The sedimentary environment
1.1 Introduction
Our goal in this chapter is to introduce you to the sedimentary environment. Throughout the book we are going to emphasise the importance of interactions between physical, chemical and biological conditions. This is enormously complicated and complex, and this chapter sets the stage for much of this interaction by considering some fundamental geological, physical and chemical processes. These processes often seem simple, ordered and structured-until life starts to complicate the situation. We start with the grains that define sediments, how they are affected by hydrodynamic processes and in turn how they interact with the strong chemical gradients in marine sediments. This approach is simple but allows us to focus on how organic matter decomposition pathways vary with depth in the sediment and also how this influences the cycling of nutrients that sustain primary production. In combination, the scene set by these important physical, chemical and geological features influences not only the evolution of seabed morphology but also a wide range of ecological processes including organism dispersal and population connectivity; primary production through altered light regimes; and the release of contaminants and nutrients.
We want you to start thinking about the environmental factors that create different seafloor habitats and how they both set the scene for, and interact with, resident species and biological processes. We focus on describing why sediment grain size is one of the first variables that researchers have used to help understand variation in the structure and functioning of benthic ecosystems (Sanders, 1958, 1960). This remains a very informative variable that
we nearly always measure but it’s not a simple situation where physical processes drive biology (Snelgrove & Butman, 1995). Nevertheless, sediment grain size helps us characterise the nearseabed flow regime and sedimentary environment and it is important in regulating the transfer of particles and solutes between the benthic and pelagic realms. Grain size, and more importantly how it affects the porosity and permeability of sediments, profoundly influences the microbial decomposition of organic matter. It lays the foundation for strong gradients in reactive solutes such as oxygen that influence the cycling of nutrients critical for primary producers (Glud, 2008; Middleburg, 2019).
For a book that focusses primarily on the ecology of the larger sediment-dwelling fauna, this brief primer on sediment biogeochemistry, given from a distinctively microbial perspective, may seem unnecessary. However, we believe it provides the reader with valuable context for later chapters discussing the functional roles of sediment-dwelling fauna, and how ecological interactions influence faunal diversity and ecosystem function.
1.2 Sediment grain size plus
Marine sediments are often described as sand or mud but there is a large variation in both median particle size and particle size variation. The median grain size is useful in characterising a habitat (e.g. sandflat, mudflat or gravel bed), but it is usually the variation in different components of the distribution of particle sizes that affects ecosystem processes. As the particle size varies from fine grained silts and clays (particles <63 µm diameter) to coarser-grained sands, dramatic shifts occur in
Norkko,
densities and types of fauna found inhabiting the sediment and how they interact to influence ecosystem function. The range of particle sizes that comprise the sediment regulates the abilities of animals to construct burrows, how easily sediments are transported by waves and currents and the exchange of solutes across the sediment–water interface. Shifts in benthic diversity and ecosystem function often correlate well with grain size although the form of these relationships can vary depending on the range of sediment grain sizes included in the study. At the extremes, the finestgrained mudflats and the coarsest beach sands, the diversity will be low but important changes occur across a range of muddy sands to medium sand habitats (e.g. Douglas et al., 2018; Pratt et al., 2014).
In coastal habitats grain size varies across multiple spatial scales as a function of sediment supply from land and the near-bed hydrodynamic forces which transport and sort sediments. These processes can be in play on a regular basis associated with tidal flows or wave action, but extreme events, such as storms, can set up sediment sources and physical structures, such as sand waves, that may structure habitats for longer periods (Green & Coco, 2014; Hall, 1994; Traykovski et al., 1999). As we will see in Chapter 2 the activities of benthic organisms are also an important driver of grain size distribution. On the continental shelf, distribution reflects bathymetry where finer grains accumulate in deeper waters and coarser grains persist in an absence of fine sediment sources and/or persistent winnowing by near-bed flows. Within estuaries, variations in grain size are correlated with different habitats where muds accumulate in some areas and coarser sands in others. These basic physical settings (sediment supply and hydrodynamics) are further modified by the presence or activities of organisms (e.g. seagrass meadows and mangrove forests may trap fine particles). A detailed review of the particle grain size analysis methods for marine sediments can be found in Kenny and Sotheran (2013).
Size classes used for characterising the sedimentary environment follow the sedimentary geology Wentworth scale, which is a geometric scale that covers 32 mm, 16 mm, 8 mm, 4 mm and so on down to 0.0039 mm (Table 1.1; Leeder, 1982). For
Table
Grain size (mm) Phi (Ø) scale Sediment type
>4 < 2 Pebbles–boulders Gravel
2 1 Granule
1 0 Very coarse sand Sand
0.5
0.25
0.125
0.0625
1 Coarse sand
2 Medium sand
3 Fine sand
4 Very fine sand
0.031 >4–8 Silt Mud
<0.0039 >8–10 Clay
convenience these size classes are often transformed to the arithmetic phi (Ø) scale which is defined as the –log2 of the particle size in mm, so that a unit change in Ø represents a factor of two change in grain size.
Unless the system is high energy, such as a beach face exposed to waves, or extremely low energy, accumulating fine silts/clays, most coastal sediments represent a mixture of size classes reflecting the complex interactions between supply, physical transport and ecology. Thus, it is important to consider not only a median grain size but also a sorting index. This index describes the distribution of particles in a sample. Although there are a number of ways to measure sorting, the recommended index is the Inclusive Graphic Standard Deviation.
Whereas the median grain size and a sorting coefficient are sterile numbers, a critical element, depending on your questions, is either the percentage of fine particles or the percentage of coarse particles. In a sedimentary matrix these two percentages fundamentally change the sediment’s properties. Thus, we frequently characterise sediment based on amount of mud (or particles <63 µm in the silt/clay fraction) according to Folk’s (1954) terminology, which provides a more informative summary of the grain size distribution and physical properties (Figure 1.1).
Grain size and sorting alone are not the only sediment characteristics useful to measure (Box 1.1). The other key measures are the sediment porosity and permeability. Porosity (Ø) is a measure of the void space between grains (VP) but is presented as a ratio
1.1 Wentworth description of sediments based on particle diameter size and the phi scale
Mud
Sandy mud
Slightly gravelly mud
Slightly gravelly sandy mud
Gravelly mud Sand
Muddy sand
Slightly gravelly sand
Slightly gravelly muddy sand
Gravelly muddy sand
Gravel sand
Gravel
Muddy gravel
Muddy sandy gravel
Sandy gravel
in terms of the bulk volume of the sediment (VT; i.e. Ø = VP/VT) with values ranging from 0 to 1. It can be easily measured gravimetrically (see Kenny & Sotheran, 2013). The relationship between grain size distribution and porosity is somewhat counterintuitive; a sandy sediment will have a greater porosity than a silty sand because the silt particles clog the interstitial voids, but then as the silt/clay content increases so does the porosity as the water content increases (e.g. Kamann et al., 2007). Permeability (K; m2) is an indication of how easily water can move through the pore spaces and can be measured as the rate at which water under a constant head of pressure passes through a sediment core of known diameter and length (Klute & Dirksen, 1986). Porosity and permeability are not necessarily correlated because permeability is a function of the size of the pore spaces and their interconnectivity. Muddy sediments can have a high porosity (water content) but, due to the very small pore spaces and lack of connectivity, have a low permeability. Sandy sediment on the other hand can have a high porosity and permeability.
Box 1.1 Sediments can also be defined as cohesive or non-cohesive
In non-cohesive sediments the individual grains can be eroded whereas cohesive sediments stick together and do not behave as separate particles. This is a property determined largely by the silt/clay fraction. Non-cohesive sediments are generally sands with a low silt/clay content (< ~10%) and are characterised by a high permeability. However, as the silt/clay content increases, filling the void spaces and ultimately binding the sediment together, due to electromagnetic forces, it creates a cohesive sediment with low permeability. The distinction between these two sediment types is an important one because it helps define the erosive behaviour when subjected to increasing currents. It also regulates the non-biologically mediated transport of solutes (such as oxygen and nutrients) across the sediment–water interface. In cohesive sediments the process is slow and dominated by molecular diffusion where solutes must follow a twisted path through the sediment matrix, whereas in permeable sediments near-bed flows and pressure variations can induce much more rapid exchanges of solutes (see section 1.4).
Figure 1.1 Simplified Folk classification system for sediments with mixed particle size ranges. From Figure 1, Long (2006).
In many coastal settings sediment grain size is not uniform. Whether these sediment mixtures behave cohesively or non-cohesively depends on many factors including the fraction of fines; compaction and the size ratio of the small and large fractions; and, of course, ecological processes (see, e.g. Bartzke et al., 2013; Le Hir et al., 2007; Jacobs et al., 2011; Staudt et al., 2017 and references therein; Box 1). The important thing is that relatively small amounts of fine sediments added to the surface of sands can cause a shift between these two behaviours—a small amount of silt and clay can make a big difference! If we measure grain size of the nearsurface sediments (0–2 cm) we may see this effect, but if we take deeper cores to assess the average grain size of the sediments we may dilute the effects of these surficial deposits. This is important because in many coastal areas fine sediment addition has changed or is changing grain size distribution (see Chapter 11).
The larger plants and animals can radically change the physical characteristics of sediments, for example, by creating faecal pellets that bind small particles into larger ones or selecting certain grain sizes to build tubes (see Chapter 2). This led to consideration about how we should actually define grain size, but measuring the size of the physical particles (involving removal of organic matter and disaggregation) gives us more consistent and comparable data. Nevertheless, it is important to remember that realised grain size distributions, porosity and permeability may be quite different than those predicted from inorganic particle size analysis.
1.3 Flow, waves and the benthic boundary layer
The characteristics of the near-bed flow play a key role in regulating solute and particle fluxes to and from the water column and thus the degree of connectivity between the benthic and pelagic environments. The interaction between flow, the sediment and organisms inhabiting the seafloor influences feeding rates of benthic organisms, supply of organic matter, resuspension potential of sediments and thus disturbance regime, recruitment and the strength of biotic interactions such as adult–juvenile
interactions. The forces generating flows on the coastal seabed vary widely in their spatial and temporal scales and include tides, waves, upwelling, storms and coastal currents. The importance of flow interactions with microbes, larger plants and animals will be covered in Chapter 2, here we will give a general overview of near-bed flows and their consequences for particle and solute fluxes. -see Nowell and Jumars (1984), Jumars (1993) and Boudreau and Jorgensen (2001) for an in-depth introduction to the physics and consequences of benthic processes.
Whether flow is likely to be turbulent or laminar can be assessed by calculating the Reynolds number (Re), a dimensionless number (i.e. it has no units so is independent of the scale of observation). It compares the factors responsible for generating turbulence (fluid speed and an appropriate length scale that the fluid is interacting with) to those that dampen it out (viscosity of the fluid). Estimating the Re for a given situation gives a first-order approximation as to how quickly mixing will occur and can also give an indication of the boundary layer characteristics (Box 1.2).
In coastal environments the thickness and structure of the benthic boundary layer (BBL) are highly dynamic, due to spatial and temporal variations in flow speed and bottom roughness. Under slow, steady currents the BBL thickness may be tens of cm but under high flows it may only extend a few cm above the seabed. Furthermore, as flow speed and/ or bottom roughness increases, the viscous and diffusive layers may disappear altogether and the logarithmic turbulent layer penetrates all the way to the bed.
The steepness of the velocity gradient in the BBL determines the force exerted at the seabed which influences particle deposition and resuspension. This force is represented by the bed shear stress (τ; N m−2 s−1) and it is probably the single most important parameter to know when thinking about flow–sea bed interactions. It is commonly estimated from vertical profiles of time-averaged velocity or measurements of turbulence in the BBL (e.g. Kim et al., 2000) and is often reported as shear velocity ( u* / = tr where ρ is the fluid density). The reason for this conversion is that the shear stress is expressed in units of velocity and therefore can be
Box 1.2 The benthic boundary layer
A fundamental characteristic of the flow close to the seabed is the presence of a benthic boundary layer (BBL) and the flow characteristics of this layer are worth describing because this is where interactions between the sediments and water column actually happen. If we assume a constant current velocity then as we approach the seabed the frictional effects of the sediment extract momentum and the flow velocity begins to slow down until at the seabed (with stationary sediment) it approaches zero (Figure 1.2).
This reduction in flow changes the characteristics of the fluid environment and the BBL can be divided up into different sections based on these. The outer logarithmic layer is characterised by a log decline in the mean flow velocity toward the bed. Nominally the top of this layer is at 99% of the freestream flow speed and represents the extent of the
Outer layer
seabed influence on flow speed. In this layer, vertical mixing is rapid because of turbulence (high-frequency irregular motions—eddies—of water parcels) that pervades the fluid undergoing shear at the bed. As we move closer to the seabed, the frictional effects of the sediment begin to exert a greater impact on the fluid and the turbulence is dampened. This is called the viscous sublayer and vertical mixing in this layer is a combination of much slower molecular diffusion and smaller-scale turbulence and as a consequence is much slower. Finally, right close to the bed we have the diffusive sublayer which is where the frictional effects have dampened out all the turbulence and viscous effects (stickiness between fluid layers) dominate. This means flow tends to follow streamlines, turbulence is absent and vertical transport of solutes is dominated by molecular diffusion.
Logarithmic layer
Fully turbulent, rapid & intense mixing
V iscous layer
Weak turbulence, vertical mixing = molecular diffusion
Diffusive sub-layer
No turbulence, mixing by molecular diffusion
used to estimate a Reynolds roughness number (Re*) which characterises the nature of the flow at the seabed. When calculating Re* the appropriate length scale will be the feature that dominates the seabed roughness at the scale of interest, most commonly median grain diameter but it can also be ripples or biogenic structure protruding off the seabed (for plants and fauna; see Chapter 2). When Re* < 3.5 the BBL will be hydraulically smooth with viscous and diffusive sublayers, but at values >100 the flow is hydraulically rough and turbulent with no sublayers. The nice thing about Re* (and other dimensionless numbers) is that it highlights how different combinations of bed roughness and u can generate similar BBL conditions.
A key modifier of BBL dynamics in coastal settings is the presence of short-period (1–5 s) winddriven waves and longer-period (10–50 s) swell waves on exposed coasts. When the wave orbitals penetrate to the bed (a function of the water depth, wave height and frequency), they add a periodic oscillation of flow acceleration and deceleration in the BBL created by the mean current. These oscillations generate additional shear in the BBL and as a consequence wave–current boundary layers are virtually always hydraulically rough. The detailed physics of the wave–current boundary layer are complicated but it is worthwhile noting their importance for ecological processes. The additional shear generated by waves is often needed to initiate
Figure 1.2 Different layers and flow characteristics of the benthic boundary layer.
sediment resuspension and the associated transport/ dispersal of organic matter, sediment porewater nutrients and organisms. As the waves are primarily generated by passing weather systems this introduces a stochastic element into the frequency and intensity of sediment transport (see Hall, 1994 for an ecologically focussed review).
1.4 Consequences of the BBL on bio-physical processes
The propensity of a sediment to be mobilised is a function of the sediment grain size (and sorting), seabed roughness and the force exerted on it by the BBL. Figure 1.3 shows a simplified relationship between beds of uniform grain size and the flow velocity required to erode them. There are fundamental differences in the response of non-cohesive and cohesive sediments to increasing bed shear stress. Non-cohesive sandy sediment erosion initially starts as bedload with grains bouncing along the sediment surface (taking microbes and animals with them). As the force increases there is enough energy to move particles into suspension and then
they can be carried along with the currents. Naturally, as grain size increases the shear stress required to erode non-cohesive sediment also increases, a simple product of the fact that larger particles are heavier and more difficult to move. The behavioural response of cohesive sediments to increasing shear stress is different and they behave more like a bed of jelly than individual particles. As the erosion threshold (the point at which the applied shear stress causes the sediment bed to release particles) is exceeded, sediment is ejected (often as aggregates) directly into the water column (i.e. there is no bedload transport). Also, somewhat counterintuitively, the force required to erode a cohesive bed increases with decreasing grain size because of increasing inter-particle electromagnetic forces and decreasing bed roughness which reduces the force that can be exerted on particles.
Much of our understanding of sediment resuspension and critical erosion thresholds is based on laboratory studies of single grain sizes. In the real world, coastal sediments are often comprised of mixed assemblages of fine and coarse grains which alters the cohesive properties, and the ability of water to penetrate the seabed, altering the erosion
Clay Silt Sand Pebble
Figure 1.3 Annotated Hjulström curve showing transport properties as a function of grain size (see Hjulström, 1935).
dynamics. For example, Bartzke et al. (2013) showed that even small amounts of silt added to a sandy sediment (a few percent by weight, and well below the amount needed to cause cohesion) filled the void spaces between sand grains and substantially increased the flow speed needed to mobilise the bed. Superimposed on this is the capability of benthic organisms to alter the erosion properties of sediments (see Chapter 2). This means it is very difficult to predict erosion thresholds for coastal sediments based only on the physical grain size distribution.
At flows below the erosion threshold, the BBL can still play a crucial role in regulating solute and particle fluxes across the sediment–water interface. In permeable sediments (i.e. K > ~ 10−12 m2) pressure differentials generated by passing waves and/or flow over topography such as ripples induce regions of high and low pressure in the sediment, which drives a porewater exchange between the sediment (to depths of several cm) and the overlying
water column. The magnitude of this exchange is a function of the sediment permeability and the flows/topography generating the pressure gradient (Huettel & Webster, 2001), but in terms of solute exchange it is many times more efficient than molecular diffusion. Advective porewater exchange also introduces organic matter (Figure 1.4) into the sediment simultaneously with oxygen where it undergoes rapid decomposition by microbes. As a result, nutrients are released back to the water column. Santos et al. (2012) have estimated that the process of advective porewater exchange in permeable coastal sediments results in the entire volume of the ocean being filtered every 3000 years. It is this process combined with high rates of primary production which contributes to the importance of coastal sediments in global biogeochemical cycles. The sediment surface is a region of strong and active vertical gradients in reactive solutes such as oxygen (see section 1.5). In non-permeable cohesive sediments (i.e. K < ~ 10−12 m2) where molecular
Figure 1.4 Diatom biomass measured by chlorophyll a accumulations with increasing bed geometry (a–c). (d) and (e) show a cross section through a large ripple trough and crest respectively. From Figure 1, Pilditch and Miller (2006). Reprinted from Continental Shelf Research, 26 (15), Pilditch C A, and Miller D C, Phytoplankton Deposition to Permeable Sediments under Oscillatory Flow: Effects of Ripple Geometry and Resuspension, 1806–1825., © 2006 reused with permission from Elsevier.
diffusion physically dominates solute exchange, BBL dynamics can still influence solute concentration gradients. The demand for oxygen in marine sediments is high due to aerobic bacterial activity and the oxidation of reduced solutes moving up from deeper in the sediments. This demand creates an oxygen gradient not only in the sediment but one that can extend into the sublayers of the BBL. The oxygen gradient in the near-bed water is known as the diffusion boundary layer (DBL) and its thickness (typically <1 mm) is regulated by a combination of the sediment oxygen demand, concentration in the overlying water column and the BBL dynamics. Because the DBL is governed in part by the sediment oxygen demand, its thickness may not correspond directly to the diffusive and viscous sublayers of the BBL. Under slow flow, this concentration gradient can extend beyond the diffusion sublayer of the BBL. Under fast, turbulent flows the DBL may disappear altogether. The presence or absence of the DBL will influence oxygen availability in surface sediments with implications for aerobic bacterial activity (see Glud, 2008 for an in-depth discussion).
Near-bed flows also play a critical role in the dispersal of benthic invertebrates and, as such, an important role in recovery from disturbances and meta-community dynamics (Chapter 3). Although dispersal is discussed in many places within the book, related to aspects of recovery, design and biotic interactions, here we briefly describe it from a flow perspective. Traditionally, the pelagic larval phase was considered most important for dispersal as this was a period that they could potentially be moved long distances by currents. However, storms may move organisms around passively and it is not uncommon to see benthic animals washed up on beaches. Under more typical conditions behaviour can influence the effect of waves and currents, e.g. crawling onto the sediment surface can increase the chances of hitching a ride with the flow (e.g. Lundquist et al., 2004). This is often particularly important for juvenile life stages where the organisms are small and easily transported. For example, the bivalve Macomona when it is 0.5–3 mm long will move onto the sediment surface and orientate itself so that its shell behaves a bit like a wing; once it is ready to go it excretes long proteinaceous threads
that act like parachutes lifting it off the sediment surface, allowing dispersal over large scales (100s–1000s of m). Many benthic species exhibit this post-settlement dispersal, and what is clear from field studies is that the diversity and magnitude of dispersers are functions of the near-bed hydrodynamics (e.g. Lundquist et al., 2006; Valanko et al., 2010; see Pilditch et al., 2015 for a review). There is still much to be done to understand how temporal and spatial variability in the interacting factors regulating dispersal (sediment properties, behaviour, species interactions and hydrodynamics) influence dispersal but such studies are critical to understanding the dynamics of benthic populations and communities (Pilditch et al., 2015).
1.5 Organic matter
The quality and quantity of the organic matter arriving at the sediment surface are critical environmental variables, providing fuel directly for metazoans or indirectly by fuelling microbes for decomposition that then become food for others. In general, the organic content of sediments increases with decreasing grain size because the lower flows allow the accumulation of low-density particles but also the increased surface area promotes high bacterial biomass.
In coastal environments, organic matter has numerous sources: primary production in the water column, detritus from macrophytes such as kelp, seagrass and mangroves as well as inputs from terrestrial ecosystems. In many benthic ecosystems, organic matter fuels secondary production. Except in the most turbid regions, the production of the microphytobenthos at the sediment–water interface is the major source of organic material (Cahoon, 2002; Hope et al., 2020; Miller et al., 1996).
The input of organic matter is the base of the food chain of marine sediments and the amount varies with many factors, including the productivity of the overlying water column. Food quality is a big issue and can be coarsely quantified by the carbon to nitrogen ratio. In general phytoplankton and bacteria have lower C:N ratios, typically <10, whereas macrophyte detritus typically has C:N ratios >20. Consumers prefer energy sources with lower C:N
ratios (ideally as close to seven as possible, the ratio in animal tissue) because higher ratios mean more time (and energy) are expended acquiring sufficient food to satisfy their N requirements. As well as these coarse measures of food quality, it is the plants that synthesise the essential fatty acids needed by animals (Antonio & Richouz, 2014; Galloway et al., 2012). Faecal pellets from animals living in the water column or suspension feeders living on the seabed are an important source of organic matter input into seafloor communities. Faecal pellets have much higher settling velocities (e.g. Giles & Pilditch, 2004) than the food that produces them and even after digestion (and any decomposition occurring on the way down) have a higher organic content than water column particles naturally settling out. This biological pump is an important component of benthic–pelagic coupling: i.e. in this context, the delivery of water column organic matter to the seabed where it is remineralised by bacteria, releasing nutrients that fuel pelagic production (Graf & Rosenberg, 1997).
1.6 Light and benthic primary production
In shallow coastal waters where light hits the seabed the microphytobenthos and larger macrophytes (macroalgae and seagrasses) can flourish and give rise to local production. The amount of light hitting the seabed is a function of atmospheric conditions, latitude, time of day and coastal topographic features that create shade, water depth and clarity. Water clarity is strongly affected by phytoplankton in the water column, dissolved organic matter like tannins from land that stain the water and importantly the input and resuspension of fine sediments. The amount of light hitting the seafloor regulates benthic primary production and in turbid estuaries the loss of this production has been a contributing factor to accelerating eutrophication (Chapter 11).
Although on a per-area basis primary production by benthic macrophytes can greatly exceed that of unvegetated habitats, it is the ubiquitous microphytobenthos (MPB) that fuel many coastal foodwebs (e.g. Christianen et al., 2017; Jones et al., 2017). MPB consist of diatoms, dinoflagellates and cyanobacteria (McIntyre et al., 1996) growing within the first several mm of the surface sediment and can often
be seen as a greenish or brownish tinge. In more oligotrophic settings MPB production can exceed pelagic production because they have ready access to nutrients stored in the sediment porewater. Their value in coastal foodwebs arises because they are easily digestible, rich in lipids and proteins, have high turnover rates and have not undergone pre-processing by pelagic consumers and so do not arrive as faecal pellets (Hope et al., 2020). In the Dutch Wadden Sea a stable isotope study showed that MPB were the most important food source for benthic invertebrates and given these organisms support higher trophic levels (fish, birds and seals) this further underlines their importance (Christianen et al., 2017).
MPB are ecologically important beyond their value as food for higher trophic levels (Hope et al., 2020). Light is attenuated rapidly in sediments and a lack of nutrients in the photic zone means MPB undertake vertical migrations to access them (Consalvey et al., 2004). To aid these migrations, diatoms in particular excrete extracellular polysaccharides (EPS), a sticky mucus-like substance that binds the sediment together. In cohesive sediments (or non-cohesive sediments with a substantial silt/ clay content) the EPS can fill the void spaces, binding sediment and particles, thus leading to a marked increase in sediment stability, making it less prone to erosion (see Chapter 2). Photosynthesis in surface sediment layers can also alter the distribution of oxygen, increasing the volume of sediment that then supports more efficient decomposition of organic matter by aerobic bacteria, affecting nutrient cycling (see section 1.8). MPB also trap nutrients at the sediment surface, preventing their release to the water column. In shallow coastal systems undergoing eutrophication, this sink of nutrients helps slow the eutrophication spiral where continual release of nutrients from decomposing organic matter in the sediments fuels further pelagic production.
1.7 Sediment biogeochemistry
Even at small scales (m2) coastal marine sediments are extremely heterogeneous with respect to their physical properties (e.g. grain size, sorting) and concentrations of reactive solutes such as oxygen.
This heterogeneity arises primarily as a result of the interactions between the activities of the resident macrofauna community and the microbial community (both the bacteria that are decomposing organic matter and the photosynthesising MPB). For example, burrow irrigation by shrimps (e.g. D’Andrea & DeWitt, 2009) may deliver oxygen deep into the sediments whereas bulldozing heart urchins leave faecal pellets that become hotspots of bacterial activity (Lohrer et al., 2004; Solan & Wigham, 2005). In Chapter 2 we explore more animal–sediment interactions but just as we have described the physical aspects of the sediment, there is value in considering the basic chemical implications of bacteria and organic matter down the sediment column.
The upper few centimeters of the sediment are one of the most extreme chemical gradients on the planet where transition occurs from an aerobic to an anaerobic environment. These gradients are driven by the physical characteristics of the sediment as well as the microbial communities and, in a world without large plants and animals, are vertically structured according to the bacterial respiration pathways during the decomposition of organic matter. The coastal sediments (which comprise <10 percent of the oceanic seafloor) are estimated to be responsible for processing ~30 percent of the world’s oceanic carbon (Smith & Hollibaugh, 1998) and during the decomposition regenertating essential nutrients. As we will see later in the book,
human-induced changes in sediment organic matter decomposition and nutrient cycling pathways can extend into the water column, negatively impacting the entire ecosystem.
If you were to carefully dig a small hole in a typical coastal sediment you would most likely see three distinct layers distinguished by their colour (and smell). These represent substantial shifts in the metabolic pathways of bacteria responsible for organic matter decomposition. The first major active zone in most coastal sediments is the surface oxic layer which appears as a yellow/brown layer that morphs into grey before immediately being replaced by reduced black anoxic sediment with a smell of rotten eggs (hydrogen sulphide). In the decomposition pathways, aerobic respiration in the surface yellow/brown layer releases the most energy and aerobes will outcompete other respiration pathways. The concentration of oxygen in a microbially dominated sediment decreases exponentially from the sediment surface to the redox potential discontinuity (RPD) where no free oxygen exists (Figure 1.5). The availability of oxygen is measured by the redox potential (Eh); once the oxygen is used up anaerobic processes dominate and control chemical speciation.
The thickness of the oxic reaction zones in sediments is variable and as a consequence so too is the efficiency with which organic matter is remineralised. The thickness of the oxic layer is a balance between supply and demand, the sediment grain
Figure 1.5 Sediment profiles of redox potential, pH, oxygen, hydrogen sulphide, iron, methane, carbon dioxide and nitrogenous species.
size which determines diffusion rates, and the quality and quantity of organic matter which determine the rates of bacterial metabolism. In finer-grained sediments diffusion rates are slower and oxygen decreases more rapidly. In permeable sediments diffusion rates are higher but there is also advective porewater exchange driven by pressure differentials. If organic loading is high, then the RPD can extend all the way to the surface and the sediments (and even the overlying water) become anoxic, limiting the higher and larger lifeforms. The loss of the oxic sediment layer also decreases the efficiency of organic matter degradation metabolism and alters the cycling of key nutrients (see section 1.8). In terms of impacts in coastal waters, increased sedimentation in many parts of the world plus eutrophication are impacting on the biogeochemistry (see Chapter 11). The presence of MPB can also induce diurnal changes in oxygen availability and during light reactions the production of oxygen can lead to a sub-surface maximum which disappears at night. Fluctuations in light intensity result in temporal variations in oxygen availability (similar fluctuations are seen through the activities of larger fauna—see Chapter 2) and may influence nitrogen cycling, a process that depends on coupled reactions in the presence and absence of oxygen (see section 1.8).
Beneath the RPD, in the absence of oxygen, other terminal electron acceptors (oxidants) are utilised by anaerobic bacteria sequentially with depth based on their standard redox potential and yield of free energy (Table 1.2; Figure 1.5). This creates a depth sequence in the oxidants consumed for the mineralisation of organic matter by bacteria, from oxygen to nitrate, manganese oxide, iron oxides, sulphate and finally carbon dioxide. The by-product of sulphate reduction is hydrogen sulphide, which gives the sediment that rotten eggs smell. Hydrogen sulphide and sulphide ions are toxic to nearly all metazoans so organisms living beneath the RPD layer obviously need to pump oxygen down. Because the oxic layer in sediments is relatively small in most of the coastal ocean it is estimated that most of the organic matter is decomposed via manganese, iron and sulphate reduction in anoxic sediments (Canfield, 1993; Jørgensen, 1982). The fate of organic matter as it passes through these different levels determines burial rates and ultimately
Table 1.2 Organic matter (CH2O) oxidation pathways mediated by bacteria in the sediment and their free energy yields (ΔG°). Table adapted from Boudreau and Jørgensen (2001)
Reaction kJ mol−1
Oxic respiration: CH2O + O2 → CO2 + H2O 479
Denitrification:
5CH2O + 4NO3 → 2N2 + 4HCO3 + CO2 + 3H2O 453
Mn-oxide reduction: CH2O + 3CO2 + H2O + 2MnO2 → 2Mn2+ + 4HCO3 349
Fe-oxide reduction:
CH2O + 7CO2 + 4Fe(OH)3 → 4Fe2+ + 8HCO3 + 3H2O 114
Sulphate reduction: 2CH2O + SO4 2− → H2S + 2HCO3 77
Methane production: HCO3 + 4H2 + H+ → CH4 + 3H2O CH3COO + H+ → CH4 + CO2 136 28
the quantity of carbon stored in the sediment (Middleburg, 2019).
The by-products of anaerobic respiration/decomposition include reduced inorganic compounds in the porewater that diffuse upwards toward the sediment surface. It is interesting to note that as you move deeper in the sediment more free energy (once oxidised) is tied up in the reduced compounds, a point highlighted by the fact that methane is a by-product of the fermentation processes at depth. As these reduced compounds diffuse toward the RPD chemosynthesis can occur whereby bacteria mediate the oxidation of reduced iron, manganese and in particular sulphides to release energy that is then used to fix inorganic carbon. This form of primary production is similar to that which fuels hydrothermal vent ecosystems in the deep sea. In shallow water depths where the RPD extends to the sediment surface and light penetrates to the seabed phototrophic sulphur bacteria and cyanobacteria are also capable of carbon fixation, utilising the diffusing sulphides (see Jørgensen et al., 2019 for a recent review).
1.8 Nutrient cycling
The decomposition of organic matter within the sediments results in the regeneration of inorganic nutrients essential for primary production. It has
been estimated that in some coastal and shelf ecosystems up to 30–50 per cent of the pelagic primary production is supported by nutrients regenerated from within the sediments (Nixon, 1981; Pilskaln et al., 1998). It is beyond the scope of this section to describe in detail the cycling of all biologically important elements and we refer readers to excellent summaries in Aller (1994), Middleburg (2019; carbon), Jørgensen et al. (2019; sulphur) and Delaney (1998; phosphorus) for entry points into this literature. We have chosen however to provide some detail on the sediment nitrogen (N) cycle because it is the main nutrient limiting primary production in coastal waters, of which the sediment is a key source (Nixon, 1981). Also, the sediment represents the most important biological pathway to remove excess N, which is the cause of eutrophication (see Chapter 11).
Bacteria in the sediment mediate multidirectional oxidation and reduction reactions that transform organic N into a number of different inorganic forms. A key feature regulating these transformations is the tight coupling between the oxic and anoxic zones in the sediment. The sediment porewater is rich in ammonium from the decomposition process (ammonification) and excretion by organisms. Nitrification, purely an aerobic process, transforms the ammonium into nitrate via nitrite. At each
stage of the transformation process inorganic N can diffuse upwards where it can be utilised by MPB at the sediment–water interface and, if not trapped there, by pelagic primary producers. Ammonium, nitrite and nitrate can also diffuse downwards (the direction is dependent on the concentration gradient and therefore a function of solute sources, sinks and diffusion rates) into the anoxic sediment where it can undergo a series of reduction reactions.
Of particular significance to systems experiencing excess N inputs are the pathways that reduce N to biologically inert forms. Denitrification converts nitrite and nitrate to nitrogen or nitrous oxide gas which are not readily assimilated by primary producers and are ultimately released back into the atmosphere. In the anoxic zone, dissimilatory nitrate reduction to ammonium (DNRA) can also occur where it can diffuse back into the oxic sediments or in the presence of nitrite be converted via anaerobic ammonium oxidation (anammox) to nitrogen gas (see Devol, 2015 for a review). The importance of anammox as a pathway of N removal remains relatively unknown in marine sediments, whereas in coastal sediments denitrification tends to be quantified more frequently. The small number of studies focussed on anammox in coastal sediments suggest it can account for 10–40 per cent of the total N2 production (Devol, 2015). The removal
Figure 1.6 A conceptual model of the sediment nitrogen cycle in coastal sediments showing transformations specific to the oxic and anoxic sediment layers. Note that nitrification is carried out by aerobic bacteria and provides the nitrate that is converted to inert forms of N in the anoxic sediments. When the oxic layer is lost from sediments (e.g. due to organic loading) the ability of the sediments to remove bio-available N is compromised. A: ammonification; Anammox: anaerobic ammonium oxidation; DNRA: dissimilatory nitrate reduction to ammonium; PON: particulate organic nitrogen. Adapted from Devol (2015).
of nitrogen through denitrification in estuaries can account for 10–80 per cent of the inputs, thereby representing a significant sink providing resilience to eutrophication (Nixon et al., 1996).
The ability of sediments to remove excess N via denitrification is dependent on the presence of an oxic–anoxic interface. The simple one-dimensional picture of N-cycling in coastal sediments shown in Figure 1.6 masks the complexity introduced by the activities of larger fauna that can greatly extend this interface through building and irrigating burrows or bulldozing their way through sediment, both activities that can stimulate denitrification (see Chapters 2 and 9). Often in eutrophic systems excess organic matter inputs result in the loss of the sediment oxic layer (RPD rises into the water column). Under these conditions denitrification is halted and the sediments become a significant source of regenerated ammonium that continues to accelerate primary production.
1.9 Close out
A lot of processes are interacting in marine sediments and the overlying water. We have highlighted the intermittent interplay between the physical processes of water flow, sediment grain size and the supply and decomposition of organic matter. These processes regulate the chemistry of this reactive layer and make a significant contribution to global cycles of nutrients and carbon, despite the biologically active layer being only a few centimeters deep. In coastal waters, the pelagic and benthic environments are tightly coupled and solute and particle exchanges across the sediment–water interface regulate important functions. Thus, the benthic and pelagic realms need to be seen as one system. Conversely, in deep-sea habitats the benthic influence on pelagic habitats is lost (but not vice versa) because of water circulation times.
The physico-chemical processes we have highlighted are often well supported by theory but we have limited empirical data. This can result in us missing important processes (e.g. anammox was only discovered about 20 years ago; Kuenen, 2008). Our understanding of how these basic processes operate can also be biased by data being collected in a restricted geographic region (Vieillard et al., 2020).
Therefore, although these processes have been under study for many years, there is still plenty of work to be done. However, physico-chemical processes do not operate in a vacuum. Although there are ‘dead’ zones in our marine systems, they are fortunately rare and instead physico-chemical processes interact with, and are often driven by, the plants and animals that live in and on the sediment (see Chapter 2). The interplay between these factors is an area rich in opportunities for future research.
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