of supergene enrichment in deposits of the Dowds Hill type. Much of the ore is gossanous, being limonitic material derived from both pyrite and magnetite. The ores contain more silica than those of other supergene zones in the Koolyanobbing area. Si0 2 is as high as 20wt% of which 12% commonly appears as free quartz and the remainder is in limonite. ORE PARAGENESIS Ores containing martite, magnetite, goethite and limonite form more than 80% of ore reserves in the Koolyanobbing iron ore deposits. This type of ore occurs as envelopes around orebodies of the Dowds Hill type and is the only type at " A " Deposit and " F " Deposit. The ore is a surficial deposit rarely extending more than 50 m below the water table and is typical of ores developed by supergene enrichment (Dorr, 1964). It is developed by leaching of silica by meteoric water combined with hydration and oxidation of iron-rich BIF components. Although there is little direct evidence, the lack of -consolidated scree ore deposits (with the exception of " A " Deposit), suggests that most of the supergene enrichment is associated with relatively recent weathering cycles. The existence of a major fault at the southeast end of " A " Deposit, in addition to the contrast in BIF mineralogy across the fault, suggests that carbonate, magnetite and sulphide BIFs were brought into juxtaposition with BIFs containing less iron and more stable mineralogy. Pyrite and carbonate decomposed rapidly under atmospheric conditions producing large concentrations of limonite and greater permeability of the host BIF, permitting oxidation and hydration of magnetite at depth. A similar situation was enhanced by brecciated F 3 fold closures at " F " Deposit. The brecciation increased permeability of the BIF, permitting deep oxidation, and hence little ore is developed outside the zone of intense F 3 folding. Ore deposits of the Dowds Hill type are distinguished from the supergene orebodies by containing specularite mineralization (associated with
brecciation of F 2 fold cores), the depth of which is independent of the oxidation depth of the host horizon. Because the breccias are contemporaneous with D2 and the specular hematite forming the breccia matrix is itself deformed, the mineralization must have commenced well before the culmination of D 2 . Undeformed specularite occurs at Dowds Hill and " D " Deposit, some of which transgresses F 2 fold closures, indicating that mineralization continued beyond the culmination of the D 2 event. Fluid-inclusion studies on quartz within the Dowds Hill and " D " Deposit specularite suggest average temperatures of formation of about 230 °C and 110°C respectively (Sullivan, 1973). In the Dowds Hill orebody the first iron-rich phase to crystallize was magnetite which exhibits a bladed habit and is overgrown by later hematite with a few pyrite inclusions. Magnetite has not been observed in the " C " or " D " Deposit hydrothermal ores, and the first observed phase is hematite instead. During hydrothermal iron enrichment, BIFs surrounding the specularite pipes underwent some replacement, with carbonate laminations being completely pseudomorphed by randomly oriented hematite blades. Magnetite was replaced by specularite preferentially along octahedral planes, although this was far less complete than carbonate replacement. Most other components, i.e. quartz and silicates, remained unchanged. Mineralization ceased until the deposits were exposed to the atmosphere and processes of supergene enrichment commenced. These processes produced envelopes of hematite, goethite and limonite around the specularite mineralization (Fig. 5). SUMMARY AND CONCLUSIONS The isoclinal folds related to the first deformation event Dj, had little influence on the sites of iron mineralization. The second deformation D2 produced three sites of intense tectonic disruption and possible associated plutonic activity. The presence of a granite mass below the Middle BIF, in areas of intense deformation, supplied a heat source which enhanced the mineralization processes. Connate water and water expelled during prograde metamorphic dehydration migrated up brecciated F2 fold cores, such as Dowds Hill, " C " Deposit and " D " Deposit. During the early stages of the D2 brecciation, the ascending fluids leached iron from the BIF host leaving a friable quartzite residue associated with specularite, as seen at Dowds Hill and in the " C " Deposit breccias, and produced hot iron-rich solutions. The effect of decreasing pressure on the ascending hydrothermal fluids can be ignored as
437
IRON ORE IN KOOLYANOBBING GREENSTONE BELT metres 50
100
150
200
250
tholeiitic
[lH.'ll'ljll
lavas
supergene
ore
brecciated
& leached B.I.F.
hydrothermal
ore
Fig. 5. Simplified geological plan of the Dowds Hill orebody showing the main areas of D brecciation and associated hydrothermal mineralization. 2
equilibrium constants are only slightly affected by pressure in the temperature range of ore formation. As the iron-rich solutions rose and cooled they became saturated in iron and iron oxides precipitated. The solutions at Dowds Hill were hot enough and contained sufficient iron to precipitate magnetite at the commencement of mineralization, whereas cooler solutions precipitated hematite and associated pyrite, as has been theoretically predicted by Helgeson et al. (1969). Hydrothermal ore outcropping at " C " Deposit contains little magnetite, abundant quartz, and little pyrite, suggesting that solutions were cooler and hence represent a higher level above the heat source. Lack of pyrite in both "C" and "D" Deposits suggests the BIF at depth was sulphurdeficient and consequently hydrothermal solutions being derived from them were also sulphurdeficient. The initial stages of ore precipitation at "C" Deposit appear to have been under slightly different chemical conditions to the other deposits. What is now seen as a "hematite mud" probably precipitated as metastable Fe(OH) and dehydrated to hematite (Garrels & Christ, 1965). This type of precipitation probably represented a change in Eh/pH conditions in comparison to the other deposits (Wedepohl, 1978). The period of hydrothermal enrichment took place over a large time-span when compared with 3
the associated D tectonic event with which it is associated. The earliest precipitation of iron oxides occurred at Dowds Hill soon after the commencement of D , with "C" Deposit mineralization taking place toward the end of D , and "D" Deposit ores being younger than "C" Deposit ores. Hydrothermal mineralization at Dowds Hill and probably "D" Deposit, continued to some extent after F folding had finished. Uplift and erosion have exposed the hydrothermal pipes which maintained high porosities, as did surrounding BIFs, resulting in supergene enrichment by the action of meteoric water. The hematite formed by hydrothermal deposition is stable in the current weathering environment. The magnetite contained within the BIFs was martitized and/or hydrated whereas silica was leached and pyrite and silicates decomposed, resulting in envelopes of supergene ore around the hydrothermal deposits. 2
2
2
2
ACKNOWLEDGMENTS The author greatly appreciates the financial support and interest in this project of the B.H.P. Company Limited and Dampier Mining Company.
438
A. C. GRIFFIN
REFERENCES
ARCHIBALD, N . J . , BETTENAY, L . J . , BINNS, R . A . , GROVES, D . I., & GUNTHORPE, R . J . , 1978: T h e
evolution of Archaean greenstone terrains, Eastern Goldfields Province, Western Australia. Precamb. Res., 6, 103-131. ARRIENS, P. A., 1971: Archaean geochronology of Australia. Spec. PubisgeoI. Soc. Aust., 3, 11-23. BLAIS, S., AUVRAY, B., CAPDEVILA, R . , & HAMEURT, J . ,
1977: Les series komatiitiques et tholeiitiques des ceintures archeennes de roches vertes de Finlande orientale. Bull. Soc. geol. France, 19, 965-970.
BINNS, R . A . , GUNTHORPE, R . J . , & GROVES, D . I.,
1976: Metamorphic patterns and development of greenstone belts in the Eastern Yilgarn Block, Western Australia; in Windley, B. F. (Ed.) The Early History of the Earth, 303-313. Wiley, London. DORR, J. V. N., 1964: Supergene iron ores of Minas Gerais, Brazil. Econ. Geol., 59, 1203-1240. ELLIS, H. A., 1958: The exploratory diamond drilling of the Koolyanobbing iron deposits for pyrite. Bull, geol. Surv. West. Aust., 111. GARRELS, R. M., & CHRIST, C. L., 1965: Solutions, Minerals, and Equilibria. Harper, New York. GEE, R. D., 1979: Explanatory notes on the Southern Cross 1:250,000 Geological Sheet, Western Australia. Rec. geol. Surv. West. Aust., 1979/5.
HELGESON, H . C . , BROWN, T . H . , & LEEPER, R. H.,
1969: Handbook of Theoretical Activity Diagrams Depicting Chemical Equilibria in Geologic Systems Involving an Aqueous Phase at One A tm and 0° to 800°C. Freeman, Cooper & Co., San Francisco. HORWITZ, R. C., & SOFOULIS, J., 1965: Igneous activity and sedimentation in the Precambrian between Kalgoorlie and Norseman, Western Australia. Proc. Australas. Inst. Min. Metall., 214, 45-61. JAMES, H. L., 1954: Sedimentary facies of iron formation. Econ. Geol., 49, 235-293. NESBITT, R. W., 1971: Skeletal crystal forms in the ultramafic rocks of the Yilgarn Block, Western Australia: evidence for an Archaean ultramafic liquid. Spec. Pubis geol. Soc. Aust., 3, 331-347. SULLIVAN, C. M., 1973: Geology of South Range, Koolyanobbing, Western Australia. Thesis, Univ. West. Aust. [unpublished]. TUREK, A., & COMPSTON, W., 1971: Rubidium-strontium geochronology in the Kalgoorlie region. Spec. Pubis geol. Soc. Aust., 3, 72. WILLIAMS, I. R., 1975: Eastern Goldfields Province; in Geology of Western Australia. Mem. West. Aust. geol. Surv., 2, 33-54. WEDEPOHL, K. H. (ED.), 1978: Handbook of Geochemistry. Springer-Verlag, New York.
ECONOMIC EVALUATION OF ARCHAEAN FELSIC VOLCANIC ROCKS USING REE GEOCHEMISTRY* P. C. Thurston
Ontario Geological Survey, 77 Grenville Street, Toronto, Ontario, Canada ABSTRACT Geochemical evaluation of the mineral potential of Archaean felsic volcanic rocks has involved trace elements (Zn, S, Cu) and major elements (Mg, Na, K, Fe). Most methods applied in the Canadian Shield involve analysis of many samples followed by statistical evaluation to separate effects of normal igneous fractionation from chemical changes brought about during formation of volcanogenic massive-sulphide bodies. Trace elements such as Nb, Y, Zr, Cr and Ni and rare-earth elements (REE) are less mobile during metamorphism than elements previously used in economic evaluation of felsic volcanic rocks and are therefore more suitable. Seventy five samples (as on Fig. 1) from three superposed cycles of tholeiitic basalt to calcalkaline rhyolite in the Confederation Lakes area of NW Ontario have been analyzed for Nb, Zr, Y and REE. The Confederation Lakes results are compared with 18 unmineralized tholeiitic rhyolites, and mineralized and unmineralized calc-alkaline rhyolites from the Abitibi belt of NE Ontario. Cycle I and II rhyolites (unmineralized) in the Confederation Lakes area have highly fractionated REE patterns with abundances of 20-150 times chondrites for light REE to 2-40 times chondrites for heavy REE. In the mineralized Cycle III rhyolites, samples from felsic flows, pyroclastics and an endogenous dome have relatively unfractionated REE patterns with abundances of 120-240 times chondrites for light REE and 53-100 times chondrites for heavy REE. Similarly, mineralized calc-alkaline rhyolites from the Kam-Kotia area within the Lower Supergroup of the Abitibi belt (Pyke, 1978) yield unfractionated REE patterns with abundances of 100-135 times chondrites for light REE and 30-50 times chondrites for heavy REE. Within Pyke's Upper Supergroup an unmineralized tholeiitic rhyolite in Dokis Township has an unfractionated pattern with 80 times chondrites for light REE and 25-30 times chondrites for heavy REE. However, unmineralized rhyolites of tholeiitic affinity within the Upper Supergroup exhibit fractionated patterns. The conventional fractionated REE patterns in all areas are interpreted in terms of partial melting of amphibolites and/or magma mixing. Unfractionated REE patterns in rhyolites associated with volcanogenic massive-sulphide deposits can be explained by the complexing and fixing of heavy REE in carbonate-rich fluids circulating through the pile during the mineralizing event. However, the calc-alkaline rhyolites associated with the mineralization in the Confederation Lakes area are underlain by variolitic, highly fractionated basalts with few rocks of intermediate composition. Analysis of the variolites reveals a similar unfractionated pattern, suggesting that the lack of fractionation of REE in the felsic rocks could also be related to production of felsic magmas by liquid immiscibility. Immiscibility may have resulted in a carbonate-rich fluid phase capable of complexing heavy REE.
INTRODUCTION This paper draws attention to the unusual patterns of rare-earth elements (REE) in felsic metavolcanic rocks associated with some Archaean volcanogenic massive-sulphide ore bodies, and offers an explanation. The work is at an early stage, and more complete documentation will follow. Geochemical exploration for volcanogenic massive-sulphide deposits (as defined by
Sangster, 1972) has involved analysis of numerous samples of felsic volcanic rocks (c/. Davenport & Nichol, 1973; Descarreaux, 1973; Wolfe, 1975). The concepts developed from this work are that the volcanogenic massive-sulphide deposits are marked by haloes of Na, Ca, and rarely K depletion (Sakrison, 1966; Franklin et al. 1975; Goodfellow, 1975) and enrichment of Mg, Fe, Si, and K (Sakrison, 1966; Simmons et aL, 1973; Franklin et al., 1975). The alteration haloes prob-
* Approved for publication by the Director, Ontario Geological Survey.
9
440
Fig. 1.
P. C. T H U R S T O N
Geological m a p of the Confederation Lakes Area, Northwestern Ontario: geology after Thurston (1980).
ably result from circulation of ocean water through porous parts of the felsic pile. The restricted extent of the alteration haloes prompted Sopuck's approach in which discriminant analysis was used to recognize hydrothermal imprints upon the magmatic chemistry of the felsic metavolcanic rocks (Sopuck, 1977). He found that there was enrichment of Fe2C>3 and Zn, and leaching of Na 2 0, over larger areas than detectable by conventional methods. The main problem inherent in this method is that many samples must be analyzed to identify successfully the effects of the mineralizing hydrothermal circulation system. Previously, REE were generally regarded as immobile during various types of hydrothermal alteration and metamorphism (Frey et al., 1968; Herrmann et al., 1974; Ferrara et al., 1976). Therefore REE have provided a means of
examining crustal petrogenetic processes such as partial melting and fractionation. Collerson & Fryer (1978) explain that trondhjemites associated with areas of high-rank metamorphic rocks are depleted in heavy REE by complexing the heavy REE with C0 3 ~ 2 ions, thereby suggesting a mechanism for removal of heavy REE. Further evidence for REE mobility during hydrous burial metamorphism has been summarized by Hellman et al. (1979). They describe four types of REE mobility: (1) gross REE and selective light REE enrichment in ultramafic, basaltic and andesitic volcanic rocks (Sun & Nesbitt, 1978; Hellman & Henderson, 1977; Chikhaoui et al., 1978) in which all REE increase in abundance, with an exaggerated increase of the light REE; (2) REE distribution about a primary mean in which enrichment and depletion in a group of
441 (f) plagioclase: Eu is concentrated in this mineral during crystallization, thus depleting the liquid. Therefore fractionation of the common mafic rock-forming minerals generally serves to deplete felsic differentiates in the heavy REE. Graf (1977) noted that the formation of volcanogenic massive-sulphide ores results in the alteration of ferromagnesian minerals and glass in felsic volcanic rocks by ore-forming hydrothermal fluids. Initially Cu and Fe are leached to form the Cu-rich base of Kuroko-type massivesulphide bodies, followed by preferential alteration of ferromagnesians and feldspar by concentrated brines enriching the ore-forming solution in Pb and Zn which tend to form the upper part of Kuroko orebodies. The heavy REE produced by the alteration of felsic volcanic glass yield unfractionated REE abundances in the circulating hydrothermal fluid. Enriched heavy REE abundances in felsic volcanic rocks may be a consequence of this model if sufficient volumes of fluid circulate through the felsic volcanic rocks. Taylor & Fryer (1980) described limited mobility of the heavy REE in systems forming porphyry copper ore, and Kerrich & Fryer (1979) noted extensive mobility of the heavy REE in hydrothermal systems associated with gold mineralization. The apparent paradox of leaching of heavy REE from the felsic volcanic rocks in the ore-forming hydrothermal system (Graf, 1977), and the enrichment in heavy REE reported here, may be analogous to the fact that depletion of base metals within the hydrothermal system of volcanogenic massive-sulphide orebodies cannot be detected in the surrounding rocks within analytical precision. REE may be removed in small amounts from large volumes of the surrounding volcanic rocks to produce the large volumes of felsic volcanic rocks with enhanced heavy REE contents described in this paper.
ECONOMIC EVALUATION USING REE GEOCHEMISTRY Fig. 1. Legend. Mafic
metavolcanics
Felsic m e t a v o l c a n i c s l l | § l | | | ] I n t e r m e d i a t e metavolcanics Granodiorite + q t z - f e l d porph Clastic metasediments G r a n i t i c rocks Chemical 1 major
metasediments
fault
anticline syncline geological contact
samples occurs about mean values brought about by textural heterogeneities and local variability of fluid activity (Sun & Nesbitt, 1978); (3) REE depletion brought about by hydrothermal leaching (Hellman & Henderson, 1977), by dilution by phases such as zeolites, which do not contain appreciable REE, or by volume increase; (4) Selective REE mobility demonstrated for several elements and groups of elements, for example Ce (Masuda & Nagasawa, 1975) Eu (Sun & Nesbitt, 1978) La and Yb (Robertson & Fleet, 1976). Fractional crystallization is the classical means of progressing from a basaltic primary magma to felsic differentiates. The effect of crystallization of various phases on the REE abundances in the liquid has been summarized by Hanson (1978) as follows: (a) olivine: crystallization of olivine raises all REE abundances fairly uniformly; (b) orthopyroxene: hypersthene preferentially raises the abundance of the light REE and contributes to a positive Eu anomaly; (c) clinopyroxene: crystallization of clinopyroxene preferentially enriches the liquid in medium REE and somewhat in the light REE and contributes to a positive Eu anomaly; (d) hornblende: the main effect of this mineral is to deplete the liquid in the middle REE and to some extent in the heavy REE, as well as contributing to a positive Eu anomaly; (e) zircon and apatite: crystallization of these minerals severely depletes the liquid in the middle and heavy REE;
VOLCANICS OF THE CONFEDERATION LAKES AREA As a test of this model the author has examined the trace-element aspects of Archaean volcanism in the Confederation Lakes area (Thurston, 1980) in which three mafic to felsic cycles comprise 8500 m of strata (Fig. 1). Metamorphic grade in the area is generally lower greenschist (Thurston & Breaks, 1978). Each cycle consists of tholeiitic pillowed basalts with minor intercalated andesites followed by calc-alkaline intermediate and felsic pyroclastics and flows. Rhyolites occur in the upper part of each cycle. The following synopsis of chemical relations in the Confederation Lake
P. C. THURSTON
442
area is summarized from Thurston (1980). Chemical variations in the basalts are consistent with fractionation of olivine and plagioclase with late accumulation of clinopyroxene. The intercalated andesites are most probably derived by mixing of tholeiitic basalt and trondhjemite liquids. The dacite to rhyolite samples of cycles I and II have fractionated REE patterns which probably originated by partial melting of amphibolite or by mixing of rhyolite and trondhjemite liquids. Cycle III differs from the lower two cycles with respect to trace-element abundances in the basalts and rhyolites. Cycle III basalts include a massive variolitic unit, which, by analogy with the textures and major-element chemical data presented by Gelinas et al. (1976) for a similar unit in the Abitibi sub-province, also 200 r
La 300
Ce
Sm Eu
Figure 2b
probably originated by liquid immiscibility. The REE pattern of the variolitic basalts (Fig. 2a, Table I) resembles those of the overlying cycle III rhyolites (Table I, Fig. 2b) in that the heavy REE are enriched. The South Bay mine, a volcanogenic Cu-ZnAg massive-sulphide ore body, is found at the contact of an endogenous dome of rhyolitic quartz-feldspar porphyry and overlying rhyolitic pyroclastic breccias of Cycle III (Pollock et al., 1972). Conventional crystal fractionation from a basaltic parent magma is unable to explain the enrichment of Cycle III rhyolites in heavy REE, and the cause may be transport of heavy REE in C0 "-rich fluids (cf. Collerson & Fryer, 1978; Kerrich & Fryer, 1979) by complexing with C0 ~ .ions. Evidence for this hypothesis is the 3
3
2
Yb Lu
Gd
CyclelK Felsic Volcanic Rocks Confederation Lake Area
200
100 80 60 40 unfractionated rhyolites La
-I
Ce
1
Nd
1
Sm
I
Eu
L
Gd
Dy
Er
_L
Yb Lu
Fig. 2. REE abundances: (a) Confederation Lake area cycle III variolitic basalts; (b) Confederation Lake area cycle III rhyolites.
ECONOMIC EVALUATION USING REE GEOCHEMISTRY
increasing degree of carbonatization, reflected in increasing modal carbonate, of the quartzfeldspar porphyry at South Bay as the orebody is approached (Sopuck, 1977; Thurston, 1978). To test the hypothesis, further felsic metavolcanic rocks from the Abitibi Sub-province were analyzed from mineralized and barren environments. Alternatively, the unfractionated rhyolites may result from the existence of a compositionally zoned felsic magma chamber unrelated to the underlying basalts (cf. Hildreth, 1979).
VOLCANICS OF THE ABITIBI SUB-PROVINCE General Geology The Abitibi Sub-province, including the Abitibi volcanic-sedimentary belt, is the world's largest greenstone belt. Metamorphic facies in the central part of the Abitibi belt is prehnitepumpellyite in the area of the samples supplied by Jensen (Table I) and greenschist in the Timmins area and the Kam-Kotia area (Jolly, 1978). The belt comprises two volcanic cycles and the upper part of a third cycle (Jensen, 1979). The two complete cycles contain a basal part dominated by peridotitic to basaltic komatiites succeeded upward by Mg-rich tholeiitic basalts which become progressively more Fe-rich by fractionation of clinopyroxene, plagioclase and olivine (Jackson, 1980). Among the tholeiitic basalts are minor intercalated rhyolitic tuff, flows and hypabyssal intrusions. Above the tholeiitic part of the cycle, calc-alkaline basalt to rhyolite sequences occur. Thickness of the cycles is variable, but a maximum of 30000 m is suggested by Jensen (1979) for one cycle in the southern part of the subprovince. The upper calc-alkaline part of the lowermost cycle is preserved in the Kirkland Lake area (Jensen, 1979). The upper two cycles have been designated the Lower Supergroup and the Upper Supergroup by Pyke (1978) as shown in Figure 3. The Lower Supergroup is known as the Deloro Group in the Timmins area, and the Upper Supergroup is known as the Tisdale Group (Pyke, 1980). Jensen (1978a) has described majorelement differentiation trends for mafic to felsic volcanic rocks in several areas of the belt, defining a tholeiitic trend ranging from basalt to distinctive high-Fe rhyolites, and a calc-alkaline trend ranging from high-alumina basalts to rhyolites characterized by lower Fe, Ti, and higher alkalis relative to tholeiitic rhyolites. Several volcanogenic, massive-sulphide orebodies occur in rhyolites of the calc-alkaline upper part of the lower Supergroup, including the
443 Kidd Creek mine, and several smaller orebodies occur in the Robb and Jamieson Township area. Tholeiitic Rocks of the Upper Supergroup Samples of tholeiitic high-silica rhyolite (terminology after Fryer & Jenner, 1978) from the Upper Supergroup of Pyke (1978) were analyzed (Nos 1, 2, 3 in Table I). The rocks are sodic and trondhjemitic as defined by Streckeisen (1976). The REE abundances (Table I, Fig. 4) of three samples are relatively unfractionated ranging from 80 to 120 times chondrites at the light REE and through 40 to 70 times chondrites at the heavy REE end, with negative Eu anomalies. These rhyolites are underlain by tholeiitic basalts for which Condie & Baragar (1974) and Jackson (1980) report flat REE patterns with abundances ranging from 10 to 22 times chondrites and negative Eu anomalies. Calc-Alkaline Rocks of the Upper Supergroup Samples of calc-alkaline andesite, dacite, and high- and low-silica rhyolite from the Upper Supergroup of Pyke (1978) were analyzed. Samples 71J444, 71J446 and 71J451 are andesite to dacite which display a relatively flat REE pattern depleted in heavy REE relative to samples plotted in Figure 4. They are similar to Condie's depleted Archaean andesite (DAA) (Condie, 1976), with abundances of about 20-50 times chondrites and positive Eu anomalies (Fig. 5). The samples have relatively high LOI values and modal carbonate, probably brought about by hydrothermal alteration, but the REE do not appear to have been affected (compare with the data of Hellman et al., 1979, and Menzies et al., 1979). The dacites and low-silica rhyolites (Table I, Fig. 6) have fractionated patterns with heavy REE abundances below 10 times chondrites, typical of Condie's depleted siliceous volcanic (DSV) group (Condie, 1976), and light REE abundances in the 10-30 times chondrites range, and variable positive and negative Eu anomalies. The high-silica rhyolites form two groups based upon REE characteristics: samples 71J26, P528-70, P8-73 and P72-76 have highly fractionated REE patterns with heavy REE abundances being 10 times chondrites, light REE abundances ranging from 25 to 50 times chondrites, and pronounced negative Eu anomalies characteristic of the DSV group of Condie (1976). Sample 71C3 has a relatively unfractionated pattern typical of Condie's undepleted siliceous volcanics (USV) (Condie, 1976). Calc-Alkaline Rocks Two high-silica rhyolites (68-664, 68-789) from the Lower Supergroup at Jamieson Township
T A B I I :
I
Major-element and REE analyses of felsic inelavo/canics and associated rocks of the Upper and Lower Abitibi Sub-province and Cycle III of Confederation Lakes area
Supergroup,
7
8
9
To
T1
\2
i"3
Hi
Ts
T5
71J30
71C3
P528-70
P72-76
71J26
P6-73
68-664
68-805
68-753
68-789
449
450
429
425
468
432
7 6 ,. 4 0
77 . 2 0
7 8 .. 2 0
62..4
7 1 •, 8
78..0
71..0
7 3 ..4
76.,8
66. 4
74. 5
71 . 6
74. 8
75. 9
77-9
8 ,. 9 7
12 . 8 0
9 .• 5 5
1 6 .• 5 0
1 4 .• 9
1 2 ,. 3
1 2 .. 8
13. 1
1 2 .• 5
13. 7
1 1. 0
12. 0
10. 9
1 1 .2
12.0
3
4
5
6
71J600
72D226
71JA5I
71J446
7U444
75 . 6 0
75 . 8 0
75 • 30
50 • 50
60 .50
57 . 10
78.40
AI2O3
10 . 5 0
10 • 5 0
10 . 4 0
19 . 7 0
16 . 9 0
17 . 2 0
1 1 .00
7
T7
TB
19
20
IT
22
Fe203
2 .54
0 • 91
2 .11
6 • 94
5 ,. 7 7
6 .00
2 • 53
0 .67
0 ,• 3 7
3 ., 5 6
1 ., 8 1
N i 11
1 .• 0 9
N i I
2. 26
1 . 10
2. 9
FeO
0 .49
2 • 79
l .18
0 .00
0 .00
0 .00
0.65
0 ,. 0 0
0 .66
2 • 3-6
0 .. 0 0
0 .. 3 3
1 •. 5 3
3 • 93
0.67
6. 50
3- 72
4. 81
5- 71
3 . 14
MgO
0 .29
0 • 69
0 • 35
1. 1 7
1.30
1.55
0.35
0 ,. 5 9
0 . 10
2 ,. 6 9
2 .• 5 3
1. 6 3
0 .. 10
0 .• 5 5
0 .. 8 1
0.
OO
1 7IJ630 SiO:
1 .9 1
1 .0 0
1 .5 5
0. 58
0. 09
1.14
CaO
1.60
1 • 30
2 .28
13 . 10
1 0 .. 3 0
10 . 10
1.53
3 .42
0 .82
2 • 52
8 .. 1 5
1. 3 9
0 .. 9 0
1 .. 2 2
1 .• 3 4
1 .. 1 0
1 . 11
2. 40
2. 35
1. 3 0
2. 82
0.29
0 .. 7 4
0 .. 3 0
4.96
Na20
5 .20
1 ., 6 4
1 .2 9
1 .6 6
1.48 1 .59
1. . 8 1
1 .. 1 9
1 .21
5-14
3 .• 1 3
5 .. 2 4
0 ,. 8 2
3 .. 3 2
0 . • 11
3 .. 4 5
5 .• 5 0
5. 25
5- 08
4. 21
2. 94
4 . 18
K2O
0 .• 3 8
1 .. 0 4
0 .. 17
0 .. 3 4
0 .. 10
0 . 10
0.53
0 .. 6 8
2 ,. 1 7
0 ,. 8 7
0 ., 2 2
1 .52
7 .. 12
7. 36
2 .. 0 5
0 .• 3 7
1 ., 2 1
0 . 51
0 . 15
1 .37
0. 53
Ti02
0 .. 2 4
0 .. 2 6
0 .. 0 2
0 .. 8 9
0 .. 8 2
0 .. 8 5
0.15
0 ,. 1 9
0 .. 0 6
1 .,8 4
0 ., 6 0
2. 00
0 .. 1 8
0 .. 2 8
0 .• 3 9
0 ., 0 6
0. 77
0. 62
0 . • 12
0. 49
0. 39
0.23
P2O5
0 .. 0 3
0. 03
0 .. 0 1
0 ., 0 0
0. 00
0 .. 0 0
0.03
0 ,. 0 0
0 .. 0 1
0 ,. 1 1
0 ., 0 0
0. 21
0 ., 0 2
0 .• 0 5
0 .. 0 5
0 ., 0 1
0 .. 0 6
0. 04
0 .. 1 3
0 ..01
0. 01
0.00
0. 57
S
1.54
0.80
3 .. 8 5
0. 28
0 . 01
0. 05
0 .. 0 0
0 .. 0 0
0 .. 0 0
0.01
0 .. 0 0
0 .. 0 4
0 .. 0 5
0 .. 0 0
0. 09
0 ., 0 1
1 .32
0 .• 0 7
0 .. 0 1
- 0 . 01
- 0 .,01
0. 06
0.02
MnO
0. 05
0. 05
2. 84
0. 00
0. 00
0 .. 0 0
0.03
0 .. 0 0
0 .. 2 9
0 ,. 0 1
0 .. 0 0
0. 00
0 .,01
0 .• 0 9
0 ,. 10
0 .. 0 4
0 ., 1 1
0 . 10
0 . , 11
0 . 15
0 ., 1 1
0.04
C02
0. 00
0. 00
0. 08
0. 00
0. 00
0 .. 0 0
0.00
0 .. 0 0
0 .• 74
0 .. 0 8
0 ., 0 0
0 ., 0 2
0. 43
0 .• 9 6
0 ,. 4 0
1 .• 7 4
0. 30
0. 30
0 .• 7 4
1 ,. 2 0
0 ., 0 8
0.24
H20+
0. 02
0. 28
0. 02
0. 00
0. 00
0. 00
0.04
0 .. 0 0
0 .. 0 6
0 .• 7 8
0 ..00
0 .,11
0 ., 8 4
1 .. 1 6
0 .• 3 8
2 .. 0 3
1. 33
1. 5 6
1 •. 3 7
0. 55
0.68
H2O"
0. 00
0. 00
0. 00
0. 00
0. 00
0. 00
0.00
0 .. 0 0
0 .. 0 0
1 .. 4 7
0 ., 0 0
0 .. 0 9
0 .. 0 4
0 .. 1 4
0 ., 0 2
3 . 10
3. 40
3. 80
0 .. 0 5
2 .. 4 0
1 .9 5
9 7 . 55
100. 28
9 7 . 91
99-,68
100. 4
LOI TOTAL:
98. 24
98. 88
99- 02
2. 50 100.08
9 8 . 41
100.87
99..2
1 0 1 .• 5
100..2
100.,8
100. 1
100,. 9
100.• 33
98..9
100. 1
8
8
-5
-5
-5
-5
Ni
-5
-5
-5
-5
-5
-5
Co
7
8
8
-5
-5
Cu
8
5
-5
-5
8
8
Zn
130
92
144
112
112
150
Pb
-10
-10
-10
48
48
14 290
Cr
99. 9
99..7
(ppm)
-5
Ba
170
140
50
240
240
Sr
90
70
50
145
145
Rb
43
12
1
19
19
44
Zr
547
438
464
545
545
385
Nb
50
32
30
36
43
43
41
140
1 10
132
152
147
202
La
27
36
25
13
9
16
28
25
8
16
12
5
45
44
33
22
46.,6
37. 1
42.. 6
45.. 7
58.0
Ce
68
97
72
28
21
37
63
56
21
31
24
14
111
107
75
53
98..7
86. 7
105
104
122
Pr
10
13
10
3
2
4
7
6
2
3
2
1
14
13
8
6
V
1
s i g n i f i e s
less
than
1
ppm.
TABLE
I
(continued)
1 71J630
2 71J600
3 72D226
4 7U451
5 7U446
6 71J444
7 71J 30
8 7IC3
Nd
53
74
50
15
9
19
37
35
Sm
15
23
15
9
11 71J26
12 P8-73
13 68-664
l4 68-805
15 68-753
16 68-789
17 449
18 450
19 429
20 425
15
12
6
68
57
42
30
58.6
54.6
63.6
66.2
77-3
4
5
19.4
20. 7
10 9 P528-70 P72-76 11
16.6
18.1
21 468
22 432
4
2
11
3
3
12
9
Eu
3
4
1
1
-1
2
2
3
-1
-1
1
-1
4
4
3
2
3-78
1.34
3-72
4.38
2.93
Gd
12
24
18
4
2
4
. 8
10
3
2
3
-1
20
12
12
8
20.0
18.3
20.3
23.5
21.5
o
-1
-1
1
-1
4
3
2
1
8
14
O
Tb
2
4
3
Dy
11
19
-1
5
-1
-1
-1
2
1
19
15.7
14
4
2
-1
-1
5
2
2
24.70
20.90
21.0
25.1
19.5
Ho
3
5
3
-1
-1
2
-1
3
-1
-1
1
-1
3
2
2
2
Er
11
14
11
3
-1
3
4
7
1
1
2
-1
12
6
8
6
15.8
13-2
13-5
15.6
10. 1
Yb
13
1 1
11
-1
-1
•-1
-1
4
-1
-1
1
-1
7
5
7
4
15-6
13.0
10.5
14.4
7.26
Lu
2
-1*
2
-1
-1
14
-1
2
-1
-1
1
-1
1
-1
-1
-1
*-] s i g n i f i e s
2
-1
19
6
m <
>
r-
c >
less than 1 p p m .
Analytical Methods: M a j o r e l e m e n t s by XRF at the M i n e r a l Research L a b o r a t o r y , O n t a r i o G e o l o g i c a l F e + 2 by titration against K M n O u . REE a n a l y s e d by the thin film XRF m e t h o d of Fryer (1977).
m
Survey.
mineralized,
c
1)
R h y o l i t e , light g r e e n , a p h a n i t i c , t h o l e i i t i c a f f i n i t y , u n m i n e r a 1 i z e d , U p p e r S u p e r g r o u p (Jensen, 1978b).
12)
Rhyolitic q u a r t z - f e l d s p a r crystal tuff, c a l c - a l k a l i c a f f i n i t y , U p p e r S u p e r g r o u p (Pyke, 1980).
2)
Rhyolite to d a c i t e , light grey cherty m a s s i v e tholeiitic a f f i n i t y , unmi neral i zed , U p p e r S u p e r g r o u p (Jensen, 1978b).
13)
Rhyolite, grey-black, cherty, calc-alkalic affinity, mineralized, Lower Supergroup (Middleton, 1973).
Z o
3)
R h y o l i t e , light g r e y , c h e r t y , tholeiitic a f f i n i t y , u n m i n e r a 1 i z e d , U p p e r S u p e r g r o u p (Jensen, 1978a).
14)
Rhyolite, brown-pink, brittle, calc-alkalic affinity, mineralized, Lower Supergroup (Middleton, 1973) -
m m
4)
A n d e s i t e , light g r e y , a p h a n i t i c , strongly a l t e r e d , c a l c - a l k a l i c a f f i n i t y , u n m i n e r a l i z e d , U p p e r S u p e r g r o u p (Jensen, 1978b).
15)
Rhyolitic w e l d e d t u f f , Jameland ore z o n e , c a l c - a l k a l i c m i n e r a l i z e d , Lower Supergroup (Middleton, 1973)
affinity,
5)
A n d e s i t e to d a c i t e , g r e e n i s h g r e y , a p h a n i t i c , calc-alkalic a f f i n i t y , u n m i n e r a l i z e d , U p p e r Supergroup (Jensen, 1978b).
16)
R h y o l i t e , cherty g r e y , banded c a l c - a l k a l i c a f f i n i t y , m i n e r a l i z e d , Lower S u p e r g r o u p (Middleton, 1973).
6)
A n d e s i t e to d a c i t e , g r e e n i s h g r e y , a l t e r e d , w i t h c a l c - a l k a l i c a f f i n i t y , u n m i n e r a l i z e d , U p p e r Supergroup (Jensen, 1978b).
17)
Matrix of variolitic b a s a l t , Cycle i l l , C o n f e d e r a t i o n Lakes a r e a , tholeiitic affinity (Thurston, I980).
7)
R h y o l i t e , quartz and feldspar p h y r i c , c a l c - a l k a l i c a f f i n i t y , u n m i n e r a 1 i z e d , L o w e r S u p e r g r o u p (Jensen, 1978b).
18)
Variole of variolitic b a s a l t , Cycle 111, C o n f e d e r a t i o n Lakes area (Thurston, I98O).
8)
Dacite to rhyolite quartz and feldspar p h y r i c , c a l c - a l k a l i c a f f i n i t y , u n m i n e r a 1 i z e d , Lower Supergroup (Jensen, 1978b).
19)
Low silica rhyolite flow, Cycle III, C o n f e d e r a t i o n Lakes area (Thurston, 19&0).
9)
T r o n d h j e m i t i c q u a r t z - f e 1 d s p a r porphyry c a l c - a l k a l i c a f f i n i t y , u n m i n e r a l i z e d , U p p e r Supergroup (Pyke, 1980).
20)
Low silica rhyolite f l o w , Cycle III, C o n f e d e r a t i o n Lakes area (Thurston, 1980).
10)
Rhyolitic quartz sericite t u f f , c a l c - a l k a l i c a f f i n i t y , u n m i n e r a 1 i z e d , U p p e r S u p e r g r o u p (Pyke, I98O).
21)
High silica rhyolite flow, Cycle III, C o n f e d e r a t i o n Lakes area (Thurston, 1980).
11)
D a c i t e , light grey t u f f a c e o u s , c a l c - a l k a l i c a f f i n i t y , u n m i n e r a 1 i z e d , U p p e r S u p e r g r o u p (Jensen, 1978b).
22)
High silica rhyolite porphyry from e n d o g e n o u s d o m e , South Bay M i n e , Cycle III, Confederation Lakes area (Thurston, I98O).
in
73
m
O o
X
m
70
IM
446
P . C. T H U R S T O N
*- v
v y "
ARCHEAN r / / J
Other granitic rocks
1°Q ° o ° | Porcupine Group UPPER SUPERGROUP P i l l ! ! ! ! Calc-Alkalic suite Tholeiitic suite I'vV'/.vlj Komatiitic suite LOWER SUPERGROUP 1-VyJr/l Unsubdivided
strike slip faults
KENOGAMISSI * BATHOLITH . X X X * * to f V X X X X x xl x\\ V'--' X X X x V x \ l\xy\ *
^ ^
transcurrent faults
x
X
X
X
X
X
X
X \ X X X X« X
X
—
* •
^ S ^ ' x X \x " x\ X X X X x\ X . yrw •* x * x * x * x \ x x
/^.y-'
X
t
x x x x v
10
x
v ^ / i ^ j ^ n V I -
kilometres 20 30 10 15 miles
Fig. 3.
€R LAKE x * x *, FAULT *
— pore;
40 20
25
T
^ ^ —
anticline syncline geological contact
"ROUND x \ - - LAKE * x BATHOLITHx'
General geology of part of the Abitibi Sub-province a f t e r Pyke (1978). Locations of samples in Figures 4, 5, 6, 7 shown approximately: 4 = tholeiitic rhyolites of Figure 4. 5 = andesites of Figure 5. 6 = calc-alkaline rhyolites of Figure 6. 7 = mineralized calc-alkaline rhyolites of Figure 7.
and two low-silica rhyolites (68-805, 68-753) (Middleton, 1973) which form the calc-alkaline upper part of the Lower Supergroup (Pyke, 1978) were analyzed (Fig. 7). The samples are f r o m the area of the Jameland Mine, a volcanogenic CuZn orebody, and show relatively unfractionated REE patterns ranging f r o m 68-130 times chondrites at the light REE and 20 to 55 times chondrites at the heavy REE end. If Condie's review of petrogenetic conclusions with respect to felsic volcanics is valid (Condie, 1976), the occurrence
of rhyolites of two different origins within the same stratigraphic unit is a n o m a l o u s . Unusual mechanisms, such as volatile complexing (Collerson & Fryer, 1978) or the existence of compositionally zoned m a g m a chambers, must be invoked (Hildreth, 1979) to explain the behaviour of REE in samples 71C-3 and 71J-30. INTERPRETATION The flat R E E pattern of the tholeiitic highsilica rhyolites f r o m the Kinojevis G r o u p (Fig. 4)
E C O N O M I C E V A L U A T I O N U S I N G REE G E O C H E M I S T R Y
447
clinopyroxene, plagioclase and olivine from a parental basalt (48.75% S i 0 2 , 7.87% M g O , 11.71% FeO) with a flat R E E pattern at about 6 times chondrites to an evolved high (15.29%) FeO1 basalt (about 32 times chondrites). Continued fractionation of clinopyroxene and plagioclase (10% liquid remaining) would produce the observed patterns for the tholeiitic rhyolites.
Lo
Fig. 4.
Fig. 5.
Ce
Pr
Nd
S;m
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
REE abundances, 3 tholeiitic rhyolites Kinojevis Group, Upper Supergroup.
Fig. 6.
REE abundances calc-alkaline rhyolites, Blake River Group, Upper Supergroup, not mineralized.
Fig. 7.
REE abundances calc-alkaline rhyolites, Deloro Group, Lower Supergroup—mineralized.
REE abundances 3 calc-alkaline andesites, Blake River Group, Upper Supergroup.
is similar to that of the dacite described by Hanson (C in Table 3, 1978), except for the high abundances present in the Kinojevis Group sample. The R E E patterns of these samples, when compared to the underlying tholeiitic basalts of the Kinojevis Group, have most probably originated by fractionation of clinopyroxene and plagioclase from a basaltic parent magma. Jackson (1980) describes the evolution of the underlying parental basalts in terms of fractionation of
448
P. C. THURSTON
The fractionated REE pattern of most of the dacites and low-silica rhyolites of the Upper Supergroup probably originated by 10-20% partial melting of garnetiferous amphibolite or eclogite, if one considers the general similarity of the REE pattern to those modelled by Condie (1976) and Arth & H a n s o n (1975). The high-silica rhyolites with less fractionated REE patterns (71J30 and 71C3) are both f r o m a small body of rhyolitic quartz-feldspar porphyry intrusive into the calc-alkaline basalts of the Upper Supergroup, and may represent 30-60% partial melting of granulites by analogy with the data of Condie (1976). The calc-alkaline high-silica rhyolites f r o m the Deloro Group are enriched in heavy REE in a fashion similar to the Cycle III rhyolites of the Confederation Lake area. While it is possible to derive unfractionated REE patterns for rhyolites of tholeiitic affinity by fractionation of clinopyroxene, olivine and plagioclase, there is no evidence of this for calc-alkaline rhyolites in either the Confederation Lakes area or the Lower Supergroup samples. The volume of rhyolite involved is much too large to have been derived by such fractionation in either area. In the Confederation Lakes samples (Thurston, 1980), 80% fractionation in the basalts estimated by Zr abundances has produced basaltic andesite with a Zr content of 199ppm, whereas the rhyolites with Zr contents ranging f r o m 400 to 650 would require an additional 50 percent fractionation which seems unrealistic on major-element grounds. In the Deloro Group, using Cr as an estimate of fractionation, rhyolitic rocks are not directly derived f r o m the underlying basalts by crystal fractionation. Therefore, some mobility of the heavy REE may be involved in the production of those unusual REE patterns, or fractionation in a high-level magma chamber (cf. Hildreth, 1979). Leaching of the medium REE (Sm-Dy) in metabasic schists accompanied by enrichment in H o through Yb is reported by Kerrich & Fryer (1979) for a C0 3 ~ 2 -rich ore-forming hydrothermal system. They note the similarity of such heavy REE enrichment to enrichment of heavy REE in pneumatolytic alteration products in
alkalic igneous complexes (Fryer & Edgar, 1977). Nesbitt (1979) also describes leaching of the heavy REE in the weathering of a granodiorite by low pH fluids. The acidity of fluids which form volcanogenic massive sulphides is well known (Fryer & Hutchinson, 1976; Urabe & Sato, 1978). The most likely transporting medium is C0 3 ~ 2 , fluoride, or sulphate complexes (Herrmann, 1970). In the Confederation Lakes area, rhyolites of cycle III are underlain by variolitic basalts, and given the similarity of the REE patterns of the rhyolites to the variolitic basalts, it could be argued that the unusual pattern of enrichment in heavy REE is the signature of a liquid immiscibility event. However, there are no variolitic basalts underlying the Deloro Group rhyolites (pers. c o m m . D. R. Pyke, 1980) which have a generally similar pattern of enrichment in heavy REE. Therefore enrichment of heavy REE in felsic metavolcanics by C0 3 ~ 2 -rich mineralizing fluids is favoured for mineralized felsic volcanics of Archaean age although the preliminary nature of these data is recognized. Other possible mechanisms include hornblende fractionation to produce the Deloro rhyolites f r o m an andesitic parent magma (Fountain, 1979). It is still an interesting coincidence that rhyolites associated with volcanogenic massive-sulphide mineralization are rather uniformly enriched in heavy REE, whatever the cause. The REE geochemistry of felsic metavolcanic rocks may indicate their potential as hosts to volcanogenic massive-sulphide ores without recourse to analysis of large numbers of samples.
ACKNOWLEDGMENTS I am indebted to C. Riddle and R. Faulkner for the analyses for rare-earth elements performed under trying circumstances. D. Pyke and L. Jensen provided access to analyzed samples for this work and gave advice on the geology of the Abitibi Sub-province. The manuscript has benefitted f r o m reviews by V. G. Milne, D. R. Pyke and L. S . J e n s e n .
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F.
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1978:
Meta-
morphic and tectonic evolution of the Uchi-English River Subprovince; in Fraser, J. A., & Heywood, W. W. (Eds) Metamorphism in the Canadian Shield, 49-62. Pap. geol. Surv. Can., 78-10. URABE, T . , & SATO, T . , 1978: K u r o k o deposits of the
Kosaka mine, Northeast Honshu, Japan—products of submarine hot springs on Miocene sea-floor. Econ. Geol., 73, 161-179. W . J., 1975: Zinc abundance in early Precambrian volcanic rocks: its relationship to exploitable levels of zinc in sulphide deposits of volcanic-exhalative origin; in Elliott, I. L., & Fletcher, W. K. (Eds) Geochemical Exploration 1974, 261-278. Elsevier, Amsterdam.
WOLFE,
CONCEPTUAL MODELS
IS OCEAN FORMATION SYNCHRONOUS WITH THE FIRST PRESERVATION OF CRUST? U. R. Costa,l W. S. Fyfe, R. Kerrich & H. W. Nesbitt On leave from Federal University of Bahia, Brazil Department of Geology, University of Ontario, London, Ontario, Canada 2
2
2
1
2
ABSTRACT The oldest preserved crust appears at about 3.8Ga. Many workers consider that meteoritic events and even the formation of magma oceans may have obliterated the first crust of the accreting planet. There is evidence from the sedimentary record and the oxygen-isotope systematics of cherts that ocean temperatures may have been near the boiling point in the Archaean. We suggest that the early Earth was in a 'greenhouse" condition maintained by volcanic processes. It seems logical that the first oceans may have boiled off producing a heavy aqueous atmosphere. Efficient quenching and partial preservation of crust may have only commenced with the formation of a significant hydrosphere. 4
INTRODUCTION Evensen (1980) has drawn attention to the possible large-scale melting of the moon caused by the energy of accretion. The melting may have produced an outer molten layer termed a magma ocean. If such a magma ocean were produced on the moon then for a larger body such as the Earth the degree of melting must have been even larger. This concept would lead to two possibilities: either a hydrosphere was present and quenched early crust that was rapidly recycled, or the hydrosphere was volatilized and the oceantemperature buffer was removed, leading to high surface temperatures as on Venus. Anders & Owen (1977) and Sagan & Mullen (1972) have discussed possible early 'greenhouse" conditions for the early Earth and Mars. In this paper we draw attention to the common presence of talcrich sediments in the Archaean and indicate that widespread formation of talc may reflect massive injection of silica into the oceans by hydrothermal processes. The same process would cause ejection of carbon dioxide into the atmosphere to maintain the 'greenhouse'' effect. 4
4
ARCHAEAN SEDIMENTARY TALC We have recently studied talc-rich sediments in the Matagami mining district of Quebec, Canada: a more detailed description of the alteration associated with this deposit is in press (Costa etal., 1980). The talc-rich sediments form part of the gangue associated with an exhalative, massive sulphide, ore body (Sharpe, 1968; Goodwin & Ridler, 1970; Roberts, 1975). In parts of this ore body there are virtually monomineralic, con-
Spec. Pubis geol. Soc. Aust., 7 (1981)
tinuous bodies of talc exhibiting sedimentary structures. Geochemistry of the talc bodies clearly shows that they are not replacing maficultramafic rocks ( e.g. they are low in Cr and Ni) or adjacent igneous rocks (see Table I). A survey of the literature shows that talc is common in ancient hydrothermal sedimentary deposits and has even been reported from modern ocean-floor hydrothermal systems as in the Guaymas Basin (Lonsdale, 1978). We have concluded a preliminary study of 180/ 160 relations within silicate phases of the talc-rich sediments and the altered zones beneath the ore. The <5 0 of coexisting quartz and chlorite in five samples from the footwall average + 8.5%o and 0.5 %>o respectively. The quartzchlorite fractionation yields an isotopic temperature of about 300 °C (Wenner & Taylor, 1971; Clayton et al., 1972) and a calculated <5*80 of 0.5 ± 0.5°/oo for fluids discharging through the footwall and ore body. Fluids implicated in mineralization are therefore isotopically similar to present-day ocean water. Oxygen-isotope fractionation for quartz-actinolite pairs within the ore body yield similar temperatures. The <5 O of talc ranges from +5.2 to +6.6%o and we have similar oxygen-isotope data from talc-bearing rocks near the vents of other seafloor discharge systems (Agnico-Eagle, Quebec, 1.8 Ga; Cape Rosier Mine, Maine, Silurian). The field observations and chemical and oxygen-isotope data support a simple model in which sea water at a temperature near 300 °C was discharged through rhyolites carrying Si-Fe-Mg and the ore metals. Chemical sediments including talc have precipitated and hosted deposition of 18
ls
U. R. COSTA, W. S. FYFE, R. KERRICH & H. W. NESBITT
454
TABLE I
Geochemistry of talc bodies and associated rocks 1
2
3
4
5
6
72.86
50.24
25.64
59-70
38.62
11.30
0.28
0.53
ND
ND
ND
Samp]e Si0 2 Ti0 2
0.27
Al 2 0 3 j.
11.02
13.13
22.21
0.05
0.17
0.16
Fe203
5.27
22.71
24.50
10.99
38.49 '
59-19
MnO
0.16
0.26
0.28
0.11
0.41
0.33
MgO
5.99
0.44
7.58
17.29
25.32
15.42
CaO
1.32
ND
ND
ND
0.96
Na20
5.21
ND
ND
ND
ND
ND
K20
0.66
ND
0. 18
ND
ND
ND
S
0.02
0.04
0.02
0.04
4.59
18.21
ZnO
NO
0.50
0.09
ND
1.21
0.21
CuO
ND
0. 13
ND
ND
0.13
L0I
1.9-2
5.43
9-54
3.96
*
99-15
100.30
100.28
100.17
100.00
Total S.G.
2.70
2.78
2.85
3.23
1.38
100.00
2.85
3.13
3.63
Zr
1*98
512
952
18
20
34
Y
218
234
451
28
31
36
Nb
46
45
81
10
10
17
Cr
2
2
10
ND
7
6
Ni
ND
ND
ND
ND
ND
ND
Sr
54
18
22
20
30
44
Rb
12
ND
11
Ba
146
ND
31
ND
4
ND
Zr/Al
85
74
81
679
222
400
Zr/Y
2.3
2.2
- 2.1
7
0.6
13
0.6
11
3-2
3.3
3.3
-
-
-
7.9
6.5
6.9
377
111
200
Zr/Nb
10.8
11.4
1 1.8
2.0
-2
0.9
Ti/Zr xio" Nb/Al
1.8
-4 Styaq.
LOG a
2.0
1 Parent unde r1y i ng rhyoli t ic rock; Moderately and highly altered rhyolitic rock; 2, 3 4, 5, 6 - Pure talc and su1phide-beariing samples from ta Ic-rich mass. Not detected. Recalculated on an L01-free basis (sul phur anallysed separately)
ND *
ores as hot fluids discharged through the sediment-throttled vent region. The formation of talc (or a sepiolite precursor) in ocean water enriched in silica is not surprising. The stability relations of some common Mg-Si phases are shown in Figure 1. The regions A and B represent the ranges of composition of modern surface sea water and interstitial waters in sediments (Wollast et al., 1968). Both water types are saturated or supersaturated with respect to sepiolite, talc and serpentine (chrysotile). The same waters are undersaturated to supersaturated with respect to quartz but undersaturated with respect to amorphous silica. Thus all the above hydrated magnesium silicates could be stable in modern marine sediments and their formation will depend on kinetic factors. It has been shown that sepiolite is readily precipitated (Fig. 2) at ambient temperatures and pressures. While talc or serpentine have not been synthesized at room temperature they are readily
Fig.
Solubility relations of magnesian silicates in aqueous solutions (25 °C, 1 bar total pressure). Solutions plotting above curve 1 will precipitate sepiolite, above curve 2 a serpentine mineral (chrysotile). Curves 3 and 4 illustrate talc and amorphous silica saturation. Waters plotting in the solution field should dissolve all of the above minerals. Area A represents the range of conditions in surface sea water and area B the range in sea water trapped in modern marine sediments.
formed in moderate-temperature hydrothermal experiments and have been observed to form at and above 90 °C from other precursors (Bricker et al., 1973). It seems that if volcanic discharge caused local increases of silica activity or the Mg + + / H + ratio, phases like sepiolite and talc could form rapidly and sepiolite would be rapidly replaced by talc in warm silica-bearing solutions. During volcanic discharge of 300 °C water into ocean water, silica may be carried at the 0.1% level (Ellis & Mahon, 1977). Magnesium concentration is highly variable but generally at low values (0.001-l.Oppm) presumably because it tends to be fixed if it passes through quartz-rich rocks. Thus during talc or sepiolite formation any major perturbation in sea water composition will be silica increase which alone will have little influence on pH.
FIRST PRESERVATION OF CRUST
1
y 1
—1—
"If
1
455
under these conditions could move 6 x 10 g a~ of silica to the oceans and form magnesiumrich sediments at a rate of 1 k m a - . Since 1958, C 0 in the atmosphere has increased by about 5% due to human activity which adds about 5 x 10 g of C 0 to the atmosphere per year (Woodwell, 1978). The past rate of volcanic emission is thus comparable. Present data have shown that the <5 0 of marine cherts has changed systematically with time. Thus at Isua cherts have maximum values of 20.4%o (Perry et al., 1978) while modern cherts have values up to 36%o (Knauth & Epstein, 1975). A corresponding change is reported for marine carbonates (Schidlowski et al., 1975; Veizer & Hoefs, 1976). Knauth & Lowe (1978) have discussed the various hypotheses that could account for such variation but, given that the <5 0 of sea water is buffered by exchange with sea-floor basalts (Muehlenbachs & Clayton, 1976; Beaty & Taylor, 1980) and that there is little evidence for much change in <5 0 of sea water with time, the most plausible explanation at present is that ocean temperatures have slowly cooled from about 80 °C at 3.8 Ga (Knauth & Lowe, 1978; Knauth & Epstein, 1975). Other evidence for high initial surface temperatures includes: (i) The palaeontological record, which suggests that the order of emergence of organisms correlates with their temperature tolerance (Hoyle & Wickramasinghe, 1979). (ii) Isotopic data for a uniform <5D, <5 O of ocean water and warmer Archaean climates (Taylor & Magaritz, 1975). (iii) Theoretical modelling of atmospheric greenhouse effects are compatible with the proposed thermal history (Sagan & Mullen, 1972; Muehlenbachs & Clayton, 1976). (iv) Sedimentological evidence favours relatively warm conditions during formation of the Onverwacht Group (Knauth & Lowe, 1978). We conclude that the evidence for warmer past oceans is compatible with the widespread formation of magnesium-rich sediments. It is possible that igneous hydrothermal events could produce a significant flux of C 0 into the atmosphere, particularly as the solubility of C 0 would be much less in hot sea water. If this flux were sufficient to maintain a ''greenhouse" effect throughout much of the Archaean period it clearly has important consequences with regard to man's present activities (Woodwell, 1978). Finally, if ocean temperatures were near boiling at 3.8Ga, then it is quite reasonable that during the early periods of intense meteoritic bombardment that total volatilization would 14
1
3
1
2
\
15
A B
18
T" 1
Vl\
1
2
1
2
•
LOG o c n StO^aq.
18
Fig. 2. Maximum range of pH and silica activities that can occur without significantly changing the mineralogical compositions of marine sediments. All calculations were made using present major element sea water composition (25 °C, 1 bar pressure). Curve 1 gibbsite-illite stability relations, curve 2 represents amorphous silica saturation. Curve 3 represents chlorite-kaolinite stability relations, curve 4 represents enstatite- solution stability relations. Areas A and B as in Figure 1.
If talc or sepiolite is precipitated, reactions which may occur are,respectively: 3Mg
ls
+ 6HCO3- + 4 S i 0 - Mg3Si O (OH) + 6C0 + 2H 0 2 M g + 3Si0 -f 4HC0 " Mg Si 0 (H 0) + 4C0 In both reactions, the net result of the fixation of sea water magnesium in talc or sepiolite is the liberation of carbon dioxide which may eventually be transferred to the atmosphere. This C 0 evolution will be enhanced by C 0 in the discharge phase itself which may be at the lOOppm level (Ellis & Mahon, 1977). Wolery & Sleep (1976) have calculated the modern fluid flux near ocean ridges and estimate that the mass flux of water of 300 °C could attain 1.3 x 10 g a - . As past heat production must have been several times greater, if one assumes that volcanism was roughly proportional to global heat production and that volcanism was dominantly submarine, the flux in the Archaean could have been in the order of five times the present, i.e. 6 x 10 g a - . A fluid evolved + +
2
4
2
2
3
8
2
2
+ +
2
10
2
3
2
2
2
2
17
1
17
1
18
2
2
U. R. COSTA, W. S. FYFE, R. KERRICH & H. W. NESBITT
456
occur. Given the resulting, much higher, surface temperatures and large-scale melting it should be no surprise that the early crust was not preserved. However, with the first production of a significant permanent hydrosphere, the water-cooled surface might lead to sufficient stabilization to preserve lighter materials.
ACKNOWLEDGMENTS We gratefully acknowledge the support of N.S.E.R.C. Canada; C.I.D.A.; the Institute of Geosciences, Salvador; and frequent discussions with R. W. Hutchinson.
REFERENCES
ANDERS, E., & OWEN, T., 1977: Mars and Earth: origin
and abundance of volatiles. Sci., 198, 453-465.
BEATY, D . W., & TAYLOR, H . P . , 1980: Early crustal
hydrothermal processes and the 0 / 0 evolution of sea-water; evidence from the Amulet mine, Noranda, Quebec (abstract). EOS, 61, 386. 1 8
1 6
BRICKER, O . P . , NESBITT, H . W . , & GUNTER, W . D . ,
1973: The stability of talc. Am. Mineral.,
58,
64-72. CLAYTON, R . N . , O'NEIL, J . , & MAYEDA, T . K . , 1972:
Oxygen isotope exchange between quartz and
water. J. geophys. Res., 77, 3057-3067. COSTA, U . R . , FYFE, W . S., KERRICH, R . , & NESBITT,
H. W., 1980: Archaean hydrothermal talc and evidence for high ocean temperatures. Chem.
Geol., 30, 341-349.
ELLIS, A . J . , & MAHON, W . A . J . , 1977: Chemistry and
Geothermal Systems. Academic Press, London.
EVENSEN, N. M., 1980: Lunar differentiation by frac-
tional crystallization and partial melting (abstract).
EOS, 61, 33.
GOODWIN, A. M., & RIDLER, R . H . , 1970: The Abitibi
Orogenic Belt; in Basins and Geosynclines of the
Canadian Shield. Pap. geol. Surv. Can., 70-40, 1-30.
HOYLE, F . , & WICKRAMASINGHE, 1979: Life
Cloud.
Sphere Books, London. KNAUTH, L. P., & EPSTEIN, S., 1975: Hydrogen and oxygen isotope ratios in silica from the Joides deep
sea drilling project. Earth planet. Sci. Lett., 25, 1-10.
KNAUTH, L. P . , & LOWE, D . R . , 1978: Oxygen isotope
geochemistry of cherts from the Onverwacht Group (3.4 billion years), Transvaal, South Africa, with implications for secular variations in the isotopic composition of cherts. Earth planet. Sci.
Lett., 41, 209-222. LONSDALE, P . , 1978:
Submersible exploration of Guaymas Basin: a preliminary report of the Gulf of California 1977 operation of DSV-4 Seacliff.
Scripps Inst. OceanogrRef. Ser., 79-1, 16.
MUEHLENBACHS, K., & CLAYTON, R. N., 1976: Oxygen
isotope composition of the oceanic crust and its bearing on seawater. J. geophys. Res., 81, 4365-4369.
PERRY, E . C . , AHMAD, S. N . , & SWULIUS, T . M . , 1978: The oxygen isotope composition of 3800 m.y. old
metamorphosed chert and iron formation from Isukasia, West Greenland. J. Geol., 86, 223-239. ROBERTS, R. G . , 1975: The geological setting of the Mattagami Lake mine, Quebec: a volcanogenic massive sulfide deposit. Econ. Geol., 70, 115-129. SAGAN, C . , & MULLEN, G . , 1972: Earth and Mars: evolution of atmospheres and surface temperatures. Sci., 177, 52-56. SCHIDLOWSKI, M . , ELCHMANN, R . , & JUNGE, C . E.,
1975: Precambrian sedimentary carbonates: carbon and oxygen isotope geochemistry and implications for the terrestrial oxygen budget. Precamb. Res., 2, 1-69. SHARPE, J. I., 1968: Geology and sulfide deposits of the Matagami area, Abitibi-East County. Geol. Rep.,
Quebec. Dep. Nat. Res., 137.
TAYLOR, H . P . , & MAGARITZ, M . , 1975: Oxygen and hydrogen isotope studies of 2 . 6 - 3 . 4 billion year old
granites from the Barberton Mountain Land, Swaziland, and the Rhodesian Craton, Southern
Africa. Geol. Soc. Am. Abstr. Progr., 7, 1293.
VEIZER, J., & HOEFS, J., 1976: The nature of Oi8/Oi«
and C / C secular trends in sedimentary carbon1 3
1 2
ate rocks. Geochim. cosmochim. Acta, 40,
1387-1395. WENNER, D. B., & TAYLOR, H. P., 1971: Temperatures of serpentinization of ultramafic rocks based on 0 /0 fractionation between coexisting serpen18
16
tine and magnetite. Contr. Mineral. Petrol., 32, 165-185.
WOLERY, T . J . , & SLEEP, N . H., 1976: Hydrothermal
circulation and geochemical flux at mid-ocean ridges.J. Geol., 84, 249-275.
WOLLAST, R . , MACKENZIE, F . T . , & BRICKER, O . P . ,
1968: Experimental precipitation and genesis of sepiolite at earth-surface conditions. Am. Mineral., 53, 1645-1662. WOODWELL, G . M . , 1978: The carbon dioxide question. Sci. Am.,
138, 34-43.
MAGMATIC CYCLES AND THE EVOLUTION OF THE ARCHAEAN GRANITIC CRUST IN THE EASTERN TRANSVAAL AND SWAZILAND C. R. Anhaeusser & L. J. Robb Economic Geology Research Unit, University of the Witwatersrand, 1 Jan Smuts Avenue, Johannesburg 2001, South Africa
ABSTRACT Three magmatic cycles are recognized in the Archaean granitic crust of the eastern Transvaal and Swaziland in terms of their geochemical, geochronological and field characteristics. The earliest cycle commenced approximately 3500 Ma ago and involved the formation of soda-rich tonalites and trondhjemites and a complex series of bimodal gneisses and migmatites. The second cycle began approximately 3200 Ma ago with the emplacement of multi-component potash-rich batholiths that enveloped vast areas occupied by the gneisses and migmatites of the earlier magmatic event. The third cycle was initiated with the intrusion of a number of late, mainly potash-rich, granite and syenite plutons, approximately 2900 Ma ago. The distinctive physical, chemical and isotopic characteristics of the components of each cycle indicate that the subdivision of the various granitic rocks in the area into three magmatic cycles provides a useful conceptual framework within which to view the evolution of the Archaean crust in the region.
INTRODUCTION The granitic terrain flanking the Barberton greenstone belt in the eastern Transvaal has been critically re-examined over the past five years as part of South Africa's contribution to the International Geodynamics Programme. The investigations have involved regional as well as detailed field mapping, supplemented by geochemical, geochronological and structural studies, all of which have been undertaken in an attempt to decipher the evolutionary history of the Archaean crust in the area. Earlier studies in the eastern Transvaal and Swaziland led to the recognition of a wide range of granitic rock types which are portrayed on various maps accompanying attempts aimed at classifying the granitic rocks of the region (Anhaeusser et al., 1968; Hunter, 1957, 1970, 1973, 1979; Viljoen & Viljoen, 1969a, b; Visser et al., 1956). Because of the nature and complexity of the granitic terrain no consensus has yet been achieved on the correlation of these rocks, and there are widely divergent views on the evolution of the primitive crust (Anhaeusser, 1973, 1981; Condie & Hunter, 1976; Glikson, 1976, 1979; Hunter, 1970, 1973, 1979; Hunter et al., 1978; Viljoen & Viljoen, 1969b). Experience gained in the Geodynamics study strip in the eastern Transvaal, coupled with reconnaissance investigations in parts of Swazi-
Spec. Pubis geol. Soc. Aust., 7 (1981)
land, has prompted this reassessment of the granitic terrain. A three-fold subdivision is proposed whereby the various granitic components are grouped together under the general heading of magmatic cycles, erected on the basis of similarities in field relations, rock geochemistry and isotopic ages. Each magmatic cycle embraces a wide range of events and granitic rock types with unifying characteristics that formed by processes operating throughout the time span defining the magmatic cycle. No attempt is made here to justify fully grouping the granitic rocks into the scheme outlined in the following sections: this justification will be made in later publications of results of the Geodynamics research programme. GEOLOGIC SETTING To the east of the Transvaal Drakensberg escarpment lies the Archaean granite-greenstone terrain of the eastern Transvaal and adjoining territory of Swaziland (Fig. 1). Centrally situated with respect to Figure 1 is the Barberton greenstone belt which consists of an assemblage of metamorphosed volcano-sedimentary rocks, including mafic and ultramafic lavas (basaltic and peridotitic komatiite, high-Mg basalt, tholeiite), mafic to felsic volcanics and pyroclastics (calcalkaline series basalt-dacite-rhyodacite) and sediments (greywacke, shale, banded iron-formation,
458
C. R. ANHAEUSSER & L. J. ROBB
THE GRANITIC ROCKS
YOUNGER SEDIMENTARY COVER SEQUENCES i
OF THE
USUSHWANA COMPLEX
c
Na-RICH PLUTON
b
YOUNGER GRANITE PLUTONS
a
OLDER GRANITE PLUTONS
2a
HOMOGENEOUS. OFTEN PORPHYRITIC
-K
GRANITES/ ADAMELUTES
+
m*L
a
+ +4 2b
+
+
+ +
TRONDHJEMITES b
SWAZILAND
+
LEUCOCRATIC BIOTITE
m
AND
+
+
b MARGINAL K-RICH MIGMATITES/GNEISSES
+
EASTERN TRANSVAAL
+
+
HORNBLENDE TONAUTES GNEISSES
N^LSPRUIT W
• • • + »+ 2b +
SWAZILAND SUPERGROUP
+j
- r < 7 X
MIGMATITES/BIMODAL
+ + + + + + + J-
+
9S69W
^ ^ ^ ^ B r a w
i
/JHSI
INTERNATIONAL BOUNDARY
Fig. 1.
Simplified geological map of the eastern Transvaal and Swaziland showing a threefold subdivision of the granitic rocks of the region.
MAGMATIC CYCLES chert, quartzite, conglomerate), described, for example, by Anhaeusser (1978), Viljoen & Viljoen (1969c, d) and Visser et al. (1956). Of particular significance are the numerous greenstone enclaves in the granitic rocks surrounding the Barberton Mountain Land. Best preserved are the xenolithic remnants in the Badplaas area, which have been traced for over 50 km beyond the southern limits of the main greenstone mass (Anhaeusser & Robb, 1980, and Fig. 1). Studies have confirmed the observations of Viljoen & Viljoen (1969c, d) that the scattered greenstone enclaves in the gneisses consist entirely of assemblages characteristic of the lower formations of the Onverwacht Group of the Swaziland succession. Detailed mapping has shown, furthermore, that these enclaves can be traced from the gneissic terrain directly into the Barberton greenstone belt in several areas. The Barberton volcanic sequences have yielded a precise Sm-Nd age of 3540 ± 30 Ma (Hamilton et al., 1979) and all available evidence suggests that these rocks were intruded by tonalite and trondhjemite ranging in age from approximately 3450-2900Ma (Barton, 1981; Oosthuyzen, 1970). Localized zones of migmatite in areas immediately flanking some greenstone enclaves (e.g. southeast of Badplaas: Anhaeusser & Robb, 1980) are interpreted as forming by the intrusion and interaction of the granitic rocks with the greenstone xenoliths. Investigations in Swaziland (Jackson, 1979; A. C. Wilson, pers. comm., 1979) have revealed numerous greenstone enclaves between the Komati River in the north and Mankaiana in the south (Fig. 1). These rocks are similar to the meta-volcanics and meta-sediments of the Onverwacht sequence in the Barberton greenstone belt (Barton et al., 1980; Jackson, 1979). Known locally as the Dwalile metamorphic suite, the successions consist mainly of mafic and ultramafic lavas and subordinate layered intrusions, overlain by chemical and clastic sediments. The Dwalile supracrustal remnants are enveloped by migmatites and a variety of tonalitic and trondhjemitic orthogneisses, and appear to be justifiable grounds for correlating the Mankaiana and Badplaas areas rather than regarding the two environments as having entirely separate origins. A major line of disagreement centres around the nature of the early crust in the BarbertonSwaziland region. One view is that the siliceous low-K leucocratic gneisses (the Bimodal Suite of Hunter, 1974a, and Hunter et al1978) predate the Onverwacht G r o u p greenstones and may represent a basement to the latter. The alternative is that orthogneisses and migmatites making up
459
the Ancient Gneiss Complex in Swaziland (the latter composed of the Bimodal Suite, the homogeneous Tsawela tonalitic gneisses and the Dwalile and Mkhondo metamorphic suites— Barton et al., 1980; Hunter, 1970, 1974a) are younger than the basal units of the Onverwacht Group and were derived by partial melting of an ensimatic source (Anhaeusser, 1973; Viljoen & Viljoen, 1969a). Rb-Sr isotopic studies by Barton et al. (1980) on orthogneisses of the Ancient Gneiss Complex in northeastern Swaziland yield an age of 3555 ± 111 Ma (2a) which is indistinguishable f r o m the Sm-Nd age of the lower Onverwacht volcanic rocks. These findings are consistent with the view, favoured here, that the orthogneiss component of the Ancient Gneiss Complex is coeval with, and was possibly derived f r o m , rocks compositionally similar to those of the basal members of the Barberton greenstone belt. The low initial 87 S r / 8 6 S r ratio of 0.6999 ± 0.0016 (2a) reported by Barton et al. (1980) further precludes any other significant prehistory for the Swaziland orthogneisses. In terms of the approach adopted in this paper the ~ 3550 Ma age represents the derivation or emplacement age and hence, the commencement of granitic magmatic activity of the first magmatic cycle. The low initial ratios indicate that the leucocratic gneisses of the Bimodal Suite cannot be a much older basement to the Onverwacht rocks but rather are the partial melt products of ensimatic parents. This type of parental material abounds in the nearby Barberton greenstone belt and in the related greenstone enclaves in the surrounding granitic terrain. The second magmatic cycle commenced approximately 3200 Ma ago when K-rich granitic magma was emplaced as large batholithic bodies both north and south of the Barberton greenstone belt (Fig. 1). Linked with this m a j o r crustforming episode were supplementary intrusive phases whose origins appear related to processes responsible for the development of the batholithic magmas. The potassic batholiths are rimmed by migmatite and gneiss aureoles which reflect zones of interaction between the batholiths and preexisting crust. Within these zones, which are best developed around the Nelspruit batholith in the north and the Mpuluzi batholith south of the Barberton greenstone belt, the earlier formed Narich gneisses, migmatites and greenstone remnants were variably influenced by processes of granitization and metasomatism. All the events ascribed to the second magmatic cycle appear to have terminated approximately 3000Ma ago, providing the crustal stabilization necessary for
460 C . R . A N H A E U S S E R & L. J. R O B B the subsequent development of the early Protero- FIRST MAGMATIC CYCLE zoie cratonic basins (Anhaeusser, 1973; Hunter, Migmatites and bimodal gneisses 1974ft). Migmatites and/or bimodal gneisses are best The third magmatic cycle began approximately 2900 Ma ago when numerous discrete granitic developed immediately southwest of the Barberton greenstone belt (Robb, 1981) and in the typeand syenitic bodies were emplaced into an already consolidated, technically stable, crustal regime. area of the Ancient Gneiss Complex in the Mankaiana district in Swaziland (Jackson, 1979). In These high-level, generally K-rich, plutons are clearly transgressive and cluster into two groups both areas, the terrain is underlain dominantly by defining a set of Older Plutons (ca 2900 Ma old) tonalite or trondhjemite gneisses (column 1, and another of Younger Plutons (<ca 2600 Ma Table I) intimately associated with subordinate amphibolites together with variable but comold). The Usushwana Igneous Complex, dated at monly significant proportions of anatectic approximately 2810 Ma (Davies et al., 1970; age material. The migmatites are invariably associcorrected using decay constant 1.42 x 1 0 for ated with greenstone supracrustal remnants Rb), overlaps with the third magmatic cycle but correlated directly with the Barberton greenstone as these rocks consist mainly of pyroxenite, gab- belt (e.g. in the region southeast of Badplaas— bro and granophyre they will not be considered Anhaeusser & Robb, 1980; Anhaeusser, 1980) or forming part of the Dwalile metamorphic suite in further. Lastly, it is emphasized that the three mag- the Mankaiana region. The migmatites in both regions (Fig. 1, see distribution of la) have been matic cycles are most clearly defined by their distinctive physical (field) and geochemical charac- repeatedly deformed so that commonly the relationships between greenstone remnants and teristics supplemented by confirmatory Rb-Sr isotopic data. The earliest event coincided with gneisses are obscure. Also characteristic of both the onset of sodic igneous activity approximately areas are syntectonic mafic dykes which probably 3550 Ma ago. This was followed by potassic ig- represent an intrusive episode that was coeval neous activity 3200 Ma ago and finally by late with the Upper Onverwacht (Geluk Subgroup) stage potassic and, to a lesser extent, sodic ig- (Jackson, 1979; Robb, 1981). neous activity 2900 Ma ago. -11
87
TABLE I
A verage major element compositions of granitic rocks representative of the three magmatic cycles in the Eastern Transvaal and Swaziland Col umn
1
2
3
4
5
6
7
8
9
10
Map Reference
la
lb
lc
2a
2a
2a
2b
3a
3b
3c
Si02
71 .16
65-05
70.67
69-72
67.70
72.38
67.59
70.42
71.20
68.97
Ti02
0
0.49
0.34
0.32
0.42
0.26
0.32
0.35
0.25
0.36
A1203
14,.84
15.75
15.75
15-52
15.83
13-95
15-59
14.79
13.73
15.80
Fe203
0 • 77
4.40"
2.42"
2 . 18*
2.83*
0.85
2.39*
0.88
1.16
2.52*
FeO
1 •, 52
-
-
-
-
1.40
-
1.38
1.8)
-
MnO
0,.02
0.06
0.06
0.05
0.05
0.09
0.05
0.05
0.10
0.06
MgO
0,• 95
2.51
1.11
0.95
1.72
0.51
1.29
0.91
0.77
1.61
CaO
3..18
4.47
3-11
2.92
2.41
1.33
2.29
1.82
1.83
2.72
Na20
4,.82
5.18
4.25
4.31
4.22
3-76
3.95
~ 4.65
3.38
4.98
K20
1..65
1-52
1.55
3.65
3.18
4.67
3-93
3.61
5.10
2.13
0..12
0.20
0.08
0.04
0.11
0.23
0.08
0.14
0.11
0.01
0.• 75
-
-
-
-
0.68
-
0.70
0.62
P2O5 H20
+
H2O~
0 . 07
L0 It
i
No. of Analyses
-
-
-
0.08
-
0.13
0.07
-
1 .01
0.48
0.59
-
0.55
-
-
0.45
41
7
18
9
8
9
3
4
9
Total iron as Fe2C>3 Loss on ignition
I . Siliceous g n e i s s e s and m i g m a t i t e s - Bimodal S u i t e ( H u n t e r , 1973)
2. 3.
A.
5. 6. 7. 8. 9-
10.
_
-
1.01
H o r n b l e n d e (± b i o t i t e ) t o n a l i t e gneiss - Kaap V a l l e y Pluton (Robb, u n p u b l i s h e d data) L e u c o - b i o t I t e t r o n d h j e m i t e gneiss - R o o i h o o g t e Pass Area ( A n h a e u s s e r & Robb H o m o g e n e o u s coarse p o r p h y r i t i c granite - N e l s p r u i t P o r p h y r i t i c G r a n i t e (Robb 1978) H o m o g e n e o u s m e d i u m - g r a i n e d , grey g r a n o d i o r i t e - Hebron G r a n o d i o r i t e (Robb 1978) Homogeneous m e d i u m - g r a i n e d , p i n k i s h grey a d a m e l l i t e - H o m o g e n e o u s Hood G r a n i t e (V M a r g i n a l K-rich gneisses and m i g m a t i t e s - N e l s p r u i t M i g m a t i t e and Gneiss T e r r a i n ( j o e n & V i l j o e n , ( R o b b , 1978)-. Late g r a n i t e plutons - O l d e r Plutons (Viljoen 6 V i l j o e n , 19692?). Late granite plutons - Y o u n g e r Plutons (Viljoen & V i l j o e n , 1969Jb). N a - r i e h pluton - Cunning M o o r T o n a l i t e (Robb, 1978).
IQRQ)
1969*0.
MAGMATIC CYCLES The interpretation of migmatites southwest of the Barberton greenstone belt is markedly different from that for migmatites or bimodal gneisses of the Mankaiana region in spite of the apparent similarities referred to above. In the Badplaas area, three types of migmatite are recognized (Robb, 1981). The first type occurs where tonalite/trondhjemite gneisses intruded and migmatized portions of recognizable greenstone crust. The second type occurs where migmatization results from the intrusion of significant proportions of anatectite and obscures relations between trondhjemite and amphibolite. The third migmatite variety developed where tholeiitic dykes intruded pre-existing tonalite/trondhjemite gneiss and were subsequently deformed and migmatized. No unequivocal relationships point to primitive komatiitic basalts that are largely correctable with the lower greenstone successions, post-dating the earliest recognized sial in the region. The syntectonic tholeiite dykes cannot be shown to have fed the greenstone successions and this, together with available geochronology (Barton, 1981) suggests that the sialic crust intruded by the dykes was not a basement to the supracrustal greenstone successions. The bimodal gneisses and migmatites in the Mankaiana area are, however, considered by Hunter (1974a, 1979) to pre-date the Barberton greenstone belt although no direct evidence for this relationship exists and the oldest age for these rocks (i.e. 3555 ± 11'lMa, Barton, 1981) suggests that they were, at best, coeval with the earliest developed ensimatic crust. The locally developed Dwalile metamorphic suite is considered to have been deposited above basement consisting of the Bimodal Suite even though repeated high strains have largely obscured relationships between the gneisses/migmatites and the supracrustals (Jackson, 1979). Jackson also suggested that the mafic dykes in the Mankaiana area might have fed an overlying greenstone supracrustal assemblage, but no direct evidence for this is available. Because of the striking similarities between the rocks in the Mankaiana area and those of the Badplaas region (including lithological likenesses between the Onverwacht Group and the Dwalile metamorphic suite and the occurrence of banded and bimodal gneisses in both areas), the authors consider that past attempts to segregate the two terrains genetically are largely unfounded. Robb & Anhaeusser (1979) indicate that differences in metamorphic grade and intensity of deformation may relate to different crustal levels for the two areas (cf. Glikson, 1979; Glikson & Lambert, 1976). The possibility of finding sialic material pre-
461
dating the ensimatic remnants in the bimodal gneiss/migmatite terrains remains, particularly since the discovery of ~ 3800 Ma old basement in the Limpopo Mobile Belt (Barton et al., 1977; Barton, 1981). However, no such material has yet been located in the areas studied here, thereby strengthening the case for primordial simatic crust in the eastern Transvaal and Swaziland. Hornblende tonalites and leucocratic biotite trondhjemites Hornblende tonalites and leucocratic biotite trondhjemites occupy large areas of the granitic terrain in both the eastern Transvaal and Swaziland (see distribution of 1 b and 1 c in Fig. 1). These gneisses are homogeneous and are generally characterized by an S > L fabric (i.e. foliation more pronounced than the lineation) which parallels the greenstone-gneiss contact. The tonalites have higher CaO, total Fe, and MgO contents than the more leucocratic trondhjemites and also have correspondingly lower S1O2 (compare columns 2 and 3, Table I). The trondhjemite gneisses have, furthermore, a very similar majorelement chemistry to the leucocratic components of the bimodal gneisses and migmatites previously described. The tonalite and trondhjemite bodies in the Badplaas region have previously been described as elliptical or rounded, their shapes being outlined by the edges of the Barberton greenstone belt or greenstone septa. The absence of welldefined, elliptical tonalite/trondhjemite gneiss plutons, particularly in the Mankaiana area in Swaziland, has been used by Hunter (1973, 1979) and Hunter et al. (1978) to suggest that the Ancient Gneiss Complex should not be correlated with the tonalitic gneisses in the Barberton region. Recent work southeast of Badplaas indicates, however, that the distribution of numerous greenstone enclaves in the granitic terrain largely invalidates the over-simplified impression that only rounded and well-defined plutons exist in the area (Anhaeusser & Robb, 1980). The few elliptical plutons are confined to the immediate edge of the Barberton greenstone belt, whereas some distance away the shapes of tonalite/ trondhjemite bodies are irregular and are not conspicuously outlined by the scattered greenstone xenoliths. Nevertheless it is still possible to recognize discrete tonalite/trondhjemite bodies or "cells" on the basis of their trace-element contents (particularly Sr). In the Badplaas region, these 4 'cells" have irregular shapes and may incorporate a number of small greenstone enclaves (Robb, 1981). In the type area of the Ancient Gneiss Complex the distribution of numerous small, scattered
462
C. R. ANHAEUSSER & L. J. ROBB
remnants of the Dwalile supracrustals resembles the greenstone enclaves in the Badplaas region. No large, homogeneous tonalite/trondhjemite plutons occur in the Mankaiana region but remapping by the Swaziland Geological Survey (A. C. Wilson, pers. comm., 1979) and Jackson (1979) indicates a number of small elliptical bodies of hornblende-biotite tonalitic gneiss. These complement the large body of hornblende tonalite recorded by Hunter (1966) and subsequently included on other maps of the territory (Hunter, 1970, 1973, 1979). These considerations, together with those of Robb & Anhaeusser (1979), suggest that the tonalite and trondhjemite plutons constitute as much a component of the Ancient Gneiss Complex (including the Granodiorite Suite of Hunter, 1973) as they do the equivalent gneiss/migmatite terrain in the Badplaas region. The tonalite and trondhjemite gneisses, together with the migmatites and bimodal gneisses described above, are thus considered to form the dominant portion of the first magmatic cycle. Geochronological and tectonic characteristics The migmatite and gneiss terrain southwest of the Barberton greenstone belt is characterized by Rb-Sr isochron ages ranging from 3447 ± 168 Ma to 2916 ± 33 Ma with a concomitant range in initial 87 Sr/ 8 6Sr ratios of 0.7000 ± 0.0019 to 0.7018 ± 0.0005 (Barton, 1981). In Swaziland, the migmatitic and gneissic units of the Ancient Gneiss Complex range in age from 3555 ±111 Ma to 3138 ± 112Ma with initial 8 7 Sr/ 8 6 Sr ratios of between 0.6999 ± 0.0016 and 0.7048 ± 0.0022 (Barton, 1981; Davies, 1970; Davies & Allsopp, 1976). In both areas, therefore, the onset of felsic magmatism took place about 3500Ma ago within the limits of geochronological error. Furthermore, the development of this early sialic material was largely coeval with the formation of oceanic crust now represented by the extrusive rocks of the Onverwacht Group. This is supported by Barton et al. (1980) who, on the basis of isotopic data, show that the Barberton volcanic rocks (Komati Formation) are coeval with the leucocratic gneisses of the Bimodal Suite and also have indistinguishable initial 8 7 Sr/ 8 6 Sr ratios. It is evident that the emplacement of units associated with the first magmatic cycle was repeated episodically for - 6 0 0 Ma after the formation of the earliest sial in the area. As Barton (1981) points out, however, the younger - 2 9 0 0 Ma ages may reflect an age of emplacement or a resetting of originally emplaced material that took place some time after the original formation of the magma and during which interval Sr
isotopes were re-homogenized. Slightly higher initial 87 Sr/ 8 6Sr ratios in the younger plutons may also reflect the extended residence of these bodies in the crust before their emplacement, possibly as gravitationally induced diapirs. Consequently, the minimum ages attributed to the first magmatic cycle are not necessarily indicative of an overlap with respect to its processes of formation and the inception of the second magmatic cycle. Tectonic considerations in the Barberton region together with data from the Canadian Archaean (Schwerdtner et al., 1978; Schwerdtner & Lumbers, 1980) indicate that the emplacement of tonalite/trondhjemite gneiss plutons into the greenstone successions was in response to gravitational instabilities that existed because of density contrasts between the two rock types. The formation of primitive sial therefore took place at the same time or shortly before this material was emplaced by diapirism into a higher crustal level than that at which it was generated. This suggests turbulent proto-cratonization and unstable tectonic conditions during the first magmatic cycle. SECOND MAGMATIC CYCLE Multi-component K-rich batholiths Whereas the first magmatic cycle is characterized by Na-rich tonalites and trondhjemites, the second cycle invariably involves more potashrich rocks that are distinctive in terms of their mineralogy, texture, field appearance, age and style of emplacement. Unlike the various units of the first cycle, which were largely emplaced into their present positions by gravity inversion, the different phases of the second magmatic cycle were probably emplaced directly into their present positions by 4 'passive" processes (Pitcher, 1979) involving stoping, cauldron subsidence and assimilation. They are referred to here as batholiths because of this distinction and the size and discrete nature of these multicomponent bodies (Fig. 1). Outlined in Figure 1 are three large K-rich batholiths, two of which exceed 1500 km 2 in area. Most detail is available from the extensive Nelspruit batholith to the north of the Barberton greenstone belt (Fig. 1, 2a) which consists dominantly of coarse-grained, relatively homogeneous, porphyritic granite or adamellite known locally as the Nelspruit Porphyritic Granite (Robb, 1978): its average composition is given in Table I, column 4. It is evident from the chemical composition that these rocks are adamellitic. The Nelspruit Porphyritic Granite is characterized by microcline megacrysts, whose development varies from intense, in certain central areas, to moder-
MAGMATIC CYCLES
ate in more marginal areas. This very extensive phase is considered to have been emplaced at a relatively high crustal level and to have slowly crystallized by a process of inward nucleation such that a pronounced fractionation trend is evident over the body (McCarthy & Robb, 1978). An isotopically coeval phase, known as the Hebron Granodiorite, occurs within the Nelspruit batholith both as veins intruding the Nelspruit Porphyritic Granite and as a small medium- to fine-grained, homogeneous pluton in the centre of the batholith. The granodioritic phase (Table I, column 5) is genetically related to the Nelspruit Porphyritic Granite but intruded the latter shortly after its emplacement (Robb, 1978). A similar intrusive granodioritic phase also forms one component of the large potash-rich Mpuluzi batholith, some 30 km to the southwest of the Barberton greenstone belt (Fig. 1). The Mpuluzi batholith, like its Nelspruit counterpart, consists principally of a coarsegrained, homogeneous, porphyritic granite. It has been invaded by linear, commonly dyke-like bodies of homogeneous, medium-grained, pinkish-grey adamellite (Table I, column 6) referred to previously as either the Homogeneous Hood or Lochiel granite (Hunter, 1973, 1974a; Viljoen & Viljoen, 1969a, b). The recent, more detailed, mapping has shown that this granite represents only a single phase within a much larger multi-component batholithic system. Marginal K-rich migmatiies and gneisses The batholiths of the second magmatic cycle contain at least three discrete and distinctive cogenetic and broadly coeval phases. In addition, the areas marginal to these batholiths are characterized by K-rich migmatites and gneisses that represent zones of interaction between the homogeneous, passively emplaced magmas of the main massifs and the crust that existed during the final stages in the evolution of the first magmatic cycle. The marginal zones are regarded as a fourth component associated with the batholiths and are best developed in the Crocodile River valley and in areas rimming the Mpuluzi batholith south of the Barberton greenstone belt and in Swaziland (Fig. 1,2b). The migmatite zones vary considerably in extent, ranging from only a few hundred metres to over 10 km wide. Proceeding from the batholith margins, which seldom display sharp contacts, there is a gradational decrease in porphyritic phases. Nebulites, consisting of migmatites with ghost-like relics of pre-existing rocks, give way to zones in which enclaves are increasingly abundant and where greenstone components ultimately become clearly recognizable.
463
Similar features are recorded in areas flanking granitoid batholiths in the Superior Province of the Canadian Shield (W. M. Schwerdtner, pers. comm., 1979; Schwerdtner & Lumbers, 1980). The migmatites (arterites) of the marginal zones were formed by the injection of "new" magma related to the K-rich batholiths. Numerous anastomosing pegmatite, granite and aplite dykes and veins add to the complexities of the batholith rims and in turn are responsible for promoting localized zones of metasomatism, palingenesis and granitization. The passive invasion process allowed preservation of pre-existing planar mineral fabrics and minor structures such that the batholith aureoles appear to have inherited, locally, the imprint of the structural (and lithologic) regime of the area before the onset of the second magmatic cycle. Chemically, the rocks in the marginal areas vary considerably and no single analysis could typify the characteristics of the zone. The average composition derived from nine samples of the Nelspruit migmatites and gneisses listed in Table I, column 7, should be viewed in such light, although it nevertheless exemplifies the dominant granodioritic nature of most rocks developed in these areas. Geochronological and tectonic characteristics Although the batholiths of the second magmatic cycle are made up of a number of compositionally and texturally discrete phases, geochronological data suggest that the rocks were emplaced into the crust over a shorter period than the sialic units of the first cycle. The various phases of the Nelspruit batholith were apparently formed between 3211 ± 133 Ma and 3149 ± 125 Ma ago with accompanying peripheral (mainly pegmatitic) phases being dated at 2927 ± 137 Ma (Barton, 1981; deGasparis, 1967). Initial Sr/ *Sr ratios in this suite of rocks range from 0.7007 ± 0.0011 to 0.7052 ± 0.0020. Data from the Mpuluzi batholith indicate a range in Rb/Sr isochron ages from 3028 ± 14 Ma to 2986 ± 69 Ma with a range in initial Sr/ Sr ratios of 0.7013 ± 0.0002 to 0.7054 ± 0.0036 (Barton, 1981; Davies, 1970). The onset of magmatism in the second cycle, therefore, apparently occurred about 3200 Ma ago and continued for some 300 million years. The initial S r / S r ratios of most associated phases are low, suggesting that they were derived from primitive material characteristic of the first magmatic cycle. The apparent vast volumes of adamellitic and granitic material constituting the batholiths precludes their derivation from the same mafic precursors considered to be responsible for the earlier tonalites and trondhjemites. The batholiths may therefore be largely 87
87
86
87
86
8
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C. R. ANHAEUSSER & L. J. ROBB
shallow-seated bodies forming relatively thin sheets over much of the earlier-formed crust in the region (a concept similar to that of a granitic 4 'hood" or capping proposed by Hunter, 1973 and Viljoen & Viljoen, 1969a). The various components of the batholiths are post-tectonic and have not undergone the intense deformations that characterize the gneissic and migmatitic rocks of the earlier magmatic cycle. The marginal migmatites and gneisses of these large batholithic bodies contrast markedly with their earlier counterparts, being characterized by potassic leucosomes and more diffuse, schlieric textures (the latter defined by Mehnert, 1968). The marginal zones furthermore display ''inherited' ' structures that were produced during the earlier magmatic cycle, being manifest as nebulites or migmatites with ghost-like relics of preexisting rocks. These characteristics suggest that the style of batholith emplacement differed from that of the first magmatic cycle when gravity inversions were probably technically dominant. The absence of a penetrative fabric in any of the batholith components as well as the ubiquitous marginal migmatite phases that accompany these units indicates passive emplacement of magma (cf. the "permissive intrusions" of Pitcher, 1979) at relatively high crustal levels and without subsequent remobilization or re-emplacement. The batholith magmas and their incumbent, more volatile, phases appear to have been passively intruded into the pre-existing crust by processes such as cauldron subsidence, stoping and assimilation. In view of the large volumes of magma which it generated, the second magmatic cycle is considered to represent the main event contributing to cratonization of the early continental masses and the one during, and subsequent to which, tectonic stability prevailed.
THIRD MAGMATIC CYCLE The late granite plutons Twelve granitic plutons have been identified in the eastern Transvaal and Swaziland, all but one appearing on the map of the area (Fig. 1). The plutons are distinguished by their topographic expression, lithology, transgressive mode of emplacement and limited metamorphic effects (Hunter, 1973; Robb, 1978; Viljoen & Viljoen, 1969a, b). Most plutons are homogeneous, coarse-grained, porphyritic, K-rich bodies that, on the basis of their geochemical characteristics, may be regarded as adamellites or granites (sensu stricto), although undersaturated syenite bodies occur 15 km southeast of Badplaas (Anhaeusser
et al., 1979) and a tonalitic pluton occurs east of Bushbuckridge (Fig. 1). The plutons are classified into two groups embracing an older and a younger category (Hunter, 1973, 19746; Viljoen & Viljoen, 1969a, b) based partly on field relations established in Swaziland (Hunter, 1973) but principally on the basis of their isotopic ages. The emplacement of the older plutons appears to have commenced about 2900 Ma ago following cessation of the second magmatic cycle. The granite plutons were emplaced into a crustal environment largely thickened and stabilized by cratonization accompanying the two previous cycles. The older plutons Based on Rb-Sr whole-rock isochron ages, the older plutons fall within the time span 2927 ± 59 Ma to 2784 ± 53 Ma (Barton, 1981; Davies, 1970). Physically there is little to distinguish the K-rich granitic bodies from those of the younger plutons, which, with the exception of the tonalite body, form a rugged terrain characteristic of all remaining plutons in the eastern Transvaal and Swaziland. The older plutons are distinguished geochemically from the younger plutons mainly on the basis of lower K/Na ratios (0.64-0.93), higher K/Rb ratios ( - 2 6 0 ) and lower Rb/Sr ratios (0.15-0.28), and are granodioritic (Condie & Hunter, 1976; Glikson, 1976; Viljoen & Viljoen, 19696). The younger plutons Rb-Sr isochron ages suggest that the younger plutons range in age from 2608 ± 123 Ma to 2496 ± 176 Ma (Barton, 1981; Davies, 1970; de Gasparis, 1967). The bodies are topographically prominent, and the rock types are generally coarse-grained, pinkish, porphyritic adamellites or granites. They differ from the older plutons in possessing higher K/Na ratios (-1.65), lower K/Rb ratios ( - 1 6 6 ) and higher Rb/Sr ratios (0.66-16) (Condie & Hunter, 1976; Viljoen & Viljoen, 19696). Major-element analyses show little chemical distinction between the older and younger plutons (Table 1, columns 8 and 9), the main differences being in the alkali elements. The Na-rich pluton The topographically subdued tonalitic pluton to the east of Bushbuckridge (Fig. 1) is an anomaly within the late granite (sensu lato) plutons. This homogeneous, medium-coarse grained body which is characterized by low K/Na ratios ( - 0 . 4 ) , Rb/Sr ratios of between 0.06 and 0.22 and a K/Rb ratio of - 180 (Robb, 1977) contrasts with both the older and younger plutons. The tonalite pluton has a Rb-Sr isochron age of 2784
MAGMATIC CYCLES
465
± 53 Ma and an initial Sr/86Sr ratio of 0.7034 and sima was responsible for the development of ± 0.0003 making it more akin to the older, complex migmatites, concurrently deformed by granodioritic plutons than to the more evolved the high stresses prevalent in the area at this younger plutons. stage. The late granite (sensu lato) plutons of the 3. The dominant processes of cratonization third magmatic cycle generally have higher initial were progressively developed during the second 87Sr/ Sr ratios than the components of the first magmatic cycle and enormous volumes of K-rich and second magmatic cycles most likely because magma were passively emplaced into the crust many, if not all, of these plutons have been de- such that by 3000 Ma ago it was at least as thick rived from the reworking of pre-existing sialic as it is at present (-35 km). These batholiths (tonalitic to granitic) crust (Condie & Hunter, were probably emplaced at a relatively high crustal level and may occur as sheet-like masses 1976). developed over the earlier-formed sialic crust (the CONCLUSIONS granite-greenstone crust produced during the first 1. From a purely descriptive viewpoint it is magmatic cycle). convenient to categorize the numerous and 4. The emplacement of smaller, discrete, postdiverse granite types in the Archaean terrain of plutons associated with the third magthe eastern Transvaal and Swaziland, into three tectonic matic cycle coincided with the termination of broad groups. The three categories also reflect cratonization. These bodies did not contribute stages in the formation and genetic evolution of either to the construction or stabilthe early sialic crust in this area. The concept of significantly ization of the early continental crust but neverthree consecutive magmatic cycles, therefore, not theless represent the ultimate cycle in the formaonly has taxonomic validity but affords an under- tion of the granitic basement. Subsequently the standing of the processes by which the Kaapvaal stable platform that had formed was progrescratori was constructed. sively denuded and a succession of cratonic-type 2. The early stages of proto-cratonization in volcano-sedimentary basins developed during the the greater Barberton area were accomplished Proterozoic Era. during the first magmatic cycle. During this event, primitive tonalitic and trondhjemitic ACKNOWLEDGMENTS material was derived by melting of an ensimatic The authors wish to thank Professor D. I. source (see Green & Ring wood, 1968 and Lam- Groves who critically read the manuscript and bert & Wyllie, 1972), which is envisaged to have offered many helpful suggestions with regard to covered most of the area under discussion. The the style of presentation and the content. Thanks emplacement of early sial was largely driven by are also due to Mrs W. A. Job and Mr N. A. de gravitational overturning (Ramberg, 1967) so N. C. Gomes who assisted with the drafting of that the primitive continental mass was tectoni- the figure and Mrs L. Tyler and Mrs D. Amaler caily unstable. Intimate interaction between sial for typing the manuscript. 87
86
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627-662.
and the Earth's Crust. Academic Press, London. ROBB, L. J., 1977: The geology and geochemistry of the Archaean granite-greenstone terrain between Nelspruit and Bushbuckridge, Eastern Transvaal. M.Sc. Thesis, Univ. Witwatersrand, Johannesburg [unpublished]. 1978: A general geological description of the Archaean granitic terrane between Nelspruit and Bushbuckridge, Eastern Transvaal. Trans, geol.
dating of Onverwacht Group volcanics, southern Africa. Nature, Lond., 279, 298-300. Soc. S. Afr., 81, 331-338. HUNTER, D. R., 1957: The geology, petrology and , 1981: Detailed studies of select migmatite outclassification of the Swaziland granites and crops in the region southwest of the Barberton gneisses. Trans, geol. Soc. S. Afr., 60, 85-120. greenstone belt and their significance concerning , (Compiler) 1966: 1:125 000 Geological Map of the nature of the early Archaean crust in the region. Swaziland. Geol. Surv. Swaziland, Mbabane. Spec. Pubis geol. Soc. Aust., 7, 337-349. , 1970: The Ancient Gneiss Complex in Swaziland. Trans, geol. Soc. S. Afr., 73, 107-150. ROBB, L. J., & ANHAEUSSER, C . R., 1979: Discussion on "The geochemical nature of the Archaean Ancient , 1973: The granitic rocks of the Precambrian in Gneiss Complex and Granodiorite Suite, SwaziSwaziland. Spec. Pubis geol. Soc. S. Afr., 3, land: a preliminary study" by D. R. Hunter, F. 131-145. Barker and H. T. Millard, Jr. Precamb. Res., 10, , 1974a: Crustal development in the Kaapvaal 153-159. craton, I. The Archaean. Precamb. Res., 1, 259-294. W . M . , SUTCLIFFE, R. H . , & TROENG, , 19746: Crustal development in the Kaapvaal SCHWERDTNER, B., 1978: Patterns of total strain in the crestal craton, II. The Proterozoic. Precamo. Res 1 region of immature diapirs. Can. J. Earth Sci., 15, 295-326. 1437-1447. , 1979: The role of tonalitic and trondhjemitic rocks in the crustal development of Swaziland and SCHWERDTNER, W . M . , & LUMBERS, S. B., 1980: M a j o r the eastern Transvaal, South Africa; in Barker, F. diapiric structures in the Superior and Grenville (Ed.) Trondhjemites, Dacites, and Related Rocks, provinces of the Canadian Shield. Spec. Pap. geol 301-322. Elsevier, Amsterdam. Assoc. Can., 20, 149-180.
MAGMATIC CYCLES M. J., & V I L J O E N , R. P., 1969a: A proposed new classification of the granitic rocks of the Barberton region. Spec. Pubisgeol. Soc. S. Afr., 2, 153-180.
VILJOEN,
, 1969b: The geochemical evolution of the granitic rocks of the Barberton region. Spec. Pubis geol Soc. S. Afr., 2, 189-218. 1969c: The geology and the geochemistry of the lower ultramafic unit of the Onverwacht Group
467
and a proposed new class of igneous rocks. Spec. Pubis geol. Soc. S. Afr., 2, 55-85. — , 1969d: The geological and geochemical significance of the upper formations of the Onverwacht Group. Spec. Pubis geol. Soc. S. Afr., 2, 113-151. D. J. L. ( C O M P I L E R ) et a!., 1956: The geology of the Barberton area. Spec. Pubis geol. Surv. S. Afr., 15.
VISSER,
GEOCHEMICAL AND ISOTOPIC CONSTRAINTS ON THE ORIGIN AND SOURCE OF ARCHAEAN GRANITES Kent C. Condie
Department of Geoscience, New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, U,S.A. ABSTRACT Archaean granites (including quartz monzonites) can be divided into three geochemical classes based chiefly on REE distributions. Geochemical model studies, initial S r / S r ratios, and oxygen-isotope data are consistent with an origin for most Archaean granites by partial melting (~ 20%) of one of three rock types in the lower crust with short crustal residence times: tonalite-trondhjemite, high-grade gneiss (of intermediate composition), or less likely, greywacke. Only a few percent of subcontinental mantle heat is needed to account for Archaean granite magma production by partial melting of the lower crust. 87
INTRODUCTION Although composing only a small amount of the sialic Archaean crust, granites are important in evaluating the composition of magma sources during the Archaean, in constraining geothermal gradients, and in enhancing our understanding of Archaean tectonic settings. Archaean granites occur both as batholiths or parts of batholiths and as isolated plutons of various sizes (Condie, 1981). Although individual plutons range from <50 to 1000 km in area, most are in the range of 250-1000 km (Hunter 1974; Goodwin, 1978). Gravity studies indicate that many small plutons extend to depths of only 2-5 km and that most plutons bottom-out by 10 or 15 km (West et al., 1977). Gravity data also suggest that most Archaean plutons have greater lateral than vertical dimensions (Goodwin, 1978). Many small plutons are discordant and appear to be post-tectonic whereas batholiths, which are mostly concordant and often foliated, appear to represent syntectonic intrusions. Compositional variants of batholiths may grade into each other or form distinct plutons with sharp contacts. Some batholith contacts are migmatitic and grade into surrounding tonalite-trondhjemite gneiss terrains (Condie, 1981). Inclusions are uncommon in most granite plutons, and are small and randomly oriented except in contact zones where they generally parallel the contact. Supracrustal remnants abound locally in batholiths where they parallel foliation. Archaean granites typically range from medium to coarse grained and some contain large K-feldspar megacrysts. They are composed of approximately equal amounts of K-feldspar, Spec. Pubis geol. Soc. Aust., 7 (1981) 2
2
86
plagioclase (An -An o), and quartz, and usually contain minor biotite. In terms of composition, they are granite (Streckeisen, 1976) or granite and quartz monzonite (Johannsen, 1939). Modal analyses from the Laramie batholith in Wyoming (Condie, 1969) indicate that it comprises 74% quartz monzonite, 14% granodiorite, and 12% granite according to the classification of Johannsen (1939). In some batholiths, however, granodiorite and tonalite may comprise >50% of the exposed rocks. Textural relationships indicate that K-feldspar, biotite, and quartz crystallized late in Archaean granites (and quartz monzonites) allowing a wide range of water contents in the magmas at the time of crystallization (1-20% H 0 ) (Maaloe & Wyllie, 1975). However, pegmatites and aplites are generally rare and contact metamorphism is minimal, suggesting that the granites crystallized at the low end of this range. Normative compositions (Barth mesonorms) also plot very close to the minimum in the system Ab-0r-Q-H 0 at low water-pressures ( < l k b ) (Condie, 1969; Hunter, 1974) reinforcing this conclusion. Archaean granites (which hereafter will also include quartz monzonites) range in age from about 3.4 to 2.5 b.y. with the largest number having been emplaced between 2.5 and 2.6b.y. (Fig. 5). It is noteworthy that the 2.5-2.6b.y. period of granite intrusion follows the very extensive period of tonalite-trondhjemite and greenstonebelt formation at 2.6-2.7 b.y. Even granites older than 2.6 b.y. followed one or more periods of tonalite-trondhjemite and greenstone magmatism. Such a time relationship suggests that the prior existence of sialic crust is a necessary condition for granite production. 10
3
2
2
KENT C. CONDIE
470
Fig. 1.
Envelopes of REE variation in major Archaean granite groups; stippled area represents Group I.
COMPOSITION Archaean granitic rocks define two geochemical trends. The tonalite-trondhjemite trend is characterized by N a 2 0 enrichment (Condie, 1981). It is noteworthy that most Archaean gneissic complexes lie on the tonalite-trondhjemite trend and most batholiths and plutons lie on the calc-alkaline trend. Most Archaean granites are also enriched in Ba and Rb relative to Sr. In terms of REE distributions, Archaean granites fall broadly into three groups (Fig. 1;, Table I) (Condie & Hunter, 1976; Birk et al., 1979; Condie, 1981). Group I, which is most widespread, is characterized by light REE enrichment (100-200x chondrites), variable negative Eu anomalies, and only small fractionation of heavy REE. Group II, which is of minor importance in Archaean terrains, is characterized by moderate light REE enrichment (50-100x chondrites), small or negligible Eu anomalies, and depleted heavy REE. Group III is represented chiefly by small post-tectonic plutons, and is characterized by
light REE enrichment similar to or greater than observed in Group I, large negative Eu anomalies, and only slightly fractionated heavy REE. G E O C H E M I C A L M O D E L STUDIES Using both major and trace elements, fractional crystallization and partial melting processes have been evaluated for the production of Archaean granite magmas. Tested sources range in composition from ultramafic to tonalitic. Fractional crystallization is tested at various depths using the appropriate liquidus phases as determined experimentally and summarized in Condie & Hayslip (1975). In all models, it is assumed that granites represent liquids and do not contain residual source material. The overall compositional homogeneity of Archaean granites tends to support this assumption. Major-element models are evaluated using a statistical mixing program modified after Wright & Doherty (1970). Trace-element modelling employs the Rayleigh fractionation law and similar expres-
GEOCHEMICAL AND ISOTOPIC CONSTRAINTS
471
TABLE I
Average compositions of Archaean granites GROUP II
GROUP I
72-9
71-3
AI2O3
0.30
14.2 0.76 1.43 0.49 1.19 3-45 4.60 1.3
14.4 0.59 1.73 0.57 1.33 3.92 4.59 1.2
7
7 6 10 196 122
Ni Co Cr Rb Sr Ba Zr La Ce Sm Eu Tb Yb Lu K/Rb Rb/Sr Ba/Sr Ni/Co La/Yb Eu/Eu*
2
4 184 112
632
130 54 120 7.5 0.76 0.78 2.4 0.41
250 1.6 5.6 3.5 23 0.37
Chi 1imanz i
Average of Cols 1-5
Giants 4 Range
73.6 0.21 13.9 0.34 0.75 0.50 1.07 3.78 4.80 1.3
72.0 0.32 14.5 0.48 1.06 0.40 1.72 3.98 4.42 1.1
75.1 0. 19 14.0 0.19 0.81 0.12 1.09 3.62 4.31 1.2
73.0 0.27 14.2 0.47 1.16 0.42 1.28 3-75 4.54 1.2
73.0 0.19 14.9 0.54 1.01 0.49 1.10 3-90 4.47 1.2
5 3 8
5
135 571 122 59 131 7.8 0.50 0.51 1-5
6 5 8 193 462 1130 187 77 149 7.0 1.1 0.46 1.5
6 4 7 235 189 700 128 59 123 7.2 0.84 0.73
121 2.4 4.2 1.7 39 0.27
190 0.42 2.4 1.2 51 0.59
Salisbury
Si 02 Ti0 2 Fe 2 0 3 FeO MgO CaO Na 2 0 K20 K 2 0/Na 2
Matopos
Lochiel 2
0.32
500
100 70 131 9.1 1.5 1.3 2.5 0.36
234 1.6 4.1 1.2
28
0.52
328
0.26
3
0.22
2
6 276 114 657 99 37
86
4.5 0.35 0.58
2.2
0.33 160 1.2 3-7 1.5 32 0.52
0.26
Major-element oxides in percent, trace elements in ppm. ( ) = values in parentheses interpolated; n = number of samples in each average 1 Condie (1969); (1981) 2 Hunter (1974); Condie & Hunter (1976) 3 New data - determined by combined X-ray fluorescence and neutron activation methods * Arth & Hanson (1975)
sions for partial melting (Hanson, 1978). Partial melting models are for non-modal, equilibrium (batch) melting and for fractional melting. Distribution coefficients are estimated from available experimental data and from phenocryst-groundmass pairs, and are available from the author upon request. Granites of Groups I, II, and III were tested separately using a range of published compositions for each group. Fractional crystallization models, in general, result in poor agreement with observed granite compositions. One problem is the rapidity with which Co, Ni, and Cr are removed; these elements are usually exhausted in the magma before granitic liquids are produced. Also, a rather poor agreement for REE is obtained when minerals are removed in proportions necessary to produce satisfactory major-element agreement. Partial melting models involving ultramafic or mafic sources were also not very successful. Such models depend upon the degree of melting and the amount of minor minerals enriched in LIL elements. Even adjusting these two variables to obtain optimum agreement for major elements does not produce satisfactory agreement for Ba,
2.0
0.39 130 2.4 5.8 2.5 17
192 202
676
(33) 75 3.6 0.61
(0.28) 0.61 0.10
193 0.95 3-3 54 0.65
region 2 Plutons 72.1 0.36 13.6 0.91 1.62 0.49 1.58 3.39 5.17 1.5 4 3 6 250 195 800 150 121 223 15 1.5 2.1 5.1 0.64 172 1.3 4.1 -1.3 24 0.35
26
Sr, and REE. It is also difficult to find a satisfactory mechanism by which small amounts of melt (<2°7o for an ultramafic source) can be segregated from the source. Consistent with geochemical models which do not favor ultramafic or mafic sources for Archaean granites (all groups), are experimental data which also seem to eliminate such sources for granites (Wyllie et al., 1976). It should be pointed out, however, that a mafic source appears necessary for the production of Archaean tonalite-trondhjemite (Condie & Hunter, 1976; Condie, 1981). The most successful models for the production of Archaean granite of Groups I and II are those involving partial melting of intermediate, tonalitic, or greywacke sources (Arth & Hanson, 1975; Condie & Hunter, 1976). Compositions of possible Archaean sources for published average analyses are given in Table II. Also shown for comparison is average modern high-K andesite. The compositions of the Archaean high-grade gneiss terrains in Table II are similar to each other except for K, Rb, and heavy REE. The high-Al2C>3 tonalite-trondhjemite averages from Archaean low-grade (granite-greenstone) and
472
K E N T C.
CONDIE
TABLE II
A verage compositions of possible sources for Archaean granites
Lewi s ian, Scotland
High-Al203 Archaean Tonalite-Trondhjemite
High-Grade Gneiss Northern Norway
East Greenland
Low-Grade
High-Grade
61 2 0 6 16 4 2 3 3 5 3 0 4 4 4 0 3-.0 0. 75
65-9 0.74 15-9 2.19 2.92 2.19 3-76 3.46 1.89 0.55
69 .4 • 0 • 35 • 15 .8 1 .2 1 .8 1,. 1 3..4 4,.7 1,.6 0,.34
67-5 0.35 15.8 1.7 1.3 1.25 3-47 4.64 1.99
56 572 1160 144 36 77 8.4 2. 2 0.81 1.9 0.36
24 1 (12) 55 39 600 1090 300 30 70 8.0 2.1 0.85 1.0 0.15
15 8 12 45 460 400 100 25 42 2. 9 0. 82 0. 28 0 . 82 0 . 12
402 0.07 1.8 2 30 1.0
295 0 . 10 0.87 1.9 31 1.0
Si0 2 Ti02 A1203 Fe203 FeO MgO CaO Na20 K20 K20/Na20
61.5 0.56 15-5 3. 4 2.8 3-5 5.9 4.0 1.0 0.25
(60-65) (0.5-0.6) (15-16) (2.5-3-5) (2.0-4.5) (2.5-^.5) (4.5-7.2) (3.5-4.5) (0.8-1.2)
Ni Co Cr Rb Sr Ba Zr La Ce Sm Eu Tb Yb Lu
69 1 (35) 90 10 495 645 165 20 4o 2.5 1.0 0.3 0.7 0.12
(40-100)
K/Rb Rb/Sr Ba/Sr Ni/Co la/Yb Eu/Eu*
830 0.02 1.3 2 29 1.4
(50-150) (8-16) (350-550) (450-780) (115-200) (14-25) (25-50) (1.5-3.0) (0.9-1.1) (0.2-0.4) (0.2-1.5) (0.05-0.18)
445 0. 10 2. 0 19 1.0
Major-element oxides in percent, trace elements in ppm. References available from the author upon request. 1 Values estimated assuming Ni/Co = 2.
Archaean Greywacke
21 26 54 540 850 215 28 58
High-K Andesite
63 .7 0 .6 15 .2 1 .0 4 .9 3 .5 2 .6 3 .0 2 .5 0.83
(62.66) (0.5-0.6) (13-16) (0.7-1.5) (4-7) (3-4.5) (2-4.5) (2-3.5) (1.5-2.5)
60.2 0.95 16.9 2.6 2.8 2.2 5.5 3.7 2.8 0.76
50 15 20 77 290 470 170 33 66 4 1. 3 0. 5 1. 5 0 . 25
(30-80) (10-25) (10-50) (50-100) (100-400) (300-600) (150-200) (23-40) (50-90) (3.5-4.8) (0.9-1.6) (0.4-0.6) (1.5-1.6) (0.22-0.3)
40 20 90 80 700 700 200 43 84 5 1.4 0.56 1.6 0.27
270 0 . 27 1.6 3. 3 22 1. 1
208 0.11 1.0 2.0 2.7 1.0
Ranges of published mean values given in parentheses.
TABLE I I I
Assumed major-element and modal compositions of sources (Ca) and calculated residues after 20 per cent equilibrium melting (Cs) Tona1ite-Trondhjemite *
High-Grade Gne i ss
Greywacke
Low-Grade
High-Grade
Co
cs
Co
cs
Co
cs
Co
cs
66.9
67.1
63-9
63.7
61.8
Si0 2
63.1
59.7
69.3
Ti0 2
0.79
0.95
0.50
1.2
0.5
1.0
0.50
1.5
A1203
16.5
16.8
15.2
15.6
16.3
16.6
16.0
16.9
Fe203
1.9
2.2
1.3
1.7
1.8
2.4
1.1
1.6
FeO
2.9
3.7
2.1
2.5
3.0
3.6
3.4
4.0
MgO
3.5
5.4
1.4
1..8
1.5
1.8
4.0
4.9
CaO
5.0
5-7
3.1
3.5
3.4
3.7
Na20
4.1
3-9
5.0
5.0
4.7
4.9
3.4
K20
2.1
3.5
1.4
1.7
1.1
M
1.5
2.4
2.0
Plagioclase
48 (An-,,)
56
53 (Anao)
58
41 (Anao)
44
-52
51
(An25)
2.7
2.8
K-feldspar
13
7
7
-
15
10
Quartz
9
2.5
15
11
27
26
19
22
20
Hornblende
10
5
6
Cordierite
_
13
_
16
Orthopyroxene
10
13
-
-
10
12
Biot ite
2
7.5
8
9
-
_
6.5
1.5
14
16
Magnetite11 men i te
2
2.5
2
2.5
3
4
2
2.5
* Modal granodiorite
_
-
6
7.5
G E O C H E M I C A L AND ISOTOPIC CONSTRAINTS
high-grade terrains are also similar to each other except for Ba and Zr. Archaean greywackes differ from both high-grade gneiss and tonalitetrondhjemite compositions by their greater Fe and Rb and lower Ca, Sr, and Na. High-K andesite is broadly similar to all three sources, differing only by higher Ti, Al, Ba, and light REE. From reported mineral abundances and compositions in the sources tested, modes are calculated which result in major-element compositions similar to those given in Table II. These are tabulated in Table III together with residue compositions after 20% melting. The tonalitetrondhjemite sources are granodiorites in their modes but not in their chemistry. Fractions of each mineral entering the melt are estimated from experimental data (Wyllie et al., 1976) and from modal analyses of Archaean granites, and are as follows: K-feldspar, 0.35; plagioclase, 0.31; quartz, 0.30; and biotite, 0.04. Of the equilibrium and fractional melting models tested, the equilibrium melting models produced the best overall agreement between calculated and observed source compositions. Trace-element
Fig. 2.
473
concentrations in the source (C0) are calculated from the relationship, C L /C 0 = 1/F + D (1 - P) for non-modal, equilibrium melting where CL is the concentration of an element in the melt, F is the fraction melted, D is the bulk distribution coefficient for the source, and P the bulk distribution coefficient for minerals entering the melt. Both minor- and trace-element results suggest melting in the range of 15-30%. Calculated traceelement concentrations in each of the three sources for Group I granites (Table I) are tabulated in Table IV. The overall agreement between the compositions of calculated sources (Table IV) and Archaean rocks that may serve as sources for granites (Table II) is, on the whole, satisfactory. The models are illustrated on a Na 2 0-K 2 0-Ca0 diagram in Figure 2. Tonalite-trondhjemite and high-grade gneiss compositions define a region of acceptable source rocks for Archaean granites. Residues after melting also fall chiefly within this region as shown by average compositions from Table III that are plotted on the diagram. It is noteworthy that the average compositions of
C a 0 - K 2 0 - N a 2 0 diagram showing source, melt, and residue relationships for Archaean granites. T-T = tonalite-trondhjemite: • = average Archaean granite of Group I; + = average compositions of sources given in Table III; o = average compositions of residues given in Table III; • = average high-K andesite (Table II); • = average Archaean granite-greenstone terrain; A = average Archaean greywacke.
474
KENT C. CONDIE TABLE I V
Calculated trace-element concentrations in model sources for Archaean granites of Group I Tonalite-Trondhjemite
Hi gh-Grade Gnei ss
Greywacke
Hi gh-Grade
Low-Grade
Cr
31
(20-35)
23
(14-27)
30
(15-40)
34
Ni
45
(40-50)
22
(18-26)
39
(35-45)
45
(35-50)
Co
22
(12-35)
14
(8-20)
13
(7-20)
21
(12-35)
(60-110)
62
(50-85)
120
(90-160)
Rb
71
80
(60-100)
(20-50)
Sr
435
(250-1000)
380
(220-900)
330
(200-800)
315
Ba
655
(500-1100)
621
(450-1000)
670
(480-1100)
990
(700-1600)
Zr
100
(75-140)
110
(85-150)
25
(15-30)
95
'(80-175)
La
27
(20-35)
24
(15-30)
23
(15-30)
22
(15-30)
Ce
66
(50-80)
55
(40-65)
48
(35-60)
45
(30-55)
(3-5-7.0)
3- 6
(2.5-4.5)
2. 3
(1.5-3.0)
2. 3
(1.5-3.0)
Sm
5. 7
Eu
1..4
(0.6-1,8)
1. 1
(0.5-2.0)
1. 1
0.85
(0.3-1.5)
Tb
0..6
(0.4-1.1)
0.36
(0.25-0.65)
0. 2
(0.1-0.4)
0. 20
(0.1-0.4)
Yb
1..6
(1.2-2.0)
0. 95
(0.7-1.2)
0. 55
(0.5-0.8)
0.65
(0.5-0.8)
Lu
0..25;
(0.2-0.3)
0. 15
(0.1-0.2)
0. 10
(0.07-0.12)
0. 1 1 (0.07-0.14)
Mean values given with ranges in parentheses.
Results are for 20 percent equilibrium melting of sources given in Table III.
Archaean high-grade gneiss and granite-greenstone terrain are similar to each other as well as to average high-K andesite. The position of Archaean high-grade gneisses between the tonalite-trondhjemite and basalt fields in Figure 2 is consistent with the interpretation of Tarney (1976) that the average composition of highgrade gneisses results from either structural or chemical mixing of these two end members, which originally were part of a bimodal association as defined by Barker & Peterman (1974). The Archaean greywacke source is distinctly more K 0-rich than the tonalite or gneiss sources and probably results from Archaean greywackes being composed of a mixture of materials derived from mafic, tonalite-trondhjemite, and granite sources. It is noteworthy that all model sources (Table IV) and especially the high-grade gneiss source have greater Rb contents than exposed rocks that may serve as sources (Table II). It is not possible to reduce the magnitude of this discrepancy by decreasing the amount of K-feldspar or biotite in the model source because the K 0 content dictates the values of these minerals as given in Table III. The residues remaining after 20% melting of the model high-grade gneiss source, however, contain Rb in the same concentration range (20-40 ppm) as observed in exposed highgrade gneisses (Table II); hence the Rb data are consistent with the high-grade gneisses representing residues from which granite magmas have been removed. Major-element and other traceelement contents of model residues also accord with this interpretation. Melting relationships for Ba-Rb-Sr are shown in Figure 3. The tonalite2
(0.5-2.0)
(180-770)
trondhjemite and high-grade gneiss sources are characterized by relatively low Rb content on this diagram. Residues after extraction of granite magma are depleted even more in Rb and may be enriched in Sr, which is housed in residual plagioclase. The large amount of Ba in the calculated greywacke source differs from that observed in Archaean greywackes as illustrated. Again this cannot be adjusted by decreasing K-feldspar or biotite in the source because K 0 dictates the amounts of these minerals present. The disparity in Ba contents may reflect preferential loss of Ba relative to Rb and Sr in greywackes during weathering and erosion, or during metamorphism. 2
2
Fig. 3, Ba-Rb-Sr diagram showing source, melt, and residue relationships for Archaean granites. Symbols as in Figure 2.
GEOCHEMICAL AND ISOTOPIC CONSTRAINTS REE patterns for the four calculated sources together with that of average Group I granite are shown in Figure 4. All REE patterns are similar for light REE but deviate f r o m each other for heavy REE and Eu anomalies. The positive Eu anomalies in the high-grade tonalite-trondhjemite and greywacke sources reflect the absence of residual hornblende, which generally has a negative Eu anomaly. It is possible to adjust the heavy REE abundances, within the limits allowed by major elements, by small changes in modal hornblende or addition of minor amounts of garnet. For instance, removal of hornblende from the mode of high-grade gneiss (Table III) (with corresponding increases in plagioclase and pyroxene) reduces the heavy REE content of the source by a factor of two, bringing the calculated source more in line with Lewisian high-grade gneisses (Table II). G r o u p II Archaean granites can be produced f r o m the same four sources as Group I by adding garnet (with K d of heavy REE = 50) to the mode at the expense of hornblende. As an example, replacement of garnet for hornblende in the high-grade gneiss source reduces the heavy REE content of derivative granite magmas by an order of magnitude without appreciably affecting other trace- or major-element concentrations. Variable Eu anomalies in Archaean granites of all three groups is explained most readily by differences in magma oxygen fugacity (Condie & Hunter, 1976).
Fig. 4.
475
Group III granites appear to require a twostage model for their production (Condie & Hunter, 1976). An acceptable model for most of the granites involves, (1) production of granodiorite magma in a manner similar to that in which Group I granites are produced (with a somewhat larger degree of melting) followed by (2) fractional crystallization at shallow depths. Archaean syenites and monzonites (and related rocks), which are rare in Archaean terrains, appear to require an upper mantle origin involving very small amounts of melting of garnet peridotite ( ± eclogite) (Arth & Hanson, 1975). STRONTIUM- A N D OXYGEN-ISOTOPE DATA ratios f r o m Published initial 8 7 S r / 8 6 S r Archaean granites are plotted on a Sr evolution diagram in Figure 5. Several features relevant to the origin of Archaean granites are apparent f r o m the distribution of points on this diagram. Most of the granites were emplaced between 2.5 and 2.6b.y. immediately following the widespread magmatism and orogeny at 2.6-2.7 b.y. Only a few plutons are greater than 3 b.y. in age; the oldest granite is the Mushandike pluton in Rhodesia dated at 3.45b.y. (Hickman, 1974). Most Archaean granites have initial 8 7 S r / 8 6 S r ratios that lie above the mantle growth envelope. Only a few, however, have ratios greater than
REE patterns of average Archaean granite of Group I and of calculated sources. T-T tonalite-trondhjemite.
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Sr/ 8 6 Sr evolution diagram showing initial 8 7 Sr/ 8 6 Sr ratios of Archaean granites. All dates are calculated with X87Rb - 1.42 x 10~ H y r ~ 1 . References may be obtained from the author upon request.
0.710 or less than 0.701, with most ratios falling in the range of 0.702 to 0.707. Any explanation for the array of initial 87 Sr/ 8 6 Sr ratios exhibited by Archaean granites must also be consistent with the geochemical data necessitating a tonalite-trondhjemite, high-grade gneiss, or greywacke source. Several models merit consideration: (1) Production of variable initial 8 7 Sr/ 8 6 Sr ratios in the mantle during core formation in a manner similar to that suggested by Hurst (1978). This model involves a chondritic earth which undergoes fractionation during core formation over an interval of 100-300 Ma after accretion. Rb follows Fe into the core under reducing conditions resulting in production of low Rb/Sr (0.01-0.03) mantle reservoirs with variable initial 8 7 Sr/ 8 6 Sr ratios which lie on a chondritic growth curve. These reservoirs are
later tapped producing tonalitic crust with variable 87 Sr/ 8 6Sr ratios (0.700-0.708) which, in turn, serve as sources for granites (with or without crustal residence times). A major problem with this model is the lack of evidence from Archaean basalts for high initial 87 Sr/ 8<Sr ratios in the mantle. Another problem with the model is that clustering of initial 8 7 Sr/ 8 6 Sr ratios from the same geographic location, which may be expected for magmas derived from a mantle source with the same Rb/Sr growth curve, is not widely recognized. (2) Partial melting of mixed older and younger tonalitic, high-grade gneiss, or greywacke sources at 2.5-2.6 b.y. For instance, if 2.7 b.y. tonalites were widely injected into 3.7b.y. tonalitic crust and the resulting mixture was partially melted in varying ratios at
GEOCHEMICAL AND ISOTOPIC CONSTRAINTS
2.5-2.6b.y., an array of initial Sr/ Sr ratios in derivative granites could be produced. Most of the initial Sr/ Sr ratios of granites lie between the 3.7 b.y. tonalite growth curve (Rb/Sr-0.2 in Fig. 5) and the mantle growth envelope at 2.6 b.y. and thus are consistent with this model. It is not possible by this mechanism, however, to explain the very low (<0.701) or very high (>0.708) initial ratios. Also, the model necessitates widespread distribution of 3.5-3.7b.y. sialic crust, evidence for which is non-existent. Available radiometric dates indicate that sialic crust of this age was of only local importance. (3) Continuous extraction of tonalite-trondhjemite from mafic mantle sources between 3.7 and 2.7b.y. such that at 2.5-2.6b.y. tonalitic and high-grade sources existed with an array of Sr/ Sr values between 0.701 and 0.707. These sources are then partially melted to produce the corresponding array of initial Sr/ Sr ratios in Archaean granites at 2.5-2.6 b.y. This mechanism neither explains the very high and very low initial ratios, nor is supported by the distribution of Archaean radiometric dates. These dates cluster at 3.5-3.7, 2.9-3.0, and 2.6-2.7b.y. and should result in corresponding clusters of Sr/ Sr ratios at 2.5-2.6 b.y. rather than a continuum of points as observed. (4) The array of initial Sr ratios does not reflect ' source materials but is produced by losses and/or additions of Sr to Archaean granites causing rotation of isochrons. Small additions or losses of Sr from Archaean granites with low Sr content can significantly change initial Sr/ Sr ratios without greatly changing the isochron age (±50Ma). Initial ratios may be either raised or lowered depending on whether samples with low or high Rb/Sr ratios are affected most by Sr losses or gains. This mechanism can explain all but the very high (>0.710) initial Sr/ *Sr ratios in Figure 5. Also differences between zircon dates and RbSr isochron dates from the same plutons seem to indicate that loss of Sr from high Rb/Sr samples in granites during or after crystallization may be a common phenomenon (Bickford & Mose, 1975). This process is tentatively preferred to explain the broad continuum of initial S r / S r ratios observed in Archaean granites. The high initial Sr/ Sr ratios in a few Archaean granites is difficult to explain by any of the above mechanisms and appears to require selective contamination of granitic plutons with radiogenic Sr. 87
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477 Oxygen-isotope distributions in Archaean granites provide another constraint on their origin. The 6 0 values (%>) in Archaean granites range from about + 2 to +9 with most falling between +6 and +9 (Taylor, 1977). The upper end of this range exceeds mantle <5C> values ( + 5.5 to +7) and hence many granites cannot be produced by partial melting of the mantle. Most Archaean greywackes have <5 0 values that exceed + 9 (Longstaffe & Schwarcz, 1977) and hence greywacke also is not favored as an important source for Archaean granites. Archaean high-grade gneiss and tonalitetrondhjemite broadly overlap Archaean granites in their 5 O values and are allowable sources for the granites. The low <5 0 values ( < + 6) in some Archaean granites are perhaps best explained by shallow interaction of magma and meteoric waters which have low 6 0 values. DISCUSSION Geochemical model studies, initial Sr/ Sr ratios, 5 0 values, and experimental petrologic data are all consistent with an origin for Archaean granites involving partial melting of sources composed of high-Al203 tonalitetrondhjemite, high-grade gneiss of intermediate composition, or greywacke. Greywacke sources, although possibly of local importance, are. not considered of general importance in the production of Archaean granites because <5 O values in most Archaean greywackes exceed values in Archaean granites and geologic evidence does not favor greywacke as a major rock type in Archaean terrains. It is not possible geochemically or isotopically to distinguish between Archaean high-grade gneiss and high-Al 0 tonalite-trondhjemite sources. Mixing of the basaltic and tonalitic end members of the bimodal suite appears to be responsible for high-grade gneisses having the overall average composition of high-K andesite. Hence, whether tonalite-trondhjemite or high-grade gneiss were of greater importance in terms of Archaean granite sources depends on the degree of chemical or deformational mixing in the lower crust. The fact that some high-grade Archaean terrains are not well mixed and still show compositional bimodality (Weaver et al., 1978), indicates that mixing is not a necessary consequence of burial in the lower crust. Some Archaean high-grade terrains that are low in Rb and K, such as the Lewisian in Scotland, are possible candidates for the residue remaining after granite extraction (Muecke et al., 1979). It is noteworthy and perhaps significant that the Lewisian terrain is also well mixed (i.e., nonbimodal), a feature which may develop during partial melting and granite removal. 18
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The array of initial S r / S r ratios observed in Archaean granites may be produced by losses or gains of Sr during or after crystallization in response to thermal metamorphism or interaction with hydrothermal fluids. If so, most tonalitetrondhjemite and high-grade gneiss source materials were probably produced during the widespread 2.6-2.7 b.y. Archaean magmatism and deformation event, and had only moderate crustal residence times (100-300 Ma) prior to the partial melting which produced granite. It is of interest to estimate the volume of lower crust involved in the production of Archaean granite. Estimates of the total volume of Archaean granite (including granodiorite) in granite-greenstone provinces range from about 15 /o in the Superior Province north of Lake Superior (Goodwin, 1978) to 40-50% in Rhodesia (Hunter, 1974) and in the western Superior Province. Overall, it would appear that 20% is a reasonable average. Let us assume an original crust 40 km thick in which the upper 10 km has been removed during uplift and erosion to produce the exposure level observed in Archaean granite-greenstone terrains. Consider a 1 km column of crust today, 30 km thick and assume the upper 10 km is characterized by the granite distribution we see at the surface. This amounts to 2 km of granite. If the 10 km of crust removed by erosion contained half as much granite (1 km ) and an equivalent amount was trapped in the lower crustal source, we have a total of 4 km of granite in the 1 km column of crust. This matches the amount that would be produced by 20% melting of the lower 20 km of crust. Even considering the uncertainty in the values estimated in this calculation, it would seem that most or all of the lower crust must have been involved in the production of Archaean granites and hence that Archaean granite-greenstone terrains are underlain by depleted crust. 87
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The formation of most Archaean granites between 2.5 and 2.6 b.y. probably reflects the sparsity of sialic crust older than 2.7b.y. Many granites are post-tectonic and appear to have been intruded during a stage that lacked significant deformation. If some of these magmas were extruded, the resulting volcanic edifices must have been removed by subsequent erosion. It appears that the continents thickened significantly at 2.6-2.7 b.y. chiefly by tonalite-trondhjemite magmatism, and that the root zones began to melt near the end of this time to produce granites. Considering the compositions of the residues remaining after one period of granite extraction (15-20% melting), the abundances of most elements allow a second period of granite production with approximately the same range of melting. However, some elements, such as Rb, do not allow a second period of melting in which granite magmas can be produced unless these elements are replenished in the source. It is of interest to evaluate how much heat is necessary to produce Archaean granites and the degree to which upper mantle heat is decreased during their production. The 1 km crustal column described above would contain 10 gm of granite which requires approximately 10 cal of heat for its production. Assuming a minimum heat flow beneath Archaean continents equivalent to the present continental average (1.5 HFU or 63 mW/m ) and a minimum time period of 50m.y. over which Archaean granites were formed, approximately 24 x 10 cal/km of heat are available from the mantle. This places an upper limit on the amount of mantle heat necessary to produce Archaean granites of about 4%. Higher Archaean heat flows and/or longer time intervals over which granite is produced result in even smaller percentages of mantle heat necessary for granite magma production. 2
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REFERENCES
ARTH, J . G., & HANSON, G. N . , 1975: Geochemistry
and origin of the early Precambrian crust of northeastern Minnesota. Geochim. cosmochim. Acta,
39, 325-362. BARKER, F., & PETERMAN, Z . E . , 1974: Bimodal tholei-
itic-dacitic magmatism and the early Precambrian crust. Precamb. Res., 1, 1-12. BICKFORD, M. E . , & MOSE, D . G., 1975: Geochronology of Precambrian rocks in the St Francois Mountains, Southeastern Missouri. Mem. geol. Soc. Am., 165.
BIRK, D . , KOLJONEU, T . , & ROSENBERG, R. J . , 1979:
Rare-earth distribution in Archaean granitoid
plutons of the Wabigoon volcanic-plutonic belt, northwestern Ontario. Can. J. Earth 5c/., 16,
270-289. CONDIE, K. C . , 1969: Petrology and geochemistry of the
Laramie batholith and related metamorphic rocks of Precambrian age, eastern Wyoming. Bull. geol.
Soc. Am.,
80, 57-82.
, 1981: Archaean Greenstone Belts. Elsevier, Amsterdam. CONDIE, K. C . , & HAYSLIP, D . L., 1975: Young bimodal volcanism at the Medicine Lake volcanic center, northern California. Geochim. cosmochim. Acta, 39, 1165-1178.
GEOCHEMICAL AND ISOTOPIC CONSTRAINTS CONDIE, K. C . , & HUNTER, D . R . , 1976: Trace element
geochemistry of Archaean granitic rocks from the Barberton region, South Africa. Earth planet. Sci. Lett., 29, 3 8 9 - 4 0 0 . GOODWIN, A. ML, 1978: Archaean crust in the Superior Geotraverse area: geologic overview; in Smith, I. E. ML, & Williams, J. G. (Eds) Proc. 1978 Archaean Geoch. Conf., 73-106. Univ. Toronto Press, Ontario. HANSON, G . N . , 1978: The application of trace elements to the petrogenesis of igneous rocks of granitic composition. Earth planet. Sci. Lett., 38, 26-43. HICKMAN, M. H . , 1974: A 3500-myr-old granite in southern Africa. Nature, Lond., 251, 295-296. HUNTER, D . R . , 1974: Crustal development in the Kaapvaal craton, II. The Proterozoic. Precamb. Res., 1,
295-326. HURST, R. W., 1978: Sr evolution in the West Green-
land-Labrador craton: a model for early Rb depletion in the mantle. Geochim. cosmochim. Acta, 42,
39-44. JOHANNSEN, A . , 1939: A Descriptive Petrography
of the Igneous Rocks, 1. Univ. Chicago Press, Chicago.
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MUECKE, G. K . , PRIDE, C., & SARKAR, P . , 1979: Rare-
earth geochemistry of regional metamorphic rocks. Phys. Chem. Earth, 11, 449-463. STRECKEISEN, A. L . , 1976: To each plutonic rock its proper name. Earth Sci. Rev., 12, 1-33. TARNEY, J . , 1976: Geochemistry of Archaean highgrade gneisses, with implications as to the origin and evolution of the Precambrian crust; in Windley, B. F. (Ed.) The Early History of the Earth, 405-418. Wiley, New York. TAYLOR, H. P., 1977: Water/rock interactions and the origin of H2O in granitic batholiths. J. geol. Soc. Lond., 133, 509-558.
WEAVER, B. L . , TARNEY, J . , WINDLEY, B. F . , SUGAVA-
NAM, E. B., & RAO, V. V., 1978: Madras granulites: geochemistry and P-T conditions of crystallization; in Windley, B. F., & Naqvi, S. M. (Eds) Archaean Geochemistry, 177-204. Elsevier, Amsterdam.
WEST, G . F . , SZEWEZYK, Z . , DUSANOWSKYJ, T . , URQUART, W . E . , WRIGHT, J . , & GODLEWSKI, M . ,
18o/160 of Archaean clastic metasedimentary rocks: a petrogenetic indicator for Archaean gneisses? Geochim. cosmochim. Acta, 41,
1977: Geophysical studies in the Geotraverse: gravity, magnetics, seismics. 1977 Geotraverse Conf., Univ. Toronto, 1-9. WRIGHT, T . L . , & DOHERTY, P . C., 1970: A linear programming and least squares computer method for solving petrologic mixing problems. Bull. geol.
granitic magma deduced from the sequence of crystallization determined experimentally with ' water-undersaturated conditions. Contrib. Mineral. Petrol., 52, 175-191.
impossible sources, water contents, and crystallization sequences. Can. J. Earth Sci., 13, 1007-1019.
LONGSTAFFE, F . J . , & SCHWARCZ, H . P . ,
1977:
1303-1312. MAALOE, S., & WYLLIE, P . J . , 1975: Water content of
Soc. Am., 81, 1995-2008. WYLLIE, P . J . , HUANG, W . L . , STERN, C . R . , & MAALOE, S., 1976: Granitic magmas: possible and
ORIGIN OF HORIZONTAL STRUCTURE IN HIGH-GRADE ARCHAEAN TERRAINS R. G. Park
Department of Geology, University of Keele, Keele, Staffordshire, United Kingdom ST5 5BG ABSTRACT Archaean terrains of the high-grade gneiss-migmatite type such as the North Atlantic craton contain large areas of dominantly sub-horizontal foliation and recumbent folds when the effects of later deformation are removed. This foliation was produced by sub-vertical flattening as shown for example by deformed agmatites etc. and typically parallels lithological layering. Strains of X: Z = 50:1 are common. Four mechanisms are discussed in detail in terms of their tectonic effects: subduction (the uniformitarian plate-tectonic model), gravity spreading, mantle decoupling and thinned-crust collision. The conventional subduction model is unattractive for two reasons: (1) the thinner Archaean lithosphere may have been incapable of sinking (2) subduction imposes a dipping zone of simple shear strain which cannot easily explain the observed horizontal structures. Gravity spreading associated with the diapiric emplacement of granite plutons in granitegreenstone terrains would produce large flattening strains in the granitic source layer. A regularly repeated pattern of radial constrictional strains pointing toward the root zones of the rising diapirs should occur in the high-grade gneiss regions if they represent such a layer, but has not so far been recognised. Mantle decoupling could take place in the absence of subduction if horizontal translations of the crust due to mantle spreading were resisted. A horizontal shear zone would be formed in the lower crust which could explain the observed strain patterns satisfactorily. Likely values for shear zone width and shear strain place limits of a few hundred km on the displacements, which would take place over relatively short periods of time. Thinned-crust collision is suggested as a method for producing high strains in the lower crust by extension over a spreading axis. Subsequent compression on collision could explain the complex interlayering of highly strained gneisses, mafic intrusives, and supracrustals seen in the thickened Archaean crust of the North Atlantic craton.
INTRODUCTION In the proliferation of models proposed over the last few years to explain the tectonic evolution of the Archaean, the evidence afforded by Archaean structures has been comparatively neglected. In an attempt to understand the tectonics of the high-grade gneiss terrains, it is proposed to review various mechanisms which have been put forward to explain the nature and distribution of Archaean rocks in general, in terms of their structural implications. Archaean structures which have now been studied in considerable detail in many parts of the world place certain limits on the feasibility of the proposed models. Archaean tectonics have been traditionally subdivided into granite-greenstone and high-grade gneiss types, although there is increasing evidence of terrains transitional between the two. The granite-greenstone terrains have been widely Spec. Pubis geol. Soc. Aust., 7 (1981)
regarded as exhibiting a tectonic pattern dominated by the effects of diapiric upwelling of basement or granite domes producing characteristically upright structures in the greenstones arranged around the granite margins. Although modified by the recognition of thrust and nappe structures in some greenstone belts, and by the clear evidence in many terrains of regional compression producing an overall structural 'grain", many geologists probably regard these as modifications or complications of a structural situation dominated, or at least initiated, by gravity tectonics and vertical movements. In contrast, the high-grade gneiss-migmatite terrains, representing deeper levels of the crust, are regarded as areas of intense deformation with the interleaving of supracrustal and basement sheets together with granites, imposing a generally sub-horizontal layering which subsequently becomes deformed by more upright structures. 4
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Thus there has arisen an Archaean structural gave simple structural styles with all rock types dichotomy expressed in terms of upright struc- possessing similar rheology. tures in greenstone-granite terrains and horizonIt is clear from the work of several authors on tal structures in high-grade gneiss terrains. the Godthaab craton that the regular banding in these rocks is the result of very high strains. Escher & Watterson (1974) show that with a bulk ARCHAEAN TECTONIC PATTERNS strain ratio of X: Z = 50:1 considered typical of such terrains, a high proportion of pre-existing Structure of High-grade Gneiss Terrains planes will be rotated to within 10° of the direcArchaean high-grade gneiss-migmatite terrains tion of maximum elongation. Such high strains have been described from a number of Pre- can most conveniently be achieved by simple cambrian shields. The best known is undoubtedly shear, as these authors point out. The alternative the Godthaab craton of South Greenland and the of flattening due to crustal load as an explanation adjoining East Nain province of Labrador (cf. for these horizontal structures is untenable in Bridgwater et al., 1973). The classic Scourian view of the space problem posed by accommocomplex of Scotland presumably formed part of dating the large strains. this craton. Other examples include the Aldan It is important to stress that the structures, and Angara blocks of Siberia, the granulite ter- although affected by later folding, do indeed rain of southern Peninsular India (Naqvi et al., appear to be horizontal or gently dipping on a 1978) the Limpopo belt of Rhodesia (Mason, regional scale. This can be checked by examining 1973) and the western Yilgarn province of the outcrop pattern produced by later upright and Western Australia (Rutland, 1973). There are in more or less symmetrical folding of dome-andaddition a number of well-described belts either basin type in the Godthaab craton, which sugmarginal to or crossing greenstone terrains (e.g. gests that over distances of several tens of km the the Pikwitonei granulites and English River enveloping surface of the later folds is approxigneiss belt of the Superior Province). mately flat and the same tectonic level is repeated The Godthaab craton is regarded as the "type several times by the later folds (e.g. see Chadwick example" of such terrains. An excellent descrip- & Nutman, 1979). So that whatever is responsible tion is provided by McGregor (1973). "Most of for the structures could be associated with a horithe rocks . . . have suffered very intense strain zontal or very gently dipping zone of simple deformation: flattening and/or stretching. In shear. most places this deformation was sufficiently intense to rotate all earlier structures until they Structure of Greenstone-granite Terrains become effectively parallel (cf. Watterson, 1968). Any model to explain the structure of the highMany contacts that originally were discordant are grade gneiss terrains ought to be compatible with now concordant. Rocks that originally were granite-greenstone tectonics, since the two types mesoscopically inhomogeneous, with different of terrain were probably developing during the lithologies cutting or enclosed by one another, same period. have been deformed into regularly and finely It would seem that some kind of diapiric model banded gneisses, amphibolites, etc.". According is most appropriate for the typical graniteto Bridgwater et al. (1974) the layering was pro- greenstone development although perhaps modiduced mainly by intense tectonic activity domi- fied by subsequent tectonic processes of a differnated by horizontal movements involving major ent type. The key to the process is the gravitathrusting and overfolding and accompanied by tional instability of thick, more dense greenstone the emplacement of very large amounts of volcanics overlying a less dense granitic crust, granitic material as sub-horizontal sheets. At resulting in a two-way flow of crustal material— least two phases of deformation were involved, the denser mafic material transformed into amseparated by the intrusion of the Ameralik dykes. phibolite flows downward toward the base of the Individual isoclinal recumbent folds are traceable and the lighter granitic material flows upfor several tens of km. This deformation was crust toward the top. The principle has been disfollowed by upright folding which produced ward cussed geologists. In a recent analysis, dome-and-basin interference structures. All these Gormanbyetmany al. show how the gravitational structures are affected by the regional 2800 Ma model (based on(1978) the experimental work of Ramgranulite-facies metamorphism. berg, 1973) explains not only the vertical tecFrom evidence of this kind, Holland & Lam- tonics of greenstone belts but also thrust and bert (1969) describe the lowest crustal levels as a nappe around the rims of the basins as region of granulite-facies metamorphic fabrics a resultformation of inward lateral movement of the diawhere deformation by "prolonged laminar flow" piric basement. The movements take place in
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essentially solid material: it is clear from there is a tectonic junction with a granitoid terRam berg's experiments that granitic melts would rain enclosing migmatised greenstone relics internot produce the required large dome structures preted as a lower crustal level. The Archaean but would form smaller sheetlike or lensoid craton of Peninsular India also exhibits a transibodies which would not be capable of controlling tion between typical greenstone terrains in the the tectonic patterns on a sufficiently large scale. north and high-grade gneisses in the south (Naqvi Gorman et al. suggest that the tight synformal et al., 1978; Glikson 1979). greenstone "keels" produced in their model deGlikson believes that the two types of terrain scend to much greater depths than the 3-6 km are coeval, merely representing different crustal indicated by gravity studies as the base of the levels, and that gneiss-granulite terrains are exgreenstone basins, and that they may swell out at posed by "regional tilting" away from them depth or perhaps even detach as they sink into the toward the greenstone terrains. The presence of lower crust. high-grade gneiss or granulite belts within The intrusive tonalitic granite plutons which granite-greenstone terrains may thus be plausibly penetrate the basement diapirs may be derived by interpreted as tectonically elevated basement melting of the amphibolitic greenstone roots on possibly brought to the surface on gently inclined their descent to the lower crust (cf. Glikson, shear zones {cf. Coward, in press). According to this view, the higher-grade zones could represent 1979). A different, but not entirely conflicting, model uplifted (possibly thickened) segments of a lower is presented by Tarney et al. (1976) who use the crustal regime which may extend everywhere analogy of the modern Pacific marginal basin (cf. below the typical greenstone terrain. Whether this Karig, 1971) to explain greenstone-belt develop- regime corresponds specifically in structure and ment. According to this model the crustal thin- metamorphism with the highly deformed horining, early basic volcanism, later acid plutonism zontally-layered granulites of the North Atlantic and subsequent tectonic closure result from sub- craton, however, is debatable. duction at a nearby oceanic margin in a fashion similar to the geologically recent behaviour of the Constraints on Thickness and Thermal Gradient in the Archaean Crust and Lithosphere Pacific margins. Wide variation in estimates of Archaean Relationship between Greenstone-granite and crustal thickness probably reflects in part a genuHigh-grade Gneiss Terrains ine variation both in space and time. There seems It was originally suggested by Windley & to be general agreement that addition of large Bridgwater (1971) that the high-grade gneiss- volumes of granitic (s.l.) material of mantle derigranulite terrain merely represented a deeper vation at around 2900 Ma resulted in a general crustal level of the greenstone-granite terrain. crustal thickening to around present-day values. Against this view it was pointed out that apart According to Condie (1973) the mean present-day from obvious differences in structural style there thickness of Archaean crustal segments is 38 km, were characteristic differences in lithology be- which is the same as the overall mean continental tween the two terrains—specifically the presence crustal thickness. Condie's estimates of crustal of layered basic-ultrabasic igneous complexes thickening during greenstone-belt development, with prominent anorthosites, and supracrustal based on the chemistry of the volcanics, are 15 to assemblages dominated by quartzite, pelite and 30 km at South Pass in the SW Superior Promarble thought to be atypical of greenstone belts. vince, and 5 to 20 km at Barberton in the KaapThe more recent models of Windley & Smith vaal craton. These figures however, may relate (1976) and Tarney & Windley (1977) ascribe the only to regions of crustal thinning and offer little high-grade complexes to tectonic and magmatic guide to the overall crustal thickness during the activity at an active Cordilleran-type continental Archaean. Assuming that the surface of a typical margin. The granite-greenstone terrains in this greenstone-granite terrain represents an original view represent a lateral equivalent which might in depth of 5-10 km, the general crustal thickness in time be transformed into the high-grade type if the late Archaean may have been in the range the subduction orogeny advanced inward across 40-50 km. Pressure estimates from the high-grade terrains (e.g. O'Hara, 1977; Wells, 1976; Dickinthe Archaean continent. & Watson, 1976) suggest that some parts of There are several examples of "transitional" son were formed at depths of 50 km or terrains known where one type grades into the these terrains If the present crustal depths are added to other or which show characteristics intermediate more. these figures, we have to consider the possibility between the two. The Archaean crust in Finland the Archaean crust was locally in excess of 80 has been re-interpreted (Gaal et al., 1978) as a that granite-greenstone terrain in amphibolite fades; km thick. However, since the material presently
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underlying the high-pressure rocks may have been emplaced subsequently to their formation by tectonic or magmatic underplating (possibly even in Proterozoic times) or at least concurrently with erosional stripping, it is not strictly necessary to invoke larger crustal thicknesses than are indicated by the actual (geobarometric) rock evidence. Although it had been previously assumed that Archaean temperature gradients would have been considerably greater than those of today, several authors have convincingly argued that the pressure-temperature data for the later Archaean are not inconsistent with present-day or only slightly higher gradients (Bickle, 1978; Burke & Kidd, 1978; England, 1979). Since it is estimated that overall Archaean heat flow was likely to have been 2-3 times present continental values, these authors argue that the additional heat is likely to have been lost by faster creation of oceanic lithosphere—with important implications for Archaean "plate" tectonics. Thus although Archaean crustal thicknesses at their maximum are unlikely to have been much less than those in present-day orogenic belts, lithosphere thicknesses may have been consider T ably reduced. It is possible that the oceanic lithosphere may have been only 30-40 km thick— about half to two-thirds its present thickness in the late Archaean (Baer, 1977; Bickle, 1978) with a corresponding reduction in the upper mantle part of the continental lithosphere. Condie (1973) estimates an average depth of 85 km to the magma source (subduction zone?) for greenstone-belt vulcanicity (compared with the 100250 km range today) which implies a considerably thinner continental lithosphere. An estimate of 80-90 km is derived by Windley & Davies (1977) from the spacing of volcanic complexes in the Abitibi greenstone belt. TECTONIC MODELS AND MECHANISMS Windley (1976) lists six mechanisms suggested as explanations for the evolution of the Archaean crust: (1) "microcontinental collision"— cf. modern plate tectonics (Windley), (2) "downsagging basins" (Glikson, 1970), (3) "basalt underplating" (Fyfe, 1974), (4) "subcrustal accretion" (Holland & Lambert, 1975), (5) "skimming orogeny" (Talbot, 1973) and (6) "collisional interthrusting" (Bridgwater et al 1974; Myers, 1976). In terms of their tectonic effects, these mechanisms may be divided into two main types: the orthodox uniformitarian type (e.g. 1, 5, 6) based on present plate-tectonic processes, where the tectonic effects are attributable mainly to sub-
duction or collision, and a non-uniformitarian type (e.g. 2, 3, 4) where the tectonic effects are produced by crustal and/or sub-crustal flow governed by gravity-driven or convective circulation. In the latter category, three specific mechanisms are discussed: gravity-spreading, mantledecoupling, and thinned-crust collision. Combinations of models are of course not excluded. The Uniformitarian Plate-tectonic Model The uniformitarian plate-tectonic model was originally proposed by Bridgwater et al. (1974) to explain the tectonics of the North Atlantic craton. Bridgwater et al. suggested that the piling up of nappes together with the injection of concordant Nuk granite sheets took place at a continental margin, giving rise to considerable crustal thickening to produce a stable continental mass. This process was seen as marking the end of an earlier mobile phase characterised by thinner crust, greater thermal activity and higher mobility. The nappe-stacking process was thought to be driven by subduction at an oceanic margin with oceanic supracrustals being interleaved with continental basement on a "series of subduction zones" . This model has been elaborated upon in subsequent papers (e.g. Burke et al., 1976; Tarney et al., 1976; Windley & Smith, 1976; and Tarney & Windley, 1977). The mechanism is discussed further by Tarney & Windley (1977). They claim that the high-grade gneiss terrains in general represent crust thickened by the underthrusting which lifts up the continental margin by the successive emplacement of younger oceanic segments at its base. The underthrusting of ocean crust causing nappe stacking and imbricate inter-thrusting is associated with the generation of extensive calc-alkaline tonaliticgranodioritic magmas. Thus according to Tarney & Windley the highgrade Archaean complexes are ascribed to subduction tectonics at an active Cordilleran-type continental margin. They conclude that a "substantial" proportion of the continental crust evolved rapidly during the late Archaean as a result of these thickening processes culminating in the formation of deep-seated granulites. The importance of crustal thickening in the formation of granulite-facies rocks has also been emphasised by Martignole (1979) but he regards the belts of Precambrian granulites as zones of crustal collision. The greenstone-belts are attributed to marginal basins developed by back-arc spreading arising from the subduction process and subsequently closed by crustal compression in a way similar to the Mesozoic example in-Chile described by
HORIZONTAL STRUCTURE IN HIGH-GRADE TERRAINS
Tarney et ml. (1976). Thus all the main features of both types of Archaean terrain are explained by a simple subduction model. Although plate collision was an aspect of former plate models it does not figure in this model. Subduction tectonics therefore, rather than collision tectonics, is the driving mechanism. The main structural problem with the subduction model is that a dipping subduction zone cannot explain horizontally elongated strain fabrics. As shown in Figure 1, a subduction zone dipping at 45 ° would produce flattening fabrics even at very high strains with dips of more than 50°. Clearly subduction dips may have been very much lower, especially with faster Archaean spreading rates, and low dips on present-day subduction zones have been observed, but these usually steepen with depth. It is difficult to reconcile the process of subduction as currently understood with a horizontal simple shear stress at the base of the crust when at the very least the subducting slab must descend from the surface to the base of the lithosphere.
485
The Gravity-spreading Model Non-uniformitarian models have achieved much less prominence in the recent literature. Nevertheless a model which does not rely on subducting lithospheric slabs offers some important advantages over uniformitarian solutions in view of the possibility that subduction as currently understood did not operate during the early to middle Precambrian (cf. Baer, 1977). The gravity-driven diapiric model for greenstonegranite terrains has already been discussed. This model has not yet been extended to the highgrade gneiss terrains. However, it is a logical consequence of the application of the model to greenstone belts that there should exist in the lower regions of the crust below the dense greenstone layer a regime of sub-horizontal mass flow in the "source layer" for the rising granitic domes (see Fig. 2A). According to the experimental work of Ramberg (1967, 1973) the material flow in this source layer results in lateral extension along flow lines pointing inward toward the domes and lateral compression in the root zones of the domes (Fig. 2B). The strain patterns in such a situation would be complex. Assuming circular diapiric domes, there would be constrictional strains along radial lines pointing inward toward the domes becoming larger inwards (Fig. 2C). Superimposed on this pattern would be a simple shear effect due to the effect of increasing flow away from the (stationary) margins of the layer. The regional structural pattern would be dominated by the wavelength of the diapirs. Various attempts been made to estimate this, yieldFig. 1. Subduction tectonics: Strain pattern pro- ing valueshave in the range 22-25 km (Ramberg, 1972; duced at the margin of a continent by a sub- Fletcher, 1972; Stephansson & Johnson, 1976). ducting slab, assuming a simple asymmetric These are not based on Archaean examples, howshear zone. ever, and the author's observations in the Superior Province suggest a figure of around It is of course possible that the subduction 50 km. angle could become shallow at depth. This possi- This model therefore would lead to a radial bility ("flat plate" subduction) has been sug- pattern of horizontal constrictional strains pointgested as an explanation of the Cape fold belt of ing inwards to regions of horizontal compression South Africa by Lock (1980). It could be argued with radially arranged upright folding, repeated that because of the "unsinkable" nature of at intervals of, say, 50 km. Such a pattern has not Archaean lithosphere, any "subduction" that yet been identified in Archaean high-grade tertook place would have to be shallow (cf. Drury, rains but ought to be easily recognisable. It is 1977). However, it is difficult to visualise exten- doubtful whether the structure of the Godthaab sive sheets of thin oceanic lithosphere penetrating craton could be interpreted in this way, but the very far horizontally below the continental litho- model offers a convenient source of large consphere while retaining their coherence and being strictional strains difficult to explain by other able to transmit shear stress. Such a model has means. more affinity with the mantle-decoupling model discussed below and should not strictly be re- Mantle Decoupling Horizontal mantle flow was originally suggarded as "subduction" tectonics. It seems to the writer that subduction models as gested by Holland & Lambert (1969) as a method presently conceived are unconvincing explana- of creating the sub-horizontal crustal flow believed by them to be responsible for the flat-lying tions of Archaean tectonic patterns.
Fig. 2.
Gravity Spreading: A. Diagrammatic crustal section after Ramberg (1967) showing the rise of lighter granitic diapirs and the sinking of a denser greenstone layer. The flow in the granitic source layer results in extension below the sinking greenstones and compression in the roots of the diapirs. B. Incremental strain ellipses for the source layer in section, showing the effect of flattening strain and simple shear in the layer, together with increasing constrictional strain toward the roots of the diapir. C. Plan view of the flow in the source layer showing the effect of increasing constrictional strains toward the root zones of two adjacent diapirs. Incremental strain ellipses change from flattening through horizontal constrictional to vertical constrictional.
early Scourian structures in the high-grade granulite terrain of NW Scotland. The process implies that the upper part of the mantle is effectively "decoupled" from the crust. Horizontal flow of mantle material with respect to a "stationary" crust will impose a horizontal simple shear stress on the base of the crust which, given a sufficiently ductile material, will become a horizontal shear zone (Fig. 3). Because of the lower viscosity of the base of the crust compared with the upper mantle, shear stresses applied to the base of the lithosphere would promote shear zone formation within the lower crust rather than within the lithosphere (see Yuen et al., 1978). With high strains, the planar fabrics in such a shear zone will eventually make a small angle with the horizontal. A horizontal shear zone produced by mantle decoupling is therefore a more acceptable method of explaining the observed structures than a dipping subduction zone. Since shear
zones are effectively ductile faults transmitting relative motion across the zone, it is appropriate to estimate the amount of motion or displacement which could be achieved given the likely constraints of the observed strains. The critical factors are the width of the zone and the value of the mean shear strain which measures the amount of deformation caused by the shear displacement. Observed crustal shear zones of the order of tens of km in width have mean shear strains in the range 4-6 (Watterson, 1979). In a study of late Archaean shear zones in the Superior Province, Park (1981) has shown, for 9 measured zones, a relationship of decreasing mean strain with increasing width. For a shear zone with 10km width, the mean shear strain was between 2 and 3 . In the lower crust, with lower viscosities and more homogeneous material, it is likely that mean shear strains would be greater. According to Escher & Watterson (1974) strain ratios at
HORIZONTAL STRUCTURE IN HIGH-GRADE TERRAINS
deeper levels in basement rocks commonly attain 50:1 which would indicate a mean shear strain of around 7. The shear-zone width is more difficult to estimate. The exposed width of horizontally layered gneisses in the Godthaab craton and in NW Scotland must be 5-10km, and 10km could be regarded as a minimum width for the shear zone. The maximum width is unlikely to be greater than 30 km given the probable crustal thicknesses of the time, and the domination of the upper 10-20 km by a different tectonic regime. If we assume a width of 20 km and a mean shear strain of 7, the displacement achieved is 140 km. Even doubling the shear strain would only double the displacement, implying that the tectonic constraints set limits to possible displacements of several hundred km. Such relatively small displacements, given the likely rates of convective motion (c/. Elder, 1976) would be achieved in a very short time, geologically. Using Elder's average figure of l l c m / y r , a displacement of 140km would only take 1.3 Ma. Thus it would seem that observed lower crustal strains are incompatible with large relative displacements of crust and upper mantle which must therefore have been effectively "coupled" for most of the time. (Decoupling at the base of the lithosphere on the other hand could have been much more important). The possibility of mantle decoupling taking place during the Archaean depends on the way that convection operated at that time, and in particular whether or not subduction (in the platetectonic sense) took place. It has been suggested by Baer (1977) that subduction did not occur until the lithosphere had become sufficiently thick for eclogite to form and provide the required negative buoyancy for the lithosphere to sink. He suggests that this was not likely until the late Proterozoic. Bickle (1978) argues, on the basis of likely Archaean temperature gradients in the oceanic lithosphere, for a thickness of around 40 km, which would be less dense than the mantle. So that either a more vigorously convecting mantle pulled down less-dense lithosphere in descending convection currents, or decoupling took place at the base of the lithosphere, with the less-dense material remaining at the surface and the denser cooler mantle material returning downwards (c/. Hargraves, 1978). If such decoupling took place, it would provide an attractive mechanism for producing horizontal shear zones in the lower crust (Fig. 3). The problem with a model which does not rely on subduction, is in providing an efficient enough method of losing heat through the oceans. The heat loss would be provided by puncturing the thin lithosphere with oceanic volcanoes which
487
would gradually thicken and expand the oceanic crust with concomitant shortening of the continents. The process would be akin to greenstonebelt formation on the continents—indeed such early Precambrian oceans might well have originated from intracontinental greenstone basins by a process of expansion and lateral squeezing of early-formed sialic crust—a process suggested many years ago by Ramberg (1964).
B t
CRUST
J_ _
LITHOSPHERE
MANTLE
Fig. 3. Mantle decoupling: A. Coupled crust and mantle spreading to provide an ocean. B. After some lithosphere thickening has taken place, mantle decoupling causes a shear zone to form in the lower crust. C. Relationship between strain and displacement in a lower crustal shear zone; strain ellipse orientation for a mean shear strain of 4.
In the absence of oceanic subduction, relative movements between crustal blocks would be accomplished by transform-type translations along steep shear zones and perhaps locally by over-riding on low-angle shear zones. Despite evidence for considerable ductility within the early Precambrian crust, uniform stress systems appear to have operated over quite large areas, particularly in the later Archaean, suggesting that the " b l o c k " concept is valid. It is suggested that a regular mantle convection system operating beneath a large number of small and variably ductile continental lithosphere blocks in mutual contact would require local decoupling at the base of some of the blocks in
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order to maintain geometric compatibility. The interactions of these blocks could explain much of the complex tectonic pattern seen in the Archaean. An analogy would be with the complex deformation pattern exhibited by the recent deformation of the Asian plate resulting from the collision with India (Molnar & Tapponier, 1975; Watterson, 1978). The size and strain patterns of the basal crustal shear zones set limits on the horizontal displacements which can be achieved at the surface, and restrict the amount of inter-plate movement. However, much larger displacements may take place within the more ductile oceanic and continental upper mantle, enabling more rapid convective mantle flow to take place without extreme crustal disturbance. Thinned-crust
Collision
Another method of obtaining sub-horizontal high-strain fabrics is by crustal thinning due to extension (Fig. 4). Models involving crustal thinning have been suggested by several authors to explain greenstone-belt formation and also figure in explanations of Proterozoic mobile belts (c/. the 4 'millipede'' model of Wynne-Edwards, 1976). An attractive aspect of the thinned-crust model is that later compression or collision can produce nappe piles or other complex compressional structures much more easily than if the whole thickness of the crust were involved. Thus for example it might be possible to explain the layered high-strain features of the high-grade gneiss belts as a result of extreme ductile thinning followed by compressive thickening by formation of a nappe pile consisting of inter-layered sialic basement and oceanic or quasi-oceanic crustal cover (Fig. 4C). Only the lower crust would behave in this way; the upper part may deform by reversing earlier extensional movements on listric faults (cf. Jackson, 1980). The actual mechanism responsible for the extensional thinning is envisaged as convectional mantle spreading. This could be the result of a local mantle diapir, a small convection cell, or of a larger-scale convection cell like that envisaged for the present plate model. However the requirements in the Archaean for faster heat dissipation are more likely to have been met by a large number of small oceans of perhaps more limited duration. The extent of the thinning is speculative, and complete separation of the continental crust may have occurred. The important feature of the model however is that only a few hundred km are needed to produce the observed tectonic effects. A crustal segment only 100km wide, if extended to 500 km would produce lower crustal strains of 1:25.
How the thinning would take place is difficult to predict, but since the extensional stresses are provided by lateral spreading, it is presumed that this would exert simple shear stress at the base of the crust which would result in the movement of material away from the extensional axis by oppositely directed shear zones. The warmed central block, being more ductile, might deform relatively homogeneously by pure tensile shear, or alternatively along low-angle shear zones with "normal"-sense displacements (see Fig. 4B). This model requires alternate crustal expansion and contraction but not necessarily subduction of lithosphere as required by plate-tectonic models. During the compressive phase, tectonic thickening might result in local melting at the base of both the crust and the lithosphere but the model does not require a subducting plate as such.
Fig. 4.
Thinned-crust
collision:
A. Commencement of extension above a spreading centre, with the development of listric faults in the upper crust. B. Formation of a small ocean basin above thinned lower crust. The upper crust has extended by normal movements on listric faults which may change at depth to low-angle shear zones with normal sense of displacement. Black ellipses show likely strains for such a model. C. Subsequent lithosphere compression produces crustal thickening by reversal of movements along the low-angle shear zones. Rapid alternations of mantle material, highly strained lower crustal gneisses and supracrustals are clearly possible in this model.
H O R I Z O N T A L S T R U C T U R E IN H I G H - G R A D E
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TERRAINS
A by-product of the model is its potentiality for thickening the Archaean crust. As pointed out by Bridgwater et al. (1974) and many others, it is convenient to have a mechanism for thickening the crust in the high-grade gneiss-granulite terrains to between 50 km and 80 km in order both to satisfy the pressure-temperature data and to explain why these areas are presently uplifted relative to the greenstone-granite terrains. According to this model, the latter would never have suffered compressive thickening to any significant extent, and their present crustal thicknesses—considerably less than the maximum thicknesses of the high-grade terrains—may have been achieved by primarily magmatic means. Structural evidence from the Superior Province suggests that quite large horizontal displacements on shear zones have resulted in 40% shortening in one direction probably balanced by comparable extension in the horizontal plane rather than by any very marked crustal thickening. Thus over an area of the size of the Superior Province, over 600 km of horizontal displacement is represented in the shortening (assuming no thickening). Displacements of this magnitude may well have accompanied the opening of ocean basins to the north or south and the simultaneous closing of ocean basins to the east or west.
thinned-crust collision. Subduction and mantle decoupling are mutually exclusive in the sense that the lateral stresses produced by mantle spreading are dissipated by subduction. Mantle decoupling alone cannot explain the tectonic mixing found in high-grade terrains between supracrustal and lower-crustal components, and for this reason the thinned crust model is preferable. However, mantle decoupling could be regarded as an additional tectonic complication affecting the thickened lower crust after collision. The gravity-spreading model is compatible with all three other models and may be regarded as a complementary intracrustal process. In the view of the writer there is no compelling evidence favouring the selection of any one of these models as the primary explanation of the lowercrustal tectonic pattern. However, on balance, the non-uniformitarian models of mantle decoupling and thinned-crust collision are more convincing explanations than subduction of highgrade gneiss tectonics. Gravity spreading or some other type of lateral diapiric flow may make a contribution to lower crustal tectonics in greenstone-granite terrains, and thus inferentially in high-grade terrains, but does not offer such an attractive explanation of large areas of horizontal highly-strained gneisses found, say, in the North Atlantic craton.
CONCLUSIONS Four basic models have been discussed in terms of their likely effects on lower-crustal tectonics. They may be termed respectively subduction, gravity-spreading, mantle decoupling and
ACKNOWLEDGMENTS The author is indebted to A. J. Baer for his critical reading of the manuscript which resulted in several improvements to this paper.
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grade complexes and modern continental margins. Nature, Lond., 260, 671-675. WYNNE-EDWARDS, H. R., 1976: Proterozoic ensialic orogenesis: the millipede model of ductile plate tectonics. Am. J. Sci., 276, 927-953. YUEN, D . A . , FLEITOUT, L . , SCHUBERT, G . , & FROIDE-
VAUX, C., 1978: Shear deformation zones along major transform faults and subducting slabs. Geophys. J.R. astron. Soc., 54, 93-119.
EVOLUTION OF ARCHAEAN CRUST IN THE EASTERN GOLDFIELDS PROVINCE OF THE YILGARN BLOCK, WESTERN AUSTRALIA N. J. Archibald, L. F. Bettenay, M. J. Bickle & D. I. Groves Department of Geology, University of Western Australia, Nedlands, Western Australia 6009 ABSTRACT The controversy concerning the crustal evolution of the granitoid-greenstone terrain of the Eastern Goldfields Province results largely from lack of critical original relationships between major components in this poorly exposed terrain for which limited geochronological data are available. However, studies of the better-preserved and less-equivocal late tectono-metamorphic events provide sensitive tests for models of crustal evolution, particularly when constrained by present crustal structure. Such studies produce compelling evidence against simatic greenstone formation and derivation of the granitoid terrains from melting of greenstones at deeper crustal levels. Models involving formation of greenstone belts in relatively shallow {ca 10 km) ensialic basins and derivation of intrusive granitoids from a pre-existing granitic (.s.l.) source satisfy most constraints. Available isotopic data do not unequivocally define the source of granitoids, although they do indicate that the widespread banded gneisses, the most likely basement among exposed rocks, could only have formed shortly before greenstone deposition. The mechanisms for the high-strain, early deformation of these gneisses are not known although a horizontal tectonic regime is suspected: this may have involved greenstone belts at higher crustal levels and was possibly responsible for crustal thickening. The characteristics of greenstone metamorphism and derived geothermal gradients appear to result, at least in part, from differential uplift of lower crustal layers to higher levels following some crustal thickening: they could not be produced by uprise of granitoid magmas alone. The tectonic features of the terrain are not directly analogous with modern plate-tectonic settings, for example oceans or back-arc basins. However, the existence of vast volumes of granitic rocks in the early crust without signs of any mafic/ultramafic residue suggest operation of a subduction-like process, and heat-flow constraints imply oceanic areas.
INTRODUCTION Models for development of Archaean granitoid-greenstone terrains of the Eastern Goldfields Province differ in basic aspects (e.g. Glikson, 1972, 1976a, b, c, 1978; Glikson & Lambert, 1976; Williams, 1974; Gemuts & Theron, 1975; Binns et al., 1976; Groves etal., 1978; Gee, 1975; 1979a). In part these differences arise from limited data, but they also hinge on the interpretation of geological relationships. Although disputed problems include the source of the large areas of granitoids and the stratigraphy of the greenstone sequences, the key factor in the evolution of the Eastern Goldfields Province relates to the tectonic setting of the greenstones: were the sedimentary-volcanic sequences deposited on an older sialic crust or did they develop as a simatic crust analogous to modern oceanic crust? Direct interpretation of the early tectonic setting of greenstones in the Province is hampered by the lack of resolution of Spec. Pubis geol. Soc. Aust., 7 (1981)
geochronological techniques, and structural complexity in poorly exposed and inadequately mapped terrains. Notwithstanding these, the Yilgarn Block has yielded important new data on Archaean metamorphism (Binns et al., 1976) and its large size may enable discrimination of regional variations not apparent in smaller cratons. Furthermore, study of better-preserved and less-equivocal late tectonic events, such as granitoid intrusion and the metamorphic episode, may provide more sensitive tests of models than direct attempts to unravel the early events. This is the approach we mainly follow here. However, the ultimate test for conceptual models rests on how adequately they account for the synthesis of geological characteristics within the terrain. The purpose of this paper is to review certain aspects of crustal development in the Eastern Goldfields Province, and to evaluate some of the models of Archaean greenstone belt formation in the light of available knowledge.
492
N. J. ARCHIBALD, L. F. BETTENAY, M. J. BICKLE & D. I. GROVES EASTERN GOLDFIELDS PROVINCE GRANITE-GREENSTONE TERRAIN
SOUTHWESTERN PROVINCE HIGH-GRADE GNEISS TERRAIN 200 km
?
?
• Darling Fault
.-Present land surface
?
?
?
?
? '
?
? ?
?
? ?
?
? ?
?
- - - __ _ ^ r ' a c e ? ? ? ~~ ~~ — — —
? ?
?
?
/Southern Cross ? '
?
?
?
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"
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Coolgardie
I i
t + + + + + + + + + + +2.78g/cc; + + + 6.12km/sec + + + + •
+ + + + + + + + + + + +
2.94g/cc; 6.66km/sec ,-Z-Z-Z-Z-Z-Z:
! 3.45g/cc; 8.25km/sec ;
"40 -50
Phanerozoic sediments
j-^ + ^-j Upper crustal layer
Archaean overburden removed by subsequent erosion
f ; ~ ~ - Z - l ] Intermediate crustal layer
Greenstone belts
Lower crustal layer
Schematic uppermost layer enriched in K, U and Th
Upper mantle
km
Fig. 1. Crustal section across the Yilgarn Block between Perth and Kalgoorlie based on data from Fraser (1974), Mathur (1974), Mathur et al. (1977), Sass et al. (1976) and Lambert (1971). The section shows the Archaean greenstone belts in relation to major crustal features of the region (note that the vertical and horizontal scales are not equal).
PRESENT CRUSTAL STRUCTURE OF THE YILGARN BLOCK The crustal structure of the Yilgarn Block has been investigated by regional gravity surveys (Fraser, 1974), deep seismic refraction (Mathur, 1974), heat-flow (Hyndman et al1968; Sa&s et al., 1976) and geochemical studies (Lambert & Heier, 1968a, 19686; Lambert, 1971). A threelayer stratification (Fig. 1) is interpreted to extend beneath both granitoid-greenstone and granulite-gneiss terrains, with a similar crustal layering characterising the northern (Drummond, 1979) and southeastern (Mathur et al., 1977) extremities of the Yilgarn Block. The average crustal composition under granitoid-greenstone terrains is probably acid to intermediate (cf. Drummond, 1979) with granitoids dominant in the upper layers and granulites at the base. The heat flow and geochemical data imply a corresponding compositional stratification, with an upper few-kilometres thick ' 'granitic" layer enriched in heat-producing elements (Sass et al., 1976). The proportion of greenstones is inferred to decrease markedly at depth./This is supported by gravity data (Peters, 1972; A. Flavell, in Archi-
bald et al., 1978; A. Flavell and P. Joppek, pers. comm., 1976; Constable, 1978) which imply that granitoids underly the greenstones at maximum depths of ca lOkms; that is, the greenstone belts are shallow structures restricted to upper crustal layers (see Fig. 1). Finally, the gravity data suggest that mafic granulites represent at best an insignificant component at the base of greenstone belts as is consistent with their restricted outcrop. They are confined to the anomalous lower crustal layer beneath the granulite-gneiss terrains which is virtually absent beneath the granitoid-greenstone terrains. Given the apparent crustal structure, it appears that the older (>3000-3300 Ma) gneisses of the western Yilgarn (e.g. Arriens, 1971; Nieuwland & Compston, 1980) extend as a component of the layer underlying the granitoidgreenstone terrains of the Eastern Goldfields Province. GRANITOIDS OF THE EASTERN . GOLDFIELDS PROVINCE Subdivision Granitoids which outcrop within the southern Eastern Goldfields Province have been sub-
493 ships have not been identified. Much attention has also been focused on the widely dispersed banded gneisses and their migmatised equivalents (Fig. 2) as relics of modified basement to the greenstone sequences (Archibald & Bettenay, 1977; Archibald et al., 1978; Bettenay, 1977). These rocks occur within regional complexes, as enclaves within and marginal to post-kinematic plutons and as phases in some synkinematic diapirs. Despite their large areal distribution, they are relatively uniform in composition (Fig. 3), being more siliceous, lower in mafic content, dominantly hornblende-free, and less tonalitic than most documented Archaean gneisses (e.g. Condie & Hunter, 1976; Bridgwater & Collerson, 1976; Dougan, 1976; Phaup, 1973). Archibald & Bettenay (1977) and Archibald et al. (1978) consider that these gneisses preserve deformation fabrics earlier than the earliest deformations recorded within the greenstone sequences and therefore conclude that the gneisses predate the greenstones. Alternatively Glikson (e.g. 1978) proposes that these more complexly deformed gneisses were derived by partial melting of greenstones and that their deformation reflects processes involved in their diapiric uprise through the crust. However the distribution and structural state of the gneisses is not so easily reconciled with diapir-associated deformations. The highly strained gneisses are not found in areas near the "hat" of synkinematic diapirs, that is in areas where highest strains would be expected (e.g. Dixon, 1975), and in addition they exhibit a fabric which predates the fabric geometrically related to the diapirs (Fig. 4). Available geochronology is insufficient to resolve the basement problem. The Rb-Sr isochron age of the gneisses (2670 ± 60 Ma) is within error of the ages of granitoids which clearly intrude the greenstones (Chapman et al., 1981). The initial S r / Sr ratio of the gneisses (0.7026 ± .0004, 26) confines their crustal prehistory (cf. Moorbath, 1978) to less than ca 200 Ma if there is assumed to have been no change in their average Rb/ Sr ratio. Despite this, there are other lines of isotopic evidence which suggest an older granitoid basement. Mineral and whole-rock Pb isotope data from some intrusive granitoids have been interpreted to indicate crustal prehistories extending back to ca 31003300Ma (Oversby, 1975). Similarly two postkinematic adamellite plutons in the central Yilgarn have rather high initial Sr/ Sr ratios (0.71 and 0.74, Chapman et al., 1981) and these ratios might either reflect melting of older gneisses (3000-3600 Ma) or rehomogenisation of Sr isotopes by post-intrusive metamorphic events.
C R U S T A L E V O L U T I O N IN E A S T E R N G O L D F I E L D S , W . A .
divided structurally by Bettenay (1977), Archibald & Bettenay (1977), and Archibald et al. (1978) into banded gneisses plus migmatised equivalents, intrusive foliated synkinematic granitoids, unfoliated post-kinematic granitoids and minor fractionated leuco-adamellites (Fig. 2). Much of the area underlain by granitoids is poorly exposed and the relative proportions of these granitoid types is uncertain; outcrops of post-deformational phases predominate but this may reflect differential weathering. The wide distribution of small outcrops of banded gneiss and migmatite, the abundance of gneiss xenoliths in intrusive granitoids, and heat flow/heat production arguments suggest that the banded-gneiss component predominates within the upper "granitic" layer at the present erosional level: the heat flow is substantially less than that expected if the upper crust mainly comprised intrusive granitoids (Table 1). Several recent regional studies (e.g. Bettenay, 1977; Gee, 19796) comment on the rarity of magmatic tonalites in outcrop so that, although the possibility advanced by Glikson (e.g. 1978) that such phases constitute a major but poorly exposed component is impossible to evaluate, it is certainly not favoured by the available evidence. TABLE I
Comparison of estimated K, U and Th in the uppermost layer of the Yilgarn crust, with mean elemental abundances in granitoid types of the southern part of the Eastern Goldfields Province ,
% K O)
3.1 3
k
A:
2
th (ppm)
U (ppm)
20
3
16
2-3*
40
8
lit 67 5 Estimated elemental abundances in the uppermost "granitic" layer of the Yilgarn Block, based on heat-flow and geochemical data (Lambert & Heier, 1968a, 19682?).
B:
11 banded gneisses from gneiss belts and migmatite complexes.
C:
28 representative post-kinematic intrusive granitoids.
D:
6 fractionated leuco-adamel1ites, including typical "K-rich" granites of Glikson & Lambert (1976) such as Mungari. Analytical data for B, C, D from Bettenay (1977), K determined by A.A.S., U and Th by XRF on duplicate pressed powder mounts.
* Uncertainty in mean resulting from s a m p l e s beneath detect ion 1imit (2 ppm).
Banded Gneisses If greenstones were formed above older sial, relicts of this basement may be preserved within the granitoid component of the crust. Interpretations of granitoid-greenstone relationships are controversial and, although some studies (e.g. Cox & Tyrwhitt, 1975; Theron, 1974) speculatively identify candidates for pre-greenstone basement, preserved basement-cover relation-
87
86
87
86
494
N. J. ARCHIBALD, L. F. BETTENAY, M. J. BICKLE & D. I. GROVES
Fractionated leucoadamellite Post-kinematic granitoids forming discrete plutons
Syn-kinematic granitoids
Banded gneiss and migmatite
• ^ ^
Concealed banded gneiss and migmatite. Sparse outcrop dominated by postkinematic phases Greenstone belts
Fault
N 100 km
Fig. 2. Interpretative regional map of granitoids in the southern part of the Eastern Goldfields Province showing distribution of kinematic types and banded gneiss-migmatite complexes (adapted from Bettenay, 1977).
There is evidence that rehomogenisation of Sr isotopes occurred adjacent to pegmatites as recently as 2300Ma (Chapman et al., 1981). Thus the isotopic evidence hints at a much older basement but suggests that the exposed banded gneisses are not part of this much older crust or, if they were, have been so modified by the addition of Rb at 2700 Ma that their prehistory is undetected. If this modification is proposed then it is necessary to postulate that the gneisses had very low Rb/Sr ratios (-0.05) characteristic of depleted granulite-facies terrains, for which there is currently no evidence. The gneisses may still have formed part of a pregreenstone basement and the errors on the isotopic ages ± 100 Ma) are great enough to allow formation and deformation of a gneiss terrain prior to greenstone deposition if this occurred shortly before metamorphism. The structural setting required for the formation of the gneisses is intriguing given the regional extent of the high-strain episode required. In some better-exposed gneiss terrains such strain has been related to horizontal tectonic regimes (Myers, 1978; Bridgwater et al1974) and there is evidence of similar tectonism in some grani-
toid-greenstone terrains (e.g. Coward et al., 1976; Bickle et al., 1980). It is therefore tempting to infer by analogy with other gneiss terrains that a horizontal tectonic regime existed in the Eastern Goldfields Province before or during the eruption of greenstone volcanic sequences. In this context it should be noted that the poorly-preserved macroscopic early isoclinal (?nappeforming) event recorded in the low-strain domains of some greenstone belts (e.g. Archibald et al., 1978) could represent a higher-level, non-fabric forming reflection of a horizontal regime that produced strong fabrics in more ductile rocks at deeper crustal levels. However, insufficient data are available to constrain these speculations. Intrusive Granitoids Widespread granitoid plutonism followed deposition of at least some, if not all, greenstone belts, and models for granitoid genesis invariably depend on preconceptions of the earliest granitegreenstone relationships. For instance, Glikson (1976#, b\ 1978) proposes that all granitoids in the south-eastern Yilgarn Block, including the banded gneisses, post-dated development of a major Lower Greenstone" sequence and belong
CRUSTAL EVOLUTION IN EASTERN GOLDFIELDS, W.A.
to (i) an older Na-rich tonalite-granodiorite group characterized by the Kambalda pluton and derived directly by partial melting of basic ''Lower Greenstone" material, or (ii) a younger K-rich adamellite group typified by the Mungari pluton (a small intrusive some 10 km NE of Coolgardie) and derived by re-melting of the older granitoid suite. The compositional evolution of such granitoids should be distinctive if the earliest bodies were derived by melting basic crust in the absence of sialic material and Glikson maintains that the Na-rich group exhibits these characteristics. However, the distribution and character of the intrusive granitoids as described by Bettenay (1977) and Archibald et al. (1978) does not bear out either the compositional range claimed by Glikson or the secular variation within this range. Most granitoids evolved as synkinematic or postkinematic plutons during the interval 2700 to 2500 Ma and available geochronology is unable to resolve age differences during this intrusive event (see summary in Chapman et al., 1981).
495
Post-kinematic fractionated leuco-adamellites with younger Rb/Sr whole-rock ages are rare and the Rb-Sr systematics in some of these plutons may have been reset 100-200Ma after intrusion. Therefore, with the available data, the only chronological division applicable to the vast majority of intrusive granitoids is into the broad, transitional synkinematic and post-kinematic groups. The synkinematic and post-kinematic phases exhibit similar ranges of composition with little evidence for advanced magmatic fractionation and no statistically significant compositional difference between the two groups (Archibald et al., 1978—Table 1). The regional studies by Bettenay (1977) have revealed that the type examples of Glikson's "Na-rich" and "K-rich" granite subdivisions (Kambalda and Mungari) are not representative of any areally significant group. Mungari represents the distinctive group of postkinematic, fractionated leucoadamellites which have a characteristic trace-element pattern (Bet-
50 CaO
Post-kinematic and synkinematic intrusive granitoids from the Kalgoorlie-Southern Cross region •
Kambalda "granodiorite"
•
Coolgardie-Norseman " Na-rich granites" after Glikson (1978) Mungari pluton
o
Coolgardie-Norseman "K-rich granites" after Glikson (1978) Field containing 90% of banded gneisses from gneiss-migmatite complexes Na20
70 K 2 0
fig. 3.
80 Na 2 0
Composition of Archaean granitoids and gneisses from the area shown in Figure 2. The granitoids attributed to "Na-rich" and "K-rich" granite groups by Glikson (1978) are shown in relation to variation shown by all granitoids of the terrain. Analyses from Bettenay (1977); data for Mungari and Kambalda from Oversby (1975).
496
N. J. ARCHIBALD, L. F. BETTENAY, M. J. BICKLE & D. I. GROVES
tenay, 1977) and compose less than 0.05% of out- Dixon, 1975; Schwerdtner etal., 1978), the ligacropping granitoids. The Kambalda granitoid is tions in both the dome and surrounding greenakin in both occurrence and composition to the stones are inferred to have formed under condisodic porphyry intrusives which may constitute tions of solid-state radial flow and as such reprean integral component of the felsic volcanic sent the maximum principal extension direction sequences within the greenstone pile (cf. of the finite strain ellipsoid. The orientation of O'Beirne, 1968; Hallberg et al., 1976). In addi- the lineations indicate that the exposed structural tion, Glikson claims that the Pioneer and Widgie- level of the Widgiemooltha Dome is the hat of a mooltha Domes (Fig. 4) are members of the "Na- diapir whereas that of the Pioneer Dome (see Fig. rich" suite and the Spargoville Dome (S of Cool- 5) is closer to the base, in agreement with higher gardie, see Fig. 2) is a member of the "K-rich" metamorphic grades in greenstones surrounding suite. These suppositions are not upheld by the the latter dome (summarised in Fig. 6). available petrographic and geochemical data (see Fig. 3). All synkinematic diapiric plutons intruding greenstones are dominantly granodiorite/ adamellite in composition and cannot be distinguished on geochemical criteria from the postkinematic major phases. WIDGIEMOOLTHA DOME There are further lines of evidence inconsistent with derivation of intrusive granitoids from melting of basic components of greenstone belts. Bettenay (1977) has presented geochemical and petrographic evidence consistent with derivation of the intrusive granitoids by partial melting of banded gneisses (cf. Gee, 1979b), a process that Foliation in greenstones (S ) occurred with limited magmatic fractionation, Bedding minor contribution of magmas produced at deep Coincident bedding and foliation crustal levels, and restricted upward movement Lineation plunge of the water-rich partial melts. Phase relations Foliation in granitoid within the granitoids, petrographic observations Theoretical maximum extension (Dixon, 1975) of crystallization sequence and the mesonormative composition of granitoids are all consistent with their derivation and crystallization at pressures of 4-5 kb and temperatures between 650 °C and 680 °C (see Bettenay, 1977 based on experimental data presented by Winkler et al., 1975; Wyllie et al., 1976 and Winkler, 1976). These conditions are constrained by the stability of quartz-muscovite in gneisses and the absence of the reaction products that would result from Fig. 4. Structural map of the Widgiemooltha Dome higher-temperature, water-deficient melting inshowing the main linear and planar elements developed in granitoids and adjacent greenvolving the breakdown of hydrous minerals such stones. The section shows a comparison beas biotite (e.g. Winkler, 1976). Melting at tween observed structure and theoretical pressures much greater than 6 or 7 kb would promaximum extension directions for model duce melts with higher modal anorthite/orthodiapirs (after Dixon, 1975). clase and albite/quartz ratios than is observed. This inference that the intrusive granitoids were derived as water-saturated melts which crystallised close to their source is supported by GREENSTONE BELTS structural evidence that the synkinematic diapirs The characteristics of volcanic and sedimenwere emplaced by dominantly solid-state defor- tary components of the greenstone belts of the mation processes. The Widgiemooltha Dome Eastern Goldfields Province are summarised by (Fig. 4) possesses a subhorizontal lineation sub- Gemuts & Theron (1975), Gee et al. (1976) and parallel to the long axis of the dome. A similar Binns & Marston (1976). The greenstones form a lineation is also present at either extremity of the series of NW-NNW-trending belts which have dome within greenstones which were deformed been grouped into three subprovinces on the basis during diapiric uprise. On the basis of compari- of contrasting sediment and volcanic components son with modelling experiments on diapirs (e.g. and tectonic style (Williams, 1974). A number of 3
C R U S T A L E V O L U T I O N IN E A S T E R N G O L D F I E L D S ,
attempts have been made to establish stratigraphies within various belts (e.g. Williams, 1969; Gemuts & Theron, 1975), but the lack of continuous exposure and structural complexity (Archibald et al., 1978; Piatt et al., 1978) make evaluation of the validity of such stratigraphies difficult. For example, estimated stratigraphic thicknesses are unrealistically high in terms of structural depths inferred from gravity data and metamorphic grade of exposed sequences. Metamorphic studies over the entire Eastern
W.A.
497
Gold fields Province (Binns et al., 1976) indicate a range of prehnite-pumpellyite to high amphibolite fades, the highest grade and highest strain rocks being characteristically confined to linear zones on present greenstone margins (Archibald et al., 1978). However, detailed structural studies are limited to the widely separated Agnew (Piatt et al., 1978) and Widgiemooltha-Norseman (Archibald, 1979) areas. The pattern of deformation and metamorphism in the latter area (Fig. 5) involved an early phase of essentially pre-
Proterozoic basic intrusions GREENSTONES - METAMORPHIC FACES
:Widgiemoolthai
High-strain mid-high amphibolite High-strain mid amphibolite Intermediate-strain low amphibolite
Widgiemooltha
Low-strain low amphibolite
Dome+\ • + + \ + + +\
!{r>v7r
Low-strain greenschist GRANITOIDS Mainly post-kinematic
[+> + +[ Synkinematic diapirs ^\ Eundynie
Banded gneiss DEFORMATION 4 •+- F4 folding 2 ^
F3 folding ?2 folding F1 folding "UPPER" AND "LOWER" GREENSTONES Boundary according to Glikson (1978) adapted from Gemuts and Theron (1975) - younging direction shown
Pioneer, Dome
20 km Norseman Fig. 5.
M a j o r m e t a m o r p h i c a n d s t r u c t u r a l f e a t u r e s of t h e W i d g i e m o o l t h a - N o r s e m a n a r e a s h o w n in r e l a t i o n t o t h e p r o p o s e d u n c o n f o r m i t y b e t w e e n " U p p e r " a n d " L o w e r " g r e e n s t o n e s ( G l i k s o n , 1978). N o t e t h a t m e t a m o r p h i c b o u n d a r i e s cut this " u n c o n f o r m i t y " a n d t h a t all f o l d i n g p h a s e s (including F , ) a r e r e p r e s e n t e d in both " U p p e r " and " L o w e r " greenstones.
498
N. J. ARCHIBALD, L. F. BETTENAY, M. J. BICKLE & D. I. GROVES
metamorphic recumbent folding (Dj) upon which was superimposed two essentially coaxial major phases of regional deformation (D2 and D3). The metamorphism was in part coincident with the D3 phase although the thermal peak normally outlasted deformation. Diapiric granitoid domes were emplaced into the greenstones along zones that now represent major greenstone belt margins and correspond in part to the subprovince boundaries: local deformation (D4) accompanied diapirism. Fundamental aspects of greenstone-belt evolution that are contentious include: (i) the age of the greenstones, particularly the oldest sequences (ii) the pre-deformational nature and shape of the greenstones (iii) the concept of distinct cycles of greenstones separated by a major tectono-metamorphic break (e.g. Glikson, 1976a, b) (iv) the disposition of the high-strain marginal zones at depth. Do these zones continue under low-grade greenstones or are they essentially linear belts restricted to marginal zones? Geochronological data relating to greenstones are dominated by Rb/Sr studies (e.g. Turek, 1966; Roddick et al., 1976; Cooper et al., 1978; Chapman et al., 1981) which indicate greenstone metamorphism at ca 2650-2550Ma (i.e. overlapping the Rb/Sr ages of intrusive granitoids). Pb/Pb and Sm/Nd studies (Oversby, 1975; McCulloch & Wasserberg, 1978) are limited and the results cannot yet be unequivocally interpreted. Archibald et al. (1978) speculate that the 2760 ±70 Ma Pb/Pb mineral age for sodic porphyries of the Kambalda Dome (Oversby, 1975) could represent the age of at least part of the Kalgoorlie greenstones. The present subparallel, elongate shape of the greenstone belts could be interpreted as reflecting either individual depositional basins of variable extent and thickness or as more continuous sequences that were subsequently disrupted, for instance by forcible emplacement of the granitoid batholiths. The former interpretation is more consistent with regional contrasts in lithology and stratigraphy that coincide with the present NNWtrending greenstone distribution (i.e. the subprovince subdivisions of Williams, 1974). These broad subdivisions are supported by more detailed comparison of the greenstone sequences between subprovinces, for example the contrast between the Southern Cross (Gee, \919b) and Kalgoorlie-Norseman (Gemuts & Theron, 1975) areas. The pattern of linear zones of high-strain, high-grade metamorphism and spatially associated diapirs that partly coincide with subprovince
boundaries are best explained if these boundaries represent major interfaces such as depositional trough margins (e.g. Archibald & Bettenay, 1977; Archibald et al., 1978). We thus favour the model first proposed by Williams (1974), in which the Kalgoorlie subprovince represents a separate trough separated from, or flanked by, independently evolving greenstone sequences (see also Gee et al., 1980). This must however remain speculative awaiting detailed sedimentological analysis of the various greenstone sequences within and outside the subprovince. The concept of "Upper" and "Lower" Greenstones separated by a tectono-metamorphic hiatus (Glikson, 1976a, b) is impossible to discuss on a Province-wide scale as relevant data or type localities have not been given. However, the concept is not substantiated in the NorsemanWidgiemooltha area, cited as an example of "Upper" and "Lower" greenstone relationships, where our work has shown that the deformation sequence is identical across the boundary proposed by Glikson, and that regional metamorphic isograds also transgress the proposed hiatus (Fig. 5). In some areas (e.g. Eundynie) the proposed stratigraphic sequence is inverted. Thus, although greenstone sequences may include unconformities they must, at least in the Widgiemooltha-Norseman area, be limited to breaks in greenstone volcanism and/or sedimentation rather than breaks separating discrete tectono-metamorphic cycles with intervening granitoid plutonism.
A BASEMENT TO THE GREENSTONE BELT As discussed above there is limited geochronological evidence of a sialic crust significantly older than the greenstone sequences. Available Rb/Sr isotopic data appear to preclude exposed banded gneisses as samples of this most ancient recorded crust, although structural evidence indicates that they predate greenstones and hence represent candidates for at least a component of a basement to the greenstone sequence. However, the poorly understood early structural history of the greenstones, combined with the occurrence of the high strain-high metamorphic grade zones along greenstone-granitoid margins, effectively preclude a direct resolution of early structural relationships between them. Even if it is established that some granitoid rocks formed prior to greenstones, the possibility that greenstones were technically juxtaposed against an older sialic crust cannot be precluded. Evidence on the tectonic setting for greenstone belt formation can only be inferred from stratigraphy.
CRUSTAL EVOLUTION IN EASTERN GOLDFIELDS, W.A. Kb f Q-Ab-Or Limit of ^quartz-muscovite stability t_600
700
°C
PARTIAL MELTING IN GNEISSES -^-GRANITOID CRUST
700PC-
800
®C
500
HIGH GRADE HIGH STRAIN DOMAIN P-T T R A J E C T O R Y
GRANITOID ROCKS
200
400
°C
LOW STRAIN DOMAIN P-T TRAJECTORIES
- L O W STRAIN DOMAIN-
- - - - - T- - - - T- •
\a a a A A A A v A A /V, , A AAA A\ A < t ' A A A A AAA\ AMAA
aVaaaaaaa,, c . a \ \ a a a a a a a A 4 K l* A/.AAAAAAA) \ A> A A A A A A V AVAAAA ^ * AAA A>.AAA v^ A AV./V /C, A '' « - \V ,/. A A^ .
Fig, 6.
°C
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I I I J Syn-kinematic
600
MEDIUM GRADE HIGH STRAIN DOMAIN P-T TRAJECTORY
- H I G H STRAIN DOMAIN-
—-
Post-kinematic
f ^ ^ ^ ^ Apparent -(Lmetamorphic " naAtharm geotherm
Cordierite quartz 600
499
Staurolite Sti Chloritoid toid + aalmandine + Q u a r t z ^ quartz
1
—
—
—
. H- -
600°C
GREENSTONES Banded gneisses
Low strain domain
Siliceous granulites
a High strain domain
Horizontal scale
10
Schematic sections of thermal regime as preserved now. Note metamorphic temperatures are not precisely determined, particularly in the lower grade areas. P-T trajectories for low strain domain inferred from supposition that rapid burial was necessary to preserve igneous minerals in higher grade zones. Chloritoidstaurolite-almandine-cordierite stability in medium grade, high strain domain from Richardson (1968) for Fe end members: the presence of Mg will extend cordierite stability to higher pressures. High grade, high strain domain conditions constrained by reaction staurolite + muscovite + quartz = andalusite + biotite 4- fluid (relict staurolite enclosed in andalusite), reaction andalusite = sillimanite, and anthophyllite stability. Al-silicate stability after Richardson et al. (1969). Conditions of melting in granitoid crust constrained by minimum water-saturated melting of quartz-orthoclase-albite, quartz 4- muscovite stability, and comparison of mesonormative compositions with experimentally-derived partial melt compositions (Winkler et al., 1975; Winkler, 1976).
Greenstones formed in a simatic setting would presumably be underlain by mantle peridotite with a tectonite fabric analogous to modern ophiolite sequences but this has not been identified in the Yilgarn Block or indeed in any greenstone terrain. Further, the stratigraphic sequence within greenstones should reflect this special setting. In modern ophiolite sequences, a simatic setting is associated with a distinctive, entirely igneous stratigraphy (except for younger overlying sediments) and with structures associated with spreading. It may be possible to envisage simatic tectonic settings rather different from modern oceanic spreading ridges in the Archaean, but none has been postulated to explain the characteristic interbedding of sedimentary and volcanic rocks throughout the stratigraphic sequences, or the lack of dykes resulting from spreading episodes. Instead, many greenstone stratigraphies resemble some modern volcanic sequences erupted above continental crust.
SUBSEQUENT TECTONIC EVOLUTIONMET AMORPHISM The regional metamorphism of greenstones and inferred widespread partial melting of the gneissic crust dominate the preserved characteristics of this granitoid-greenstone terrain. By analogy with younger orogenic belts, metamorphic processes reflect the nature of energy transport in the crust, and the style of metamorphism in such terrains thus provides a sensitive record of the orogenic processes involved. However, the Archaean metamorphic patterns of the Eastern Goldfields Province are distinctive and cannot be directly compared with any postArchaean tectonic province (Binns et al., 1976). As such, they place significant constraints on tectonic models for the area. Significant facets include: (i) Peak metamorphic conditions were established late in the deformation sequence; post-D 2 and syn- and post-D 3 .
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N. J. ARCHIBALD, L. F. BETTENAY, M. J. BICKLE & D. I. GROVES
(ii) A crustal thickness of ca 50 km was established before or during the tectonic events responsible for the deformation and metamorphism (we reject mechanisms for later crustal thickening by processes at depth that produce no near-surface tectonic expression). (iii) The earliest deformation in some greenstone belts involved overturned to recumbent, macroscopic, pre-metamorphic folding providing a possible mechanism for crustal thickening at least within greenstone belts. (iv) The distribution of relict igneous mineral phases in mafic-ultramafic rocks (Binns et al., 1976) implies that water became a widespread mobile phase only late in the metamorphic cycle, perhaps due to rapid early burial of precursors to high-grade areas. (v) Areas of low metamorphic grade within greenstone belts correspond to areas of high gravity anomaly (Binns et al., 1976). Highergrade greenstones are restricted to linear belts along present margins of greenstone areas that are characterised by high strain and numerous synkinematic granitoid diapirs: in places, hornfelsing associated with post-kinematic granitoids overprints these metamorphic fabrics. The high- grade zones are inferred to continue at depth below the low-grade greenstones but whether they are also high-strain zones analogous to the exposed margins is not known. (vi) Metamorphic assemblages containing equilibrium textures that can be employed to constrain P-T trajectories are rare (see Fig. 6). Assemblages containing staurolite-garnet and chloritoid rimmed by cordierite (Archibald, 1979) are inferred to imply isothermal uplift of medium grade areas. In high-grade areas, a significant temperature increase and a slight pressure drop are inferred on the basis of staurolite rimmed by andalusite, itself rimmed by cordierite and sillimanitemuscovite-biotite (Archibald, 1979). Derived thermal gradients range from ~ 60°C/km in high-grade regions, to less than 45°C/km in medium-grade regions (chloritoid-staurolitegarnet-cordierite assemblages), and possibly as low as 30°C/km in lower-grade greenschist-facies areas (on the basis of stable kyanite in greenschist assemblages inferred to be underlain by ca 10 km of greenstone within hydrous metamorphic fades). Models for greenstone metamorphism must therefore take into account the ca 20 km structural thickness of greenstone sequences, at least in low-grade areas, the variation in implied geothermal gradient from possibly as low as ca
30°C/km in low-grade areas to ca 60°C/km in high-grade areas, a crustal thickness of some 50 km, and the uniformly high temperatures experienced by the granitoid crust at exposed levels. The main uncertainty in thermal models arises from lack of data on the relationship of the high-grade, highly strained marginal belts to the lower-grade, apparently less-strained areas. In particular, were the marginal zones a focus of some special tectonic activity which resulted in especially high geothermal gradients? There is some evidence that the present greenstone margins were important tectonic interfaces early in the history of the terrains (Archibald & Bettenay, 1977). However, models which require the high gradients on greenstone margins to result from in situ heating, are difficult to reconcile with the overall crustal structure. Simple conduction of heat from below would be incapable of maintaining the thermal gradients along the highgrade greenstone margins without complete melting of the basal 30 km or more of crust. Intrusion of igneous melts could give rise to such gradients, but basic plutonic bodies are too few to contribute significant heat, and intrusion of granitic material at most one or two hundred degrees hotter than the required temperatures would have been a very inefficient process. For example, simple calculations (cf. Oxburgh and Turcotte, 1970) suggest that granitoid magmas would have to constitute at least 90% of the crust before the required temperatures were reached. Models in which the high-grade marginal belts and granitoid crust were derived by differential uplift of a uniformly buried terrain provide more realistic crustal thermal structures. Cooling of thick rock masses is a slow process, and uplift at geologically reasonable rates would lead to enhanced geothermal gradients as hot rock was brought near^surface (England & Richardson, 1977). The inferences for rapid early tectonic burial and of falling pressures at peak metamorphic conditions are both consistent with thermal models requiring rapid early crustal thickening with the subsequent thermal relaxation being accompanied by uplift and erosion. Extrapolation of thermal gradients in the lowgrade areas (say 3kb and 300°C-400°C with 6-10 km of greenstones below the present surface) indicates that high-grade greenstones (600 °C800 °C and ca 6kb) should be present at depth. Rapid local uplift of a terrain with such thermal gradients could transpose the granitoid-greenstone marginal zones into the 3-5 kb (10-15 km depth) region with relatively little temperature drop. Direct evidence that this mechanism was responsible for the relatively high-temperature low-pressure assemblages in the high-grade areas,
501 stones may record a tectonic event of regional significance. This early structural complexity and poor exposure make us cautious about attempts to establish stratigraphic correlations on a province-wide scale. In particular there is no evidence of "Upper" and "Lower" greenstones or of major unconformities within the greenstone sequence in the areas that we have studied. A major tectonic episode thickened the greenstone sequences to at least 20 km locally and following this both the greenstone and granitoid components of the crust underwent a sequence of tectonothermal events which included further deformation, metamorphism and partial melting of the granitoids. Early rapid burial of precursors to high-grade greenstones is indicated by the preservation of certain relict igneous minerals only in these suites (c/. Binns et al., 1976) and by metamorphic textures reflecting P-T trajectories towards lower pressures. Regional variation of metamorphic grade may reflect variable and relatively rapid uplift of the granitoid basement and associated higher-grade high-strain marginal zones of greenstone belts. Metamorphic conditions exceeded minimum melting temperatures across an extensive thickness of the granitoid crust and the prominent intrusive granitoids formed during and after the deformation events are interpreted as water-saturated partial melts of pre-existing sial that included the exposed banded gneisses. Extensive melting of basic components of preserved greenstone sequences seems highly unlikely as seismic and gravity data restrict mafic granulites to a minor proportion of the crust and the range of granitoid compositions does not match that expected of a mafic source. In particular there is no evidence of a secular variation of granitoid composition from "Na-rich" to "Krich" previously claimed to support models involving progressive partial melting of greenstones and derived products; data upon which this speculation rests are largely irrelevant to the perspective of regional studies. The later deformations and metamorphism of the Archaean crust are inferred to have taken place with moderate mantle thermal gradients and a relatively thin lithosphere although it is not possible to exclude some transfer of heat by intrusion of mantle-derived igneous melts. Significant convection of heat by partial melts may have occurred within the lower parts of the sialic crust although products of such melting have not been recognized. None of the tectonic features is clearly inconsistent with plate-tectonic mechanisms although a distinctive asymmetry of volcanic, plutonic or metamorphic events, as in modern converging plate margins, is not apparent. It is significant that the high-grade
CRUSTAL EVOLUTION IN EASTERN GOLDFIELDS, W.A.
such as preservation of relict mineralogies from high-pressure assemblages (e.g. kyanite), is however absent. The mechanism for metamorphism of the upper crustal layers is thus not yet resolved. Both early crustal thickening and igneous intrusion may have contributed significantly to the thermal picture and at present neither process alone can account for the observed metamorphic conditions. Perhaps the most significant aspect of the metamorphism is the inferred thermal gradient for the lower-grade areas (ca 30 to 40°C/km). Such gradients would require a steady-state heat flow of 1.8 to 2.4h.f.u. (10~ cal. c m - sec- ). Given present estimates of radioactive heat production in the Yilgarn granitoid crust (Sass et al., 1976), and allowing for a two-fold increase at 2700Ma, requires that mantle heat flow was 1.2 to 1.8h.f.u. (20-30°C/km) (assuming steadystate conditions). Such a gradient could only be maintained by conduction across the basal 30 km of the granitoid crust if the bulk of this crust was anhydrous, and even so temperatures of 1250°C would be predicted for the base of such a crust (note that O'Hara & Yarwood, 1978, predicted such metamorphic conditions—1150°±100°C at 15±3kb, for Scourian. Archaean metamorphism). Perhaps a tectonic setting with a thin (10-30 km) lithosphere below the granitoid crust and moderate mantle heat flow provides the most reasonable tectonic setting for this segment of Archaean crust. 6
2
1
ARCHAEAN CRUSTAL DEVELOPMENT IN THE YILGARN: CONCLUSIONS The high-grade gneisses in the west and southwest comprise the oldest known rocks in the Yilgarn Craton. The seismic structure implies that these high-grade siliceous granulites may continue as a layer at depth under the predominantly amphibolite-facies granitoid-greenstone terrain to the east. The main constituent of the granitoid crust enveloping greenstone belts is inferred to be banded gneiss derived from a low Rb/Sr source {i.e. mantle or basic crust) and deformed by a regional high-strain tectonic event or events immediately prior to greenstone deposition. Gravity and seismic data imply that greenstone belts in the Eastern Goldfields Province are preserved as a relatively thin (less than 10 km) discontinuous layer on a thick ca 40 km sialic crust. The sedimentary-volcanic stratigraphy of the greenstone sequences is consistent with formation above a sialic crust, possibly as an active trough (or troughs) flanked by more stable basins. Early overturned and recumbent folds within the green-
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metamorphic areas are best explained as areas in which thermal gradients are enhanced, either by uplift, or by a combination of uplift and intrusion. Gradients in the lower-grade areas imply mantle thermal gradients not much more than those in Archaean high-grade gneiss terrains and taken by Bickle (1978) to indicate that much of the Earth's heat must have been lost by platetectonic or analogous processes at 2700 Ma. However, neither the stratigraphy nor tectonic setting of the greenstones in the Eastern Goldfields Province or others in the Yilgarn suggest oceanic crust. The earlier tectonic history of the Eastern Goldfields Province is perhaps the more intriguing, if least understood. There is tenuous isotopic evidence of a significantly older sialic crust in the Province but the widespread banded gneisses were apparently derived from a mantle or low Rb/Sr source only one or two hundred million years before greenstone metamorphism. Such a vast volume of sialic crust would require melting of even greater volumes of basic crust, but there are neither signs of preserved source rocks nor the residue of partial melting in the exposed terrains or in crustal profiles. An early subduction-like process in which a renewable basaltic slab provides a source of the initial granitoids and a sink for residues of melting is one potential solution. The regional high-strain event that yielded the banded gneisses is equally perplexing. Even if this does relate to major sub-horizontal deformations as suggested for other gneiss terrains, what caused the deformation and what crustal masses were involved? Finally the early folding and thickening of the greenstone sequences is still poorly understood. This early folding might
represent a high-level equivalent of the deformation of the gneiss terrain although there is no direct evidence to support this. An understanding of the critical early period in the development of this craton is essential, and more precise geochronology and isotopic data are clearly required. Given the severity of events within the interval 2700-2550 Ma, together with the poor outcrop on any standards, unravelling the early history of the Eastern Goldfields Province will not prove to be an easy task, so that debate concerning the Archaean evolution of this crustal segment will continue. Despite the important constraints provided by the later tectonometamorphic history of the terrain, attempts to infer processes on a global scale seem premature until outstanding problems are at least partially resolved. ACKNOWLEDGMENTS The early studies upon which this paper was based were partly funded by the Australian Mineral Industries Research Association, the Australian Research Grants Committee, and the University of Western Australia. Postgraduate Studentships awarded by CSIRO (N.J.A.) and the Commonwealth of Australia (L.F.B.) are gratefully acknowledged. Various exploration and mining companies have provided access to company maps and other information, and we are indebted to the many students whose unpublished B.Sc. Honours projects at the Department of Geology, University of Western Australia have supplied corroborative information to the studies specifically acknowledged here. We thank Dr R. D. Gee for his critical review of the manuscript.
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AUTHOR INDEX Citations in Italics refer to pages on which the references are set out. Authors may be cited more than once on the pages listed. A few authors have varied the initials under which they have published, and are accordingly listed under each combination of initials used. Abercrombie, I. E., 178, 186 Adie, A. J., 203 Ahmad, S. ML, 253 Ahmad, S. N., 253, 456 Ahrens, L. H., 325 Akhmetov, R. N., 119 Albarede, F., 299 Alexander, E. C., 353, 360 Algarte, J. P., 141 Allaart, J. H., 191, 220, 221, 302, 303, 310, 311, 313, 325 Allchurch, P. D., 17, 186, 504 Allegre, C. J., 192, 296, 299, 325 Allison, I., 107, 108 Allsopp, H. L., 28, 31, 156, 205, 220, 462, 466 Almeida, F. F. M., 133, 137, 140, 141, 142 Alves, F. R., 137, 139, 141 Amit, O., 337, 348 Anders, E., 453, 456 Andersen, L. S., 156, 173, 175, 186, 304, 310, 503 Anderson, C. A., 203 Anderson, D. L., 5, 15 Anderson, J. L., 93, 98, 337, 349 Anderson, P.,17 Anhaeusser, C. R., 24, 26, 27, 28, 30, 31, 50, 55, 58, 62, 67, 88, 96, 235 , 243, 272, 272, 310, 310, 335, 335, 337, 338, 339, 343, 344, 348, 379, 382, 457, 459, 460, 461, 462, 464, 465, 466 Appel, P. W. U., 10, 15, 316, 325 Archibald, N. J., 11, 75,.48, 55, 173, 174, 178, 181, 184, 186, 282, 285, 368, 369, 372, 379, 382, 399, 404, 411, 419, 429, 432, 434, 438, 491, 492, 493, 494, 495, 497, 498, 500, 502, 503 Arkhipov, Yu. V., 117, 119 Armstrong, R. L., 8, 13, 15, 54, 55, 218, 219 Arndt, N. T., 192, 269, 272, 290, 296, 297, 299, 303, 305, 306, 310, 310, 321, 325, 347, 348 Arriens, P. A., 12, 15, 33, 40, 44, 55, 59, 63, 67, 145, 148, 149, 150, 154, 155, 155, 156, 160, 168, 170, 171, 189, 191, 195, 197, 198 , 201, 203, 394, 395 , 398 , 404, 429, 438, 492, 502 Arth, J. G., 92, 96, 97, 217, 219, 281, 285, 448, 448, 471, 475, 478 Arthur, A. C., 141 Ashley, G. M., 236, 243 Ashworth, J. R., 337, 348 Auvray, B., 438 Awramik, S. M., 9, 15 Ayres, D. E., 426, 428
Badham, J. P. N., 106, 107, 108 Baer, A. J., 11, 13, 14, 15, 484, 485, 489 Bahnemann, K. P., 329, 330, 335 Bai Jin, 132 Baikova, V. S., 119 Baksi, A. K., 393, 405 Balashova, Z. N., 119 Ballard, R. D., 266, 272 Bamford, D., 38, 40 Banks, P. O., 98 Baragar, W. R. A., 15, 99, 107, 108, 109, 110, 286, 295, 299, 305, 306, 311, 443, 449 Barbosa, J. F., 140 Barbosa, O., 135, 140 Barghoorn, E. S., 18 Barker, F., 17, 31, 97, 217, 219, 220, 221, 272, 272, 281, 285, 360, 385, 392, 404, 466, 474, 478 Barley, M. E., 59, 61, 63, 67, 225, 226, 227, 228, 231, 232, 232, 235, 243, 263, 265, 266, 267, 268, 269, 271, 272, 361, 572,409, 416, 417, 419 Barlow, R. B., 449, 450 Barnes, H. L., 428 Barnes, R. G., 50, 55 Barrett, F. M., 419 Barton, J. M., 21, 22, 23, 24, 25, 26, 27, 30, 31, 86, 96, 330, 331, 332, 333, 334, 335, 337, 338, 343, 348, 459, 461, 462, 463, 464, 465, 466 Basumallick, S., 137, 140 Bau, A. F. S., 101, 108 Bauer, R. L., 95, 96 Baxter, J. L., 43, 47, 49, 55, 156, 409, 415, 419, 503 Bayley, R. W., 249, 253 Beaty, D. W., 455, 456 Beckinsale, R. P., 245, 253 Bell, C. K., 138, 140 Bell, I., 101, 110 Bell, P. M.,392, 504 Benow, C. D., 141 Berry, H., 203 Beryozkin, I., 114, 119 Best, M. G., 284, 285 Beswick, A. K., 289, 290, 299 Betekhtin, A. G., 119 Bettenay, L. F., 15, 48, 55, 61, 66, 67, 156, 173, 175, 177, 178, 181, 184, 186, 263, 272, 285, 361, 372,382, 399, 404, 419, 438, 491, 493, 494, 495, 496, 498, 500, 502, 503 Bettencourte, J. S., 140 Bickford, M. E., 477, 478 Bickle, M. J., 61, 66, 67, 156, 173, 186, 272, 300, 311, 361, 364, 367, 368, 369, 370, 372, 484, 487, 489, 491, Baadsgaard, H., 9, 15, 101, 109, 110, 205 , 210, 219, 494, 502, 502, 503 220, 313, 325
506
INDEX
Bilibina, T. V., 115,119, 120 Billings, G. K., 249, 253 Binder, A. B., 5, 15 Binns, R. A., 75, 49, 53 , 55, 156, 173, 186, 285, 372, 352, 390, 391, 392, 479, 429, 430, 433, 438, 491, 496, 497, 499, 500, 501, 502, 503 Binns, R. E., 231, 232 Birk, D., 470, 478 Birkett, T. C., 16, 17, 69, 300, 372,:490 Bishop, F. C., 18 Black, L. P., 9, 15, 71, 81, 82, 83, 205, 213, 215, 220, 353, 360 Blais, S., 429, 438 Blake, D. H., 221 Blake, T. S., 67, 272, 361 Blanchard, D., 449 Blatt, H., 162, 171 Blight, D. F., 49, 55, 198, 199, 201, 203 Blockley, J. G., 33, 40, 57, 59, 61, 63, 67, 68, 150, 151, 155, 755, 156, 189, 191, 362, 372, 373, 382, 382, 409, 415, 416, 419 B.M.R., 33, 40 Boese, M., 504 Bohlen, S. R., 390, 392 Boothroyd, J. C., 236, 243 Boronikhin, V. A., 119 Bostock, H. H., 107, 108 Boulter, C. A., 67, 272, 361, 372, 502 Bowden, A. A., 360 Boyle, R. W., 106, 108 Brakel, A. T., 35, 40 Branston, J. C., 56, 504 Brandt, R. T., 57, 67 Brandt, S. B., 112, 113, 117, 119 Braybrooke, J. C., 77 Breaks, F. W., 441, 450 Breitbart, R., 346, 349 Bricker, O. P., 454, 456 Bridgwater, D., 9, 10, 12, 14, 75, 16, 17, 72, 83, 85, 86, 96, 98, 108, 110, 191, 205 , 206, 207, 210, 219, 220, 221, 255, 260, 302, 303, 310, 311, 313, 314, 316, 319, 320, 321, 322, 325, 347, 348, 349, 351, 356, 358, 360, 385, 386, 391, 392, 482, 483, 484, 489, 489, 490, 493, 494, 503 Briqueu, L., 219, 220 Brito Neves, B. B., 135, 137, 140, 141, 142 Brook, W. A:, 59, 61, 69, 151, 153, 757, 189, 192, 225, 233, 263, 273 Brooks, C., 156, 171, 186, 192, 215, 217, 221, 290, 299, 303, 305, 306, 310, 310, 449 Brown, D. A., 57, 67 Brown, G. C., 208, 220, 305, 310 Brown, T. H., 438 Browne, W. R., 68 Bryan, M. P. D., 109 Buick, R., 9, 15, 16, 60, 68, 225, 233, 271, 272 Bunting, J. A., 49, 50, 51, 52, 53, 55, 150, 154, 755, 211,285, 411,416, 419 Burger, A. J., 29, 37 Burk, C. A., 18 Burke, K., 11, 14, 16, 52, 53, 55, 62, 67, 484, 489 Burley, A. J., 25,37 Burnol, L., 420 Burns, M. S., 428
Busch, W., 208, 220 Bye, S. M., 16, 55, 156, 186, 285, 419, 503 Calk, L., 233 Cameron, A. G. W., 5, 16 Cameron, M., 167, 169, 171, 212, 213, 214, 216, 220 Cameron, R. L., 198,203 Caminos, E., 142 Campbell, F. E., 261 Campbell, F. H. A., 99, 107, 108 Campbell, K. S. W., 67 Cann, J. R., 63, 69, 300, 306, 308, 377 Cannon, W. F., 87, 93, 96, 97 Cant, D. J., 237, 243 Capdevila, R., 438 Card, K. D., 97 Carignan, J.\ 299 Carmichael, D. M., 109 Carmichael, I. S. E., 324, 325 Carneiro, C. D. R., 141 Carter, D. N., 10, 16, 285 Carter, J. D., 47, 55 Cawthom, R. G., 281, 285, 310, 310 Cecile, M. P., 99, 107, 108 Chadwick, B., 14, 16, 205, 220, 303 , 310, 358, 360, 482, 489 Chamberlain, V. E., 732 Chapman, H. J., 152, 156, 173, 205, 213, 220, 493, 494, 495, 498, 503 Chappell, B. W., 62, 67, 69, 138, 142, 156, 178, 186, 381, 382, 383 Charlesworth, H. A. K., 339, 349 Chayapathi, N., 253 Cheng Yu-chi, 121, 122, 732 Chikhaoui, M., 440, 448 Chin, R. J., 51, 150, 755, 372, 383, 392, 411, 416, 419 Choudhuri, A., 140, 140 Christ, C. L., 437, 438 Chung Fu-Dao, 732 Church, W. R., 295, 299 Clark, L. D., 97 Clarke, D. B., 310, 310, 311 Clarke, E. de C., 57, 67 Clayton, R. N., 393, 395, 404, 453, 455, 456 Clifford, T. N., 140 Cloud, P. E., 8, 9, 15, 16 Coe, K., 16, 358, 360 Coimbra, A. M., 141 Coleman, L. C., 109, 110, 286 Coleman, R. G., 11, 16 Collerson, K. D., 54, 55, 86, 96, 171, 205 , 206, 207, 210, 212, 217, 219, 220, 221, 255, 260, 260, 294, 296, 299, 325, 385, 391, 392, 440, 442, 446, 448, 493, 503 Collins, V. P. M., 416, 419 Collinson, J. D., 239, 243 Comaford, D. J., 193 Compston, W., 47, 54, 56, 59, 63, 67, 69, 145, 148, 150, 153, 155, 156, 157, 159, 167, 777, 181, 186, 189, 191, 195, 203, 205, 210, 212, 220, 221, 249, 250, 253, 429, 438, 492, 504 Compton, P., 210, 220 Condie, K. C., 9, 16, 29, 37, 62, 67, 255, 260, 289, 290, 294, 299, 308, 310, 382, 382, 443, 446, 448, 448,
INDEX
507
449, 457, 464, 465, 466, 469, 470, 471, 475, 478., 479 Dixon, J. M., 369, 372, 493, 496, 503 483, 484, 489, 493, 503 Doe, B. R., 92, 96
Cong Bolin, 132 Conrad, R. L., 203 Constable, S. C., 492, 503 Coomer, P. G., 426, 428
Doepel, J. J. G., 194, 195, 203 Doherty, P. C., 470, 479 Doig, R., 109 Donakov, V. I., 119 Cooper, J. A., 53, 55, 59, 61, 63, 68, 150, 153, 154, Donaldson, C. H., 273 156, 174, 179, 184, 186, 188, 189, 190, 191, 203, 282, Donaldson, J. A., 9, 16, 98 285, 498, 503 Donnelly, T. H., 233, 419 Cooper, R. A. 203 Dorr, J. V. N., 436, 438 Cordani, U. G., 135, 137, 139, 140, 140 Dostal, J., 260, 448 Costa, L. A. M., 137, 139, 140, 140 Dougan, T. W., 493, 503 Costa, U. R., 453, 456 Douglas, R. J. W., 86, 96, 98, 109 Cousineau, P., 232, 272 Doust, Q., 34, 40 Coward, M. P., 72, 83, 483, 489, 494, 503 Drake, C. L., 18 Cox, K. G., 311, 329, 335 Drummond, B. J., 33, 35, 38, 40, 40, 44, 55, 66, 68, Cox, R., 310, 493, 503 382, 383, 492, 503 Cray, E. J., 132 Drury, S. A., 105, 108, 253, 392, 485, 489 Crockett, J. H., 8, 17 Dubois, J. F., 93, 96 Crook, K. A. W., 67 Duk, V. L., 117, 119 Dunbar, G. J., 4, 9, 16 Cull, J. P., 38, 40 Dunlop, D. J., 14, 16 Currie, K. L., 47, 55, 98 Dunlop, J. S. R., 9, 15, 16, 59, 60, 67, 68, 225 , 227, 228, 230, 232, 232, 233, 243, 271, 272, 372, 419 Dallwitz, W. B., 83 Dunning, F. W., 339, 349 Daly, R. A., 332, 335 Dupuy, C., 448 Dalziel, I. W. D., 37, 55, 244, 490 Durney, D. W., 157, 186, 504 Daniels, J. L., 34, 37, 40 Dusanowskyj, T., 479 Darracott, B. W., 25, 31 du Toit, M. C., 329, 335, 335 Dashkova, A. D., 119 Dymond, J., 7, 16 da Silva, L. C., 141 Eade., K. E., 255, 257, 258, 260 Davenport, P. H., 439, 449 Eckstrand, O. R., 17 David, T. W. E., 57, 68 Edgar, A. D., 448, 449 Davidson, A., 100, 101, 106, 107, 108, 109 Egorova, T. V., 119 Davidson, L. R., 44, 55, 385, 392 Eichelberger, J. C., 218, 220 Davies, F. B., 484, 490 Eichmann, R., 15, 16, 18, 456 Davies, R. D., 26, 27, 28, 31, 156, 205, 220, 462, 463, Elder, J., 487, 489 464, 466 Elias, M., 41, 48, 55, 157, 385, 392 Davis, J. C., 377, 380, 382 Elliott, I. L., 449, 450 Davy, R., 61, 62, 63, 68, 263, 272, 373, 379, 3£2 Ellis, A. J., 454, 455, 456 Dawson, J. B., 8, 16 Ellis, D. J., 71, 83 Dearnley, R., 339, 349 Ellis, H. A., 432, 435 , 438 de Bie, I., 109 Embleton, B. J. J., 40, 40 Deer, W. A., 321, 325 Emslie, R. F., 98, 321, 322, 325 de Gasparis, A. A. A., 26, 31, 463, 464, 466 Engel, A. E. J., 131, 132, 305, 306, 311 De Jong, K. A., 372, 490 Engel, C. G., 132, 311 de Laeter, J. R., 33 , 34, 40, 41, 44, 55, 59, 63 , 68, 69, 145, 150, 151, 153, 154, 155, 156, 157, 160, 170, 171, England, P. C., 484, 489, 500, 503 173, 174, 178, 183, 186, 188, 189, 191, 362, 372, 373, England, R. N., 83 Epstein, S., 455, 456 382, 382, 392, 418, 419, 503 Eriksson, K. A., 9, 16, 24, 29, 31, 61, 68, 235 , 242, de la Hunty, L. E., 37, 40, 41 244, 263, 271, 272, 409, 411, 417, 419 Delaney, J. S., 18 Erlank, A. J., 31, 191, 347, 349, 466 Delevaux, M. H., 92, 96 Ermanovics, I., 325 Denham, D., 38, 40 Ermanovics, I. F., 335 DePaolo, D. J., 190, 191, 214, 215, 220 Ermolaev, B. A., 119 , Derrick, G. M., 289, 295, 299 Escher, A., 325, 360, 482, 486, 489 Descarreaux, J., 439, 449 Escher, J. C., 354, 360 Devadu, G. R., 253 Essene, E. J., 390, 392 Dewey, J. F., 16, 55, 67, 489 Esson, J., 310 de Wit, M. J., 27, 31, 244, 490 Evans, C. R., 339, 349 Dickinson, B. B., 391, 392, 483, 489 Evans, R. B . , 3 i Dickinson, W. R., 4, 16 Evensen, N. M., 31, 68, 187, 191, 192, 220, 221, 233, Dietz, R. S., 201, 202, 203 273, 299, 311, 325, 360, 372, 453, 456, 466 Dimroth, E., 8, 17, 226, 232, 266, 272 Eyal, Y., 337, 348 Divakara Rao, V., 253, 490
508
INDEX
Fahrig, W. F., 98, 255, 257, 258, 260 Fairbridge, R. W., 44, 55 Fanale, F. P., 7, 8, 16 Farhat, J., 77 Faure, G., 27, 31, 178, 179, 186 Felton, E. A., 416, 419 Ferguson, J., 260 Ferrara, G., 440, 449 Filho, V. O., 141 Finkl, C. W., 44, 55 Fiori, A. P., 140 Fisher, D. E., 7, 16 Fitton, M. J., 57, 58, 68 Flanagan, F. J., 305, 311 Fleet, A. J., 441, 450 Fleitout, L., 490 Fletcher, C. L, 503 Fletcher, R. C., 485, 490 Fletcher, W. K., 449, 450 Flinn, G. W., 349 Floyd, P. A., 290, 294, 299, 300 Folinsbee, R. E., 110 Folk, R. L., 227, 228, 233 Fornari, D. J., 232, 233 Foster, J. J., 171, 220 Fountain, J. C., 448, 449 Fourcade, S., 393, 403, 404 Francis, P. W., 218, 220 Franklin, J. ML, 439, 449 Fraser, A. R., 34, 40, 492, 503 Fraser, J. A., 97, 99, 106, 108, 109, 110, 449, 450 Frater, K. M., 421, 424, 427, 428 Fratta, M., 347, 349 Frey, F. A., 279, 285, 296, 299, 440, 449 Friend, C. R. L., 301, 303, 304, 307, 310, 311, 356, 360 Friend, C., 310 Fripp, R. E. P., 22, 31, 329, 332, 333, 335, 465 Frith, R., 109 Frith, R. A., 99, 100, 101, 105, 107, 109 Froidevaux, C., 490 Frumkin, I. M., 117, 119 Fryer, B. J., 54, 55, 217, 220, 255, 260, 260, 294, 296, 299, 440, 441, 442, 443, 445, 446, 448, 448, 449, 450 Fuchs, K., 38, 40 Fuck, R. A., 140 Fujimore, S., 135, 139, 141 Furnell, R. G., 10, 18, 25, 31 Fyfe, W. S., 11, 13, 16, 133, 137, 138, 139, 140, 140, 141, 208, 220, 255, 260, 453, 456, 484, 490 Fyson, W. K., 105, 107, 109, 369, 372 Gaal, G., 483, 490 Gale, G. H., 10, 16, 310, 311 Gale, N. H., 15, 17, 97, 220, 221, 311 Garrels, R. M., 437, 438 Gass, I. G., 300 Gee, R. D., 33, 34, 39, 40, 43, 47, 48, 49, 50, 55, 145, 148, 152, 153, 156, 173, 186, 194, 195, 203, 242, 244, 385, 392, 399, 400, 404, 409, 411, 419, 429, 438, 491, 493, 496, 498, 503 Gelinas, L., 295, 299, 442, 449 Gemuts, I., 52, 55, 409, 419, 491, 496, 497, 498, 503 Geological Survey of Western Australia, 145, 156, 409, 419
Gerasimov, N. S., 119 Ghisler, M., 353,360 Gibbins, W., 99, 100, 101, 109 Gibbs, A. D., 16 Gilbert, M. C., 392, 504 Giles, C. W., 275, 410, 419 Gill, J. B., 281, 285, 347, 348, 349 Gill, R. C. O., 12, 76, 313, 319, 320, 321, 322, 325 Gilligan,- L. B., 419 Gillingham, J. M., 31 Girardi, V. A. V., 137, 139, 140 Glagolev, A. A., 117, 119 Glikson, A. Y., 6, 11, 12, 16, 39, 40, 44, 47, 55, 57, 62, 63 , 66, 68, 138, 141, 153, 156, 187, 189, 191, 217, 220, 233, 257, 258, 260, 264, 269, 272, 272, 273, 282, 285, 287, 289, 290, 294, 295, 296, 297, 298, 299, 310, 311, 367, 369, 370, 371, 372, 373, 378, 379, 382, 383, 457, 461, 464, 466, 483, 484, 490, 491, 493, 494, 495, 497, 498, 503 Glover, J. E., 67, 68, 69, 232, 243, 272, 372, 419, 428 Glukhovsky, M. Z., 112, 113, 119 Godlewski, M., 479 Godvinsky, G. P., 119 Goff, S. P., 108 Goldich, S. S., 86, 91, 92, 93, 94, 95, 96, 96, 97, 98, 203 Golding, S. D., 403, 404 Golding, L. Y., 51,55 Goldschmidt, W. M., 255, 260 Goleby, A. B., 260 Goode, A. D . T . , 34, 40 Goodfellow, W. D., 439, 449 Goodwin, A. M., 5, 11, 12, 16, 62, 68, 86, 88, 96, 252, 253, 257, 260, 284, 285, 295, 299, 453, 456, 469, 478, 479 Goodwin, J. A., 109 Goodwin, R., 312, 335 Gorman, B. E., 11, 16, 17, 69, 300, 325, 370, 372, 482, 483, 490 Gorsline, D. J., 242, 244 Gorton, M., 156, 173, 503 Gorton, M. P., 186 Govil, P. K., 245 Gower, C. F., 51, 52, 55, 56, 273, 275, 277, 285 Graf, J. L., 441, 449 Grant, J. A., 93, 94, 95, 96 Grasso, F., 299 Gray, C. M., 171, 220 Green, A. A., 41 Green, D. C., 101, 109, 393, 405 Green, D. H., 5, 10, 13, 16, 18, €2, 68, 135, 141, 279, 280, 284, 285, 296, 297, 298, 299, 310, 311, 390, 391, 392 Green, J. C., 88, 96 Green, T. H., 54, 55, 281, 285, 390, 392, 465, 466 Greenhalgh, D., 59, 68 Greenwood, H. J., 391, 392 Grew, E. S., 72, 82, 83 Griffin, A. C., 429 Griffin, W., 325 Griffiths, J. R., 201, 203 Grikurov, G. E., 71, 72, 83 Gross, G. A., 103, 109 Groves, D. I., 75, 55, 67, 68, 69, 156, 173, 186, 227,
INDEX 232, 233, 243, 267, 272, 273, 285, 361, 372, 352, 404, 409, 415, 416, 417, 419, 428, 438, 491, 502, 503 Gruner, J. W., 90, 91, 96 Guest, J. E., 271,273 Gunter, W. D . , 4 5 6 Gunthorpe, R. J . , 1 5 , 55, 186, 285, 372, 382, 404, 419, 438,502 Hales, P. O., 219 Hall, J. M., 109, 110, 286 Hall, R. P., 301, 303, 305, 306, 310, 311, 356, 360 Hall, W. D. M., 34, 40 Hallam, A., 201, 203 Hallberg, J. A., 10, 16, 50, 51, 55, 150, 156, 175, 178, 184, 185, 186, 260, 261, 275, 276, 277, 278, 279, 281, 282, 284, 285, 286, 287, 295, 296, 299, 416, 419, 420, 496, 503 Hama, M., 137, 141 Hameurt, J., 438 Hamilton, P. J., 25, 26, 27, 31, 59, 68, 153, 187, 189, 191, 191, 192, 205, 213, 215, 220, 221, 225, 233, 263, 273, 287, 299, 303, 311, 313, 316, 325, 361, 372, 459, 466 Hanson, G., 325 Hanson, G. N., 88, 92, 96, 97, 215, 221, 441, 447, 448, 448, 449, 471, 475, 478, 479 Hanumantha Rao, T., 253 Hao Chungrong, 132 Hargraves, R. B., 138, 141, 255, 260, 487, 490 Hari Narain, 253 Hart, S. R., 7P2, 215, 217, 221, 289, 290, 299, 303, 305, 306, 310, 310 Hartmann, L. A., 138, 139, 141 Haskin, L. A., 255, 261, 449 Haskin, M. A., 449 Hasui, Y., 137, 140, 141, 142 Hauer, L. M., 100, 101, 109 Havens, R. G., 311 Hawkesworth, C. J., 214, 221, 335, 335 Hayslip, D. L., 470, 478 Heacock, J. G., 392 Hedge, C. E., 96, 97, 337, 349 Heier, K. S., 249, 253, 260, 260, 503, 504 Heiken, G., 267, 273 Heinrichs, T. K. 251, 253 Helgeson, H. C., 437, 438 Hellman, P. L., 440, 441, 443, 449 Helmstaedt, H., 101, 109 Henderson, J. B., 4, 9, 11, 12, 16, 99, 100, 101, 103, 104, 105, 106, 107, 108, 109 Henderson, P., 440, 441, 449 Herd, R. K., 360 Herrmann, A. G., 440, 448, 449 Herve, F. A., 142 Hess, H. H., 311, 312 Heywood, W. W., 97, 101, 109, 110, 449, 450 Hibberson, W. O., 299 Hickman, A. H., 48, 55, 57, 58, 59, 60, 61, 62, 63, 66, 68, 145, 151, 153, 156, 187, 189, 190, 191, 192, 225, 232, 233, 263, 264, 265, 267, 269, 271, 272, 272, 273, 287, 289, 290, 294, 296, 297, 298, 299, 361, 362, 365, 369, 371, 372, 373, 374, 380, 382, 383, 409, 410, 411, 419, 475, 479 Hickman, M. H., 205, 221
509
Hildreth, E. W., 443, 446, 448, 449 Hill, J. D., 101, 109 Himmelberg, G. R., 94, 95, 97 Hinthorne, J. R., 194, 203 Hocking, R. ML, 56, 392 Hoefs, J., 455, 456 Hoffman, J. H.,203 Hoffman, P., 13, 16 Hoffman, P. F., 99, 109 Hofmann, A. W., 215, 227 Hogan, L., 7, 16 Holdaway, M. J., 391, 392 Holland, H. D., 7, 8, 17 Holland, J. G., 132, 285, 482, 484, 485, 490 Holler, H., 208, 222 Holt, R. W., 253 Honey, F. R., 41 Hooper, P. R., 91, 97 Hopgood, A. ML, 205, 221, 304, 312 Hor, A. K., 13, 17 Horrocks, P. C., 335 Horwitz, R. C., 33, 34, 35, 39, 40, 41, 57, 68, 429, 438 Howie, R. A., 321, 325 Hoyle, F., 455, 456 Huang, W. L., 286, 479, 504 Hubbard, N. J., 306, 308,37/ Hudson, D. R., 412,416, 419 Hughes, D. J., 301, 303, 310, 311 Hughes, D. R., 110 Humm, M., 335 Hunter, D. R., 12, 77, 27, 30, 31, 235, 244, 299, 382, 382, 457, 459, 460, 461, 462, 463, 464, 465, 466, 469, 470, 471, 475, 478, 479, 493, 503 Hurst, R., 9, 77 Hurst, R. W., 216, 220, 221, 476, 479 Hussain, S. M., 251,253 Hutchinson, R. W., 252, 253, 448, 449 Hutt, D. K., 77 Hyde, R. S„ 101, 109 Hyndman, R. D., 492, 503 Hypolito, R., 137, 139, 141 Inda, H. A. V., 140 Ingersoll, R. V., 239, 244 Ingram, P. A. J., 57, 58, 68 Innocenti, F., 449 Irvine, T. N., 305, 306, 377 Irving, A. J., 320, 325 Isachsen, Y. M., 777 Issler, R. S., 141 I to, K., 390, 391, 392 Itson, S. P., 732 Iyer, S. S., 135, 139, 140, 140 Jackson, G. D., 9, 16 Jackson, J. A., 488, 490 Jackson, M. C., 443, 447, 449 Jackson, M. P. A., 30, 37, 459, 460, 461, 462, 466 Jacobsen, S. B., 191, 192 Jaeger, J . C . , 48, 55, 96, 98, 503, 504 Jafri, S. H., 253 Jager, E., 25, 26, 30, 37, 148, 757, 189, 192, 212, 227 Jago, J. B., 203 Jahn, B.-M., 91, 92, 97, 308, 377
510
INDEX
Jahns, R. H., 349, 392 Jakes, P., 255, 260, 306, 308, 311 James, B. A., 67, 272, 361 James, H. L., 93, 95, 97, 153, 156, 249, 253, 430, 438 James, P. R., 68, 71, 72, 81, 82, 83, 138, 141, 191 Jaques, A. L., 299 Jaramillo, J. M., 142 Javoy, ML, 393, 403, 404 Jefferson, C. W., 101, 109, 110 Jeffery, P. M., 59, 68, 148, 156, 157 Jenner, G. A., 257, 260, 260, 443, 449 Jensen, L. S., 310, 443, 445, 449 Jensen, M. L., 428 Jensen, S. B., 310 Jesseau, C. W., 220 Jocelyn, J., 162, 171 Johannsen, A., 469, 479 Johnson, K., 485, 490 Johnson, R. L., 335 Johnston, C., 16, 55, 156, 186, 285, 419, 503 Johnstone, M. H., 162, 171 Jolly, W. T., 289, 299, 368, 372, 443, 449 Jones, J. G., 226, 233, 267, 273 Jones, M. J., 499 Jopling, A. V., 243 Junge, C. E., 18, 456 Juniper, D. N., 415, 416, 419 Kable, E. J. D., 299 Kaefer, L. Q., 141 Kalsbeek, F., 210, 221, 304, 312 Kamenev, E. N., 71, 75, 76, 83 Karig, D. E., 483, 490 Karsakov, L. P., 112, 119 Kasarda, J., 449 Kastrykina, V. M., 112, 119 Katz, M. B., 139, 141 Kaul, P. F. T., 137, 141 Kay, R. W., 306, 308,5/7 Kazansky, V. I., I l l , 116, 119 Keays, R. R., 260 Kelly, P. R., 193 Kennedy, G. C., 390, 391, 392 Kerr, A., 205, 208, 221 Kerrich, R., 107, 108, 109, 441, 442, 448, 449, 453, 456 Keto, L., 325, 360 Key, R. M., 22, 23, 24, 25, 31, 335 Kidd, W. S. F., 16, 55, 67, 484, 489 Kimberley, M. M., 8, 17 King, A. C., 175, 186 King, B. C., 337, 349 King, J., 109 Kiselev, G. N., 119 Kiselev, Yu. V., 119 Kishida, A., 141 Kleeman, A. W., 345, 346, 349 Kleeman, J. D., 415, 416, 419 Klein, G. deV., 232,233 Knake, D., 449 Knauth, L. P., 60, 69, 251, 252, 253, 455, 456 Knight, C. L., 55, 203, 419, 503 Koljonev, T., 478 Komaristyy, A. A., 83 Konig, ML, 201, 203
Korzhinsky, D. S., 116, 119 Kravchenko, V. M., 119 Kriewaldt, M. J. B., 58, 68, 69 Krogh, T. E., 99, 100, 101, 109, 160, 171 Kroner, A., 11, 13, 14, 17, 31, 109, 138, 141, 261, 465 Krupicka, J., 220 Krylov, V. V., 119 Kudryavtsev, V. A., 118, 119, 119 Kuno, H., 305,311 Kusakabe, M., 424, 428 Kustra, C. R., 449 Kuzmin, M. I., 119, 120 Kvenvolden, K. A., 17 Laird, M. G., 201, 203 Lambert, I. B., 39, 40, 44, 47, 48, 55, 138, 141, 195, 197, 203, 225, 227, 228, 232, 233, 282, 285, 297, 409, 417, 419, 421, 427, 428, 461, 465, 466, 491, 492, 493, 503, 504 Lambert, M. B., 99, 101, 103, 105, 109, 271, 273 Lambert, R. St J., 130, 132, 220, 281, 283, 285, 300, 482, 484, 485, 490 Lamey, C. A., 97 Lancelot, J., 192, 325 Lancelot, J. R., 219, 220 Leduc, M., 232, 272 Leech, A. P., 110 Leeper, R. H., 438 Leggo, P. J., 59, 69 Le Maitre, R. W., 331, 335, 379, 383 Leonardos, O. H., 133, 137, 138, 139, 140, 140, 141 Lewis, J. D., 40, 55, 61, 62, 63, 68, 69, 138, 141, 150, 154, 156, 178, 186, 191, 263, 272, 297, 300, 372, 373 Libby, W. G., 48, 55, 56, 145, 150, 155, 156, 157, 160, 170, 171, 183, 186, 419 Liebermann, R. C., 284, 285 Li Fuyuan, 132 Li Jiliang, 130, 132 Lilly, P. A., 335 Lindemann, W., 346, 349 Lintern, B. C., 503 Lipple, S. L., 47, 49, 55, 57 , 58 , 59, 66, 68, 69, 190, 192, 232,233, 264, 265, 267, 269, 271, 273, 287, 299, 300, 361, 362, 365, 368, 371, 372, 373, 383 Lister, L. A., 140, 382 Lock, B. E., 485, 490 Lofgren, G., 269, 273 Longstaffe, F. J., 393, 397, 402, 403, 404, All, 479 Lonsdale, P., 453, 456 Lovering, J. F., 193, 203, 220 Low, G. H., 47, 55, 56, 159, 162, 171 Lowe, D. R., 60, 69, 232, 233, 251, 252, 253, 455, 456 Lowman, P. D., 337, 349 Lowry, D. C., 195, 203 Lugmair, G. W., 190, 192 Lumbers, S. B., 97, 462, 463, 466 Luth, W. C., 345, 346, 349, 391, 392 Maaloe, S., 286, 469, 479, 504 Macgregor, A. M., 329, 335 Macke, J. E., 16, 31 Mackenzie, D. H., 337, 339, 349 Mackenzie, F. T., 456 MacLeod, W. N., 37, 40, 41
INDEX Magaritz, M., 455, 456 Mahon, M. W., 227 Mahon, W. A. J., 454, 455, 456 Maitland, A. G., 57, 69 Manson, V., 305, 306, 311 Manton, W. I., 31, 72, 82, 83, 466 Marchetto, C. M. L., 137, 140, 140 Marcopoulos, T., 504 Margulis, L., 18 Marston, R. J., 11, 17, 51, 56, 267, 273, 409, 412, 415, 419, 496, 503 Marti, K., 190, 192 Martignole, J., 484, 490 Martin, A., 300, 311 Martin, W., 312 Martineau, M. P., 107, 109 Mascaranhas, J. F., 141 Mason, B., 205, 221, 303, 311 Mason, G. D., 297, 300 Mason, R., 55, 67, 311, 329, 335, 481, 490 Masson Smith, D., 31 Masuda, A., 441, 449 Mathur, S. P., 39, 41, 44, 56, 138, 141, 492, 504 Matson, C . R . , 47P Matsuhisa, Y., 393, 404 Ma Xingyuan, 121, 123, 132 Mayeda, T. K., 393, 404, 456 McBirney, A. R.,299 McCall, G. J. H., 3, 4, 5, 7, 8, 9, 10, 13, 16, 17, 68, 269, 273 McCarthy, T. S., 463, 466 McCormick, G. R., 18 McCulloch, M. T., 93, 97, 150, 153, 156, 498, 504 McDonald, B. C., 243 McDougall, I., 203 McElhinny, M. W., 40, 41, 261, 299 McEwen, G., 335 McGregor, V. R., 9, 10, 15, 17, 83, 205, 206, 210, 220, 221, 303 , 310, 311, 313 , 325, 356, 358 , 359 , 360, 385, 391, 392, 482, 489, 490, 503 McGlynn, J. C., 98, 99, 100, 105, 107, 108, 109 Mclntyre, G. A., 148, 156, 167, 171, 178, 186, 212, 213, 214, 227 McLean, N., 335 McLennan, S. M., 6, 7<S, 255, 256, 257, 258, 259, 260, 260, 261 McLeod, H. A., 109 McLimans, R. K., 88, 97 McQueen, K. G., 419 McWhae, J. R. H., 162, 171 McWilliams, M. O., 40, 41 Mead, J., 300 Medaris, L. G., 98 Mehnert, K. R., 208, 220, 221, 464, 466 Mekhanoshin, A. P., li5, 117, 120 Menzies, M., 443, 449 Miall, A. D., 236, 237, 243, 244 Michard-Vitrac, A., 187, 192, 313, 325 Middleton, D. D.,17 Middleton, G., 171 Middleton, R. S., 445, 446, 449 Migdisov, A. A., 129, 132 Mikkola, A., 490 Millard, H. T., 17, 31, 466
511
Miller, L. J., 57, 58, 69 Milne, V. A., 68, 232 Milne, V. G., 449, 450 Minioli, B., 137, 141 Minster, B., 299 Minster, J.-F., 299 Misch, P., 337, 349 Mitchell, R. H., 8, 17 Miyashiro, A., 126, 752/289, 300, 306, 307, 311 Molnar, P., 488, 490 Monkman, L. J., 335 Montgomery, A., 57, 69 Moorbath, S., 9, 10, 15, 17, 86, 97, 138, 139, 141, 170, 171, 179, 186, 192, 205, 213, 215, 217, 219, 220, 221, 222, 303, 304, 310, 311, 312, 313, 325, 335, 335, 360, 493, 504 Moore, H., 4, 77 Moore, J. G., 233, 266, 272 Moralev, V. M., I l l , 113, 119, 120 Morant, P., 67, 272, 361, 372, 502 Morey, G. B., 86, 87, 88, 93, 96, 97, 98 Morgan, J . , 9 7 Morgan, W. C., 206, 210, 227 Morrice, M. G., 276, 286, 295, 300 Morris, G. B., 38, 41 Morrison, E. R., 335 Morrison, K., 15, 16 Mortimer, G. E., 55, 156, 186, 285, 503 Morton, R., 220 Morton, R. D., 108, 109, 110, 171 Mose, D. G., 477, 478 Moss, F. J., 56, 504 Mota, L., 35, 41 Mudrey, M. G., 87, 91, 97 Muecke, G. K„ 256, 260, All, 479 Muehlenbachs, K., 395, 404, 455,456 Muhling, J. R., 385 Muhling, P. C., 35, 40, 150, 154, 156, 392 Muir, M. D., 68, 232 Mullen, G., 453, 455, 456 Mullins, O., 273 Munroe, R. J., 504 Muratori, A., 140 Murray, R., 777 Murthy, V. R., 7, 77, 91, 92, 97, 311, 360 Mutti, E., 239, 241, 244 Muysson, J. R., 261 Myers, D., 107, 109 Myers, J. S., 9, 10, 12, 15, 17, 83, 210, 220, 221,310, 325, 337, 349, 351, 353, 356, 358, 359, 360, 386, 392, 489, 490, 494, 503, 504 Nagasawa, S., 441, 449 Nagy, B., 14, 15, 77 Nagy, L. A., 14, 77 Naidu, P. R. J., 131, 132 Naldrett, A. J., 17, 51, 56, 272, 297, 299, 300, 310, 325, 348, 419 Nance, W. B„ 255,260 Naqvi, S. M., 156, 245, 246, 248, 251, 252, 253, 303, 311, 479, 482, 483, 490, 503 Narain, H., 490 Narayana, B. L., 252, 253 Nardi, L., 138, 139, 141
512
INDEX
Nardi, L. V. S., 138, 747 Natrajan, R., 253 Nesbitt, H. W., 448, 449, 453, 456 Nesbitt, R. W., 55, 156, 186, 227, 250, 251, 253, 257, 258, 267, 278, 279, 280, 282, 255, 256, 289, 290, 294, 295, 297, 303, 306, 310, 377, 347, 349, 415, 420, 430, 435, 440, 441, 450, 503 Neumann, H., 298, 300 Newall, G., 273 Newton, R. C., 18 Nichol, I., 439, 449 Nicholls, I. A., 299 Nicholson, R., 369, 372 Nikic, Z., 99, 100, 104, 770 Nicolaysen, L. O., 148, 756 Nielsen, P. A., 99, 105, 770 Nieuwland, D. A., 47, 54, 56, 153, 155, 756, 159, 162, 165, 167, 168, 777, 492, 504 Nisbet, E. G., 290, 300, 303, 306, 377 Nockolds, S. R., 127, 732, 332, 335 Noldart, A. J., 57, 69 Norris, A. W., 98 Norrish, K., 178, 756 Norry, M. J., 306, 372 Nutman, A. P., 205, 220, 303, 370, 325, 482, 459 O'Beirne, W. R., 496, 504 O'Driscoll, E. S., 105, 770 Oehler, D. Z., 15, 77 Offe, L. A., 53 O'Hara, M. J., 218, 227, 281, 255, 483, 490, 501, 504 Ohmoto, H., 425, 426, 427, 428 Ojakangas, R. W., 88, 91, 97 Olade, M. A. D., 13, 77 Oliveira, M. A. F., 137, 139, 140, 141 Oliver, R. L., 198, 199, 201, 203 O'Neil, J., 456 O'Nions, R. K., 77, 37, 65, 97, 187, 189, 191, 192, 214, 220, 227, 233, 255, 260, 260, 273, 299, 310, 377, 325, 335, 372, 382, 353, 466 Oosthuyzen, E. J., 26, 37, 343, 349, 459, 466 Orridge, G. R., 195, 203 Osadetz, K., 97 Oversby, V. M., 33, 41, 54, 56, 59, 61, 67, 69, 150, 151, 153, 756, 188, 189, 192, 205, 212, 220, 227, 362, 372, 395, 404, 493, 495, 498, 504 Owen, P. F., 255 Owen, T., 453, 456 Oxburgh, E. R., 500, 504 Padgham, W. A., 99, 100, 101, 103, 106, 107, 109, 110 Page, R. W., 169, 777, 179, 756, 210, 227 Pankhurst, R. J., 75, 77, 97, 205, 215, 217, 220, 227, 255, 260, 260, 304, 377, 325, 360, 382, 353 Park, R. G., 52, 481, 486, 490 Patchett, P. J., 218,227 Pattenden, G. E., 267 Patterson, J. G., 109 Pavlovsky, E. V., 113, 119 Pearce, J. A., 63, 69, 290, 298, 300, 306, 308, 377, 372 Pearce, T. H., 11, 16, 17, 63 , 69, 290, 298, 300, 372, 490 Pedreira, A. J., 140, 747 Penalva, F., 137, 742
Percival, J. A., 101, 106, 770 Perminov, A. V., 119 Perrault, G., 299 Perry, E. C., 251, 253, 455, 456 Pessoa, D. A. R., 740 Peterman, Z. E., 85, 86, 91, 92, 93, 95, 97, 98, 272, 272, 474, 475 Peters, W. S., 492, 504 Petrov, V. P., 720 Petrova, Z. J., 119 Petrykowski, A. C., 372 Pettijohn, F. J., 97, 127. 732 Phaup, A. E., 493, 504 Philpotts, J. A., 308, 372 Phinney, W. C., 94, 97 Pichamuthu, C. S., 138, 747 Pidgeon, R. T., 33, 47, 59, 61, 69, 151, 153, 757, 162, 163, 777, 188, 189, 190, 192, 205, 227, 225, 233, 263, 273, 304, 372, 362, 372 Piper, J. D. A., 138, 747 Pitcher, W. S., 349, 462, 464, 466 Piatt, J. P., 11, 77, 55, 756, 175, 756, 255, 359, 497, 503, 504 Piatt, R. G., 351,360 Poetz, J. A., 449 Poldervaart, A., 377, 372 Pollock, G. D., 442, 449 Polovinkina, Yu. Ir., 112, 114, 720 Ponnamperuma, C., 75, 76, 77, 75 Porter, D., 175, 756 Potts, M. J., 449 Poulsen, K. H., 449 Powell, D., 78, 53 Powell, J. L., 27, 37, 227 Price, R. A., 709 Pride, C., 260, 479 Prider, R. T., 159, 162, 163, 777, 385, 392 Prince, L. A., 92, 97 Prinz, M., 306, 372 Purvis, A. C., 411, 420 Pyke, D. R., 10, 77, 272, 299, 370, 325, 345, 439, 443, 445, 446, 450 Rafter, T. A., 425 Raitt, R. W., 47 Ramakrishnan, M., 246, 253 Rama Rao, P., 253 Ramberg, H., 95, 97, 369, 372, 465, 466, 482, 485, 486, 487, 490 Ramiengar, A. S., 248, 253 Ramsay, J. G., 25, 37, 72, 53, 334, 335 Rankama, K., 747 Rao, V. V., 479 Rast, N., 271, 273, 339, 349 Ravich, M. G., 71, 72, 53 Reading, H. G., 243 Reilly, G. A., 267 Reimer, T. O., 9, 15, 76, 77, 37, 251, 253 Reineck, H. E., 231, 233 Rhodes, J. M., 203 Ricci, C. A., 449 Ricci Lucchi, F., 239, 241, 244 Riccio, L., 295, 299
INDEX Richards, J. R., 59, 69, 151, 153, 157, 189, 192, 195, 197, 203, 263, 273 Richardson, S. W., 391, 392, 499, 500, 503, 504 Ridge, P., 137, 141 Ridler, R. H., 62, 68, 453, 456 Rigot, W., 273 Riley, G. H., 157 Ringwood, A. E., 5, 13, 18, 28, 31, 54, 55, 135, 141, 268, 273, 296, 299, 390, 391, 392, 465, 466, 503 Rivereau, J. C., 140 Robb, L. J., 26, 30, 31, 335, 335, 337, 338, 339, 343, 344, 348, 457, 459, 460, 461, 462, 463, 464, 465, 466 Roberts, G., 41 Roberts, I. W. N., 310 Roberts, R. G., 453, 456 Robertson, A. H. F., 441, 450 Robertson, I. D. M., 138, 141 Robertson, J. A., 450 Robinson, B. W., 424, 428 Roddick, J. C., 150, 153, 154, 157, 174, 181, 184, 185, 186, 498, 504 Roddick, J. C. M., 157, 395, 404 Roeder, P. L., 321, 322, 325 Roering, C., 465 Rogers, J. J. W., 245 Rogers, N. W., 256, 260 Ronayne, E. A., 109 Ronnenson, B. M., 116, 117, 120 Ronov, A. B., 129, 132 Rose, W. I., 273 Rosenberg, R. J.,478 Rosman, K. J. R., 69, 150, 157, 186 Roy, S. D., 285, 299 Ruberti, E., 137, 139, 141 Rubey, W. W., 7, 18 Rudnik, V. A., 83, 112, 120 Runcorn, S. K., 490 Rutland, R. W. R., 11, 17, 18, 47, 56, 68, 186, 191, 369, 372, 482, 490, 504 Rvan, B., 335, 465 Ryan, G. R., 57, 58, 69 Rye, R. O., 425, 426, 427, 428 Sadowski, G. R., 137, 141 Sagen, C., 453, 455, 456 Sakrison, H. C., 439, 450 Salye, M. E., 119 Sanchez, J., 271, 273 Sanford, B. V., 98 Sangster, D. F., 59, 61, 69, 151, 153, 157, 189, 192, 225, 233, 263, 273, 416, 420, 421, 426, 428, 439, 450 Sankaran, A. V., 10, 18 Sanschagrin, Y., 232, 272 Santini, P., 137, 140 Sarkar, P., 260, 479 Sass, J. H., 492, 501,504 Sato, T., 421, 428, 448, 450 Satyanarayana, K., 253 Schau, M., 10,18, 108, 110 Schidlowski, M., 7, 8, 15, 16, 18, 455, 456 Schneider, G., 220 Schnetzler, C. C., 312 Scholten, R., 372, 490 Schopf, J. W., 4, 14, 17, 18
513
Schubert, G., 490 Schulz, K. J., 88, 96, 97 Schwarcz, H. P., 393, 397, 402, 403, 404, 404, 426, 428, 477, 479 Schwerdtner, W. M., 91, 97, 369, 372, 462, 463 , 466, 496, 504 Scott, S. D., 421, 428 Seribbins, B. J., 109 Seccombe, P. K., 421, 426, 428 Selvan, T. A., 13, 18 Serri, G., 449 Seyfried, W., 449 Shackleton, R. M., 14, 18 Sharaskin, A. Ya., 221 Sharpe, J. I., 453, 456 Sharpe, M. R., 16 Shaw, D. M., 255, 257, 260, 261, 347, 349 Shegelski, R. J. 109, 110 Sheraton, J. W., 62, 68, 71, 72, 81, 83, 385, 392 Shido, F., 126, 132, 306, 307, 311 Shieh, Y. N., 393, 402, 404, 404 Shih, C.-Y., 311 Shimizu, M., 7, 8, 18 Shor, G. G., 41 Shvetsov, E. S., 115, 117,120 Sidorenko, A. V., 116, 120 Siever, R., 8, 18 Sighinolfi, G. P., 135, 139, 141 Silver, L. T., 162, 171 Simmons, B. D., 439, 450 Simmons, G. C., 93, 96 Simonds, C. H., 7,18 Simonian, K. O., 300 Sims, P. K., 85, 86, 87, 88, 90, 91, 93, 95, 96, 97, 98 Sinclair, I. G. L., 449 Singh, I. B., 231, 233 Sklarew, D. S., 17 Sleep, N. H., 455, 456 Smewing, J. D., 295, 300 Smith, A. G., 201, 203 Smith, H. S., 31, 191, 347, 349, 466 Smith, I. E. M., 41, 96, 98, 157, 171, 192, 349, 372, 479 Smith, J. V., 4, 5, 17, 18, 218, 222, 310, 312, 360, 483, 484, 490 Smith, J. W.,428 Smith, R. E., 33, 34, 35, 37, 39, 40, 41, 289, 298, 299, 300, 449 Smith, S. E., 289, 296, 298, 300 Snowden, P. A., 369,372 Sobotovich, E. V., 71, 72, 83, 112, 120 Soderholm, B., 490 Sofoulis, J., 429, 438 Sohnge, P. G., 329, 335, 335 Sokolov, Yu. M., 120 Solomon, M., 426, 427, 428 Sopuck, V. J., 440, 443, 450 Soucie, G., 289, 290, 299 South wick, D. L., 91, 98 Spooner, C. M., 138, 141 Sproll, W. P., 201, 202, 203 Steele, I. M.,18 Steiger, R. H., 25, 26, 30, 31, 96, 98, 148, 157, 189, 192, 212, 221
514
INDEX
Stemprock, M., 420 Stephansson, O., 369, 372, 485, 490 Stephenson, N. C. N., 159, 162, 777, 195, 203 Sterenberg, V. Z., 109 Stern, C. R., 27, 37, 479, 504 Stern, T. W., 95 Stevens, B. J. P., 419 Stickney, D. M., 132 Stiliman, C. J,, 335 Stockwell, C. H., 85, 86, 98 Stolz, G. W., 278, 279, 285 Stone, D., 97 St-Onge, M., 109 Stormer, J. C., 135, 137, 139, 142 Stott, G. M.,97 Stowe, C. W., 14, 18 Streckeisen, A., 331, 335, 375, 383, 450, 469, 479 Strong, D. F., 310, 310 Stuckless, J. S., 93, 98 Sugavanam, E. B., 479 Sullivan, C. M., 435, 436, 438 Sun Dazhong, 121, 132 Sun, S.-S., 215, 227, 250, 251, 255, 257, 258, 261, 278, 279, 280, 282, 285, 255, 289, 290, 294, 295, 300, 303, 306, 310, 377, 347, 349, 420, 440, 441, 450 Suslova, S. N., 10, 18 Sutcliffe, R. H., 372, 455, 504 Sutton, J., 138, 142 Su Yun-Jun, 732 Swami Nath, J., 246, 253 Swulius, T. M., 253, 456 Sylvester, G., 68 Szewezyk, Z., 479 Talbot, C. J., 310, 372, 484, 490 Talwani, M., 201,203 Tan Yingjia, 732 Tan, F. C., 251,253 Tapponnier, P., 488, 490 Tarling, D. H., 220, 325, 455, 490 Tarney, J., 29, 37, 53, 55, 62, 59, 53, 797, 220, 235, 244, 391, 392, 417, 420, 474, 479, 483, 484, 485, 490 Taubeneck, W. H., 279 Taylor, H. P., 251, 253, 453, 455, 455, 477, 479 Taylor, P., 325 Taylor, P. N., 205, 227, 222, 304, 372 Taylor, R. P., 441, 450 Taylor, S. L., 109 Taylor, S. R., 6, 75, 255, 256, 257, 258, 259, 260, 250, 257, 281, 255 Terekhova, R. V., 119 Theron, A., 52, 55, 409, 479, 491, 496, 497, 498, 503 Theron, A. C., 493, 504 Thorn, R., 55, 273 Thompson, P. H., 99, 106, 709, 770 Thorpe, R. S., 220 Throstad, L. E., 709 Thurston, P. C., 439, 441, 442, 443, 445, 448, 450 Tilley, C. E., 281, 255, 390, 391, 392 Tingey, R. J., 53 Tirrul, R., 101, 770 Tischendorf, G., 415, 416, 420 Titov, V. K., 779 Torquato, J. R., 747
Touret, J., 404, 405 Travis, G. A., 51,55, 193 Treagus, J. E., 78, 53 Trein, E., 740 Trendall, A. F., 39, 47, 55, 59, 55, 59, 145, 151, 755, 755, 757, 755, 189, 797, 479 Treuil, M., 299 Troeng, B., 372, 455, 504 Trundel, P., 299 Trzcienski, W. E., 449 Turcotte, D. L., 500, 504 Turek, A., 150, 755, 777, 755, 195, 203, 227, 429, 435, 498, 504 Turner, A. R., 51, 55 Turner, F. J., 325 Tuttle, O. F., 349, 392 Tveten, E., 325 Tyrwhitt, D. S., 195, 203, 493, 503 Uday Raj, B., 253 Ulbrich, H. H. G. J., 137, 139, 740 Urabe, T., 448, 450 Urquart, W. E., 479 Usselman, T. M., 273 Vail, J. R., 335 Vandor, H., 709 Vandoros, P., 740 van Niekerk, C. B., 29, 37 van Reenen, D. D., 329, 335, 335 Van Schmus, W. R., 86, 93, 95, 95, 337, 349 Vakey, M. J., 705 Veizer, J., 75, 249, 250, 253, 455, 455 Verhoogen, J., 14, 75, 325 Vernon, R, H., 390, 392 Viljoen, M. J., 24, 25, 37, 55, 57, 220, 271, 273, 290, 299, 300, 303, 305, 310, 372, 337, 343, 347, 349, 457, 459, 460, 463, 464, 455, 457 Viljoen, R. P., 24, 25, 37, 55, 57, 271, 273, 290, 299, 300, 303, 305, 310, 372, 337, 343, 347, 349, 457, 459, 460, 463, 464, 455, 457 Visher, G. S., 231, 233 Visser, D. J. L., 24, 37, 457, 459, 457 Viswanatha, M. N., 246, 253, 377 Viswanathan, S., 10, 75, 88, 95, 377 Vitaliano, C. J., 300 Volkova, N. V., 779 von Platen, H., 208, 222, 337, 346, 349 Wakefield, J., 77 Walker, G. P. L., 277, 255 Walker, I., 175, 755 Walker, I. W., 55 Walker, J. C. G., 7, 8, 75 Walker, K. R., 296, 300 Walshe, J. L., 426, 427, 425 Walter, M. R., 15, 75, 51, 55, 59, 59, 225, 227, 232, 233 Walton, B. J., 370 Wang, H. C., 732 Wang Hongzhen, 121, 732 Wang Kuande, 732 Warburton, A. F., 449 Ward, W. C., 228, 233
INDEX Wasserburg, G. J., 4, 18, 93, 97, 150, 153, 156, 191, 191, 192, 214, 215, 220, 498, 504 Watson, J., 13, 18, 325, 391, 392, 409, 420, 483, Watson, J. V., 138, 142 Watt, W. S., 325, 360 Watterson, J., 482, 486, 488, 489, 490 Weaver, B., 392 Weaver, B. L., 417, 420, 477, 479 Webb, A. W., 198, 203 Wedepohl, K. H., 253, 437, 449 Wells, P. R. A., 16, 219, 222, 359, 3(50, 391, 392, 483, 490 Wenner, D. B., 453, 456 Wernick, E., 133, 135, 137, 139, 140, 142 West, G. F., 469, 479 West, K. N., 10, 16, 285 Wetherill, G., 17 Wetherill, G. W., 196, 203, 227 Wheatley, T. J.,83 White, A. J. R., 62, 67, 69, 138, 142, 306, 308, 311, 337, 349, 381, 382, 383 White, O. L.,449, 450 Wickramasinghe, 455, 456 Wierzbicki, V., 449 Wilde, S. A., 43, 47, 56, 156, 157, 159, 162, 171, 419, 503 Wildeman, T. R., 255, 261 Williams, D. A. C., 10, 18, 25, 31, 278, 279, 285, 306, 310, 312 Williams, H. R., 138, 142 Williams, I. R., 43, 44, 45, 47, 50, 52, 56, 156, 272, 273, 392, 411, 419, 420, 429, 438, 491, 496, 497, 498, 503, 504 Williams, I. S., 160, 171 Williams, J. G., 41, 96, 98, 157, 171, 192, 349, 372, 479 Williams, R. J., 273 Williams, S. J., 34, 41, 49, 50, 52, 53, 55, 56, 150, 154, 157, 277, 285, 385, 386, 389, 392 Wilson, A. € . , 31, 466 Wilson, A. F., 145, 157, 385, 392, 393, 403, 404, 405 Wilson, H. D. B., 276, 286, 295, 300 Wilson, J. F., 335 Winchester, J. A., 290, 299, 300 Windley, B. F., 13, 16, 17, 18, 31, 55, 62, 67, 69, 85, 98, 108, 110, 119, 120, 132, 141, 156, 186, 218, 220, 221, 222, 244, 253, 300, 310, 310, 311, 312, 325, 353, 358, 360, 385, 392, 417, 420, 438, 479, 483, 484, 489, 490, 502, 503
515
Winkler, H. G. F., 176, 186, 346, 349, 496, 499, 504 Wise, D. U., 8, 18 Wolery, T. J., 455, 456 Wolf, K. H., 419, 428 Wolfe, W. J., 439, 450 Wollast, R., 454, 456 Wood, B. J., 391, 392 Wood, D. A., 215, 222 Wood, D. N., 335 Wooden, J. L., 86, 93, 94, 96, 96, 98 Woodruff, L. G., 271,273 Woodwell, G. M., 455, 456 Woolsey, L. L., 93, 98 Worden, J. M., 150, 157 Wright, A. E., 83 Wright, J., 479 Wright, J. V., 277, 286 Wright, L., 83 Wright, L. I., 503 Wright, T. L., 470, 479 Wu Changhua, 121 Wu Zhengwen, 132 Wyatt, J. D., 57, 69 Wyllie, P. J., 217, 222, 282, 286, 370, 372, 465 , 466, 469, 471, 473, 479, 496, 504 Wynne-Edwards, H. R., 488, 490 Xie Guanghong, 130,132 Yan Hongquan, 132 Yardley, B. W. D., 337, 349 Yardley, D. H.,97 Yarwood, G., 501, 504 Yoder, H. S., 281, 286, 390, 391, 392 York, D., 178, 179, 186, 189, 192, 212, 213, 214, 222 Young, G. M., 255,250, 261 Yuen, P. A., 486, 490 Zartman, R. E., 97, 335 Zhang Wenhua, 132 Zhang Yixia, 130, 132 Zhong Fudao, 130, 132 Zindler, A., 191, 192 Zoneshain, L. P., 114, 120 Zumberge, J. E., 17 Zussman, J., 186, 321, 325