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Sediment Retention and Release
3172 Ecological Processes | Sediment Retention and Release
Sediment Retention and Release F J Black, C Gallon, and A R Flegal, University of California Santa Cruz, Santa Cruz, CA, USA ª 2008 Elsevier B.V. All rights reserved.
Introduction Chemical Reactions Transport Processes
Measuring Nutrient and Contaminant Fluxes in Sediments Further Reading
Introduction
Chemical Reactions
Both the retention and release of nutrients and contaminants by sediments are controlled by complex physical and biogeochemical interactions. Sediments, porewaters, and overlying waters are dynamic matrices that alternatively serve as sources and sinks for elements and compounds, as physical and biogeochemical conditions vary. Some of those conditions markedly change over small spatial scales, from nanometers to millimeters, and over time, from nanoseconds to millennia. Consequently, rates of scavenging and mobilization of elements and compounds to and from sediments, respectively, may also exhibit pronounced spatial and temporal variability. This summary, therefore, briefly describes some of the principal factors involved in the uptake and release of nutrients and contaminants from sediments, as well as the diffusive and advective dispersion of those constituents in sediment porewaters and overlying waters. This includes a synopsis of chemical reactions that affect the speciation, cycling, and bioavailability of constituents in sediments that can result from the degradation of organic matter, other redox transformations, the dissolution and precipitation of mineral phases, and sorption processes, including cation exchange. It also includes brief comments on transport processes related to geochemical and physical characteristics of the media, as well as the influence of benthic organisms on the distribution and dispersion of those constituents in sediments and associated waters. The complex integration of all the preceding factors determines the physical and biogeochemical cycling of nutrients and contaminants in sediments. Directions and rates of those cycles are governed by prevailing chemical and physical conditions, and are often mediated or driven by microbial communities. While those constituents may initially be entrained in sediments by geochemical sorption, authigenic precipitation, and biological scavenging, they may subsequently be remobilized from the sediments into dissolved and colloidal phases by mineral weathering, desorption, and decomposition.
Organic Matter Decomposition When suspended particulate, colloidal, and dissolved organic matter (DOM) are scavenged from surface waters and deposited to benthic sediments they encounter the sediment–water interface, an important region in the biogeochemical cycling of many elements and compounds. Most biologically mediated oxidation and mineralization of organic matter, referred to as early diagenesis, occurs within that interface, which typically extends a few centimeters into the surface sediments. This early diagenesis is carried out by microbes using the most energetically favorable oxidant available, which in aerobic environments is molecular oxygen, O2. This is illustrated by the simplified oxidation of organic matter, based on Redfield’s ratio of C:N:P of 106:16:1: ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 138 O2 ! 106 CO2 þ 16 HNO3 þ H3 PO4 þ 122 H2 O
The G of the above reaction is 3190 kJ mol1, which is thermodynamically favorable and yields substantial free energy. The oxidation of organic material via this reaction will be dominated by aerobic bacteria as long as sufficient O2 is available. However, the rate of O2 consumption will exceed the rate of O2 diffusion from overlying waters to sediments at some sediment depth. In highly productive aquatic systems and those impacted by relatively large discharges of organic waste, the depletion of O2 may occur at or above the sediment–water interface. Microbial depletion of O2 leads to the formation of suboxic conditions in deeper sediments, where further oxidation of organic material continues with microbes using the next most energetically favorable terminal electron acceptor. As one oxidant is consumed, the next most energetically favorable oxidant is utilized. MnO2 and NO 3 reduction occur following the loss of O2, sequentially followed by iron reduction, sulfate reduction, and methanogenesis. This sequence is illustrated by the following balanced equations for the oxidation of organic material and the Gibbs
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free energy associated with each terminal electron acceptor: Manganese oxide reduction: G ¼ 3090 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 236 MnO2 þ 472 Hþ ! 236 Mn2þ þ 106 CO2 þ 8 N2 þ H3 PO4 þ 366 H2 O
Nitrate reduction: G ¼ 3030 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 94:4 HNO3 ! 106 CO2 þ 55:2 N2 þ H3 PO4 þ 177:2 H2 O
Iron reduction: G ¼ 1330 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 424 FeOOH þ 848 Hþ ! 424 Fe2þ þ 106 CO2 þ 16 NH3 þ H3 PO4 þ 742 H2 O
Sulfate reduction: G ¼ 380 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 53 SO24 – ! 53 S2 – þ 106 CO2 þ 16 NH3 þ H3 PO4 þ 106 H2 O
Methanogenesis: G ¼ 350 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ ! 53 CH4 þ 53 CO2 þ 16 NH3 þ H3 PO4
The distribution of redox zones with depth below the sediment–water interface where each of the above respiration pathways occurs varies with physical and biogeochemical conditions. In naturally eutrophic and contaminated sediments, where inputs of organic matter are relatively high, O2 penetration by diffusion from overlying waters may be limited to a few millimeters. In oligotrophic freshwater and deep-sea sediments, where organic matter inputs are much lower, O2 may diffuse a few centimeters below the sediment–water interface. However, the advection of surface waters into the sediments due to bioturbation commonly accounts for greater and uneven penetration of O2 and other terminal electron acceptors in benthic sediments. Consequently, the distribution of redox zones in sediments may be highly heterogeneous. Multiple microenvironments often exist within millimeters of each other where aerobic respiration, nitrate reduction, and sulfate reduction are carried out simultaneously at comparable depths. Since the rate of organic matter oxidation is temperature dependent due to its effect on microbial respiration rates, early diagenesis in shallow water sediments also displays seasonal variability. The oxidation of organic material in sediments during early diagenesis causes a number of chemical changes beyond those in the organic matter itself. The decomposition of the organic matter also releases other nutrients, trace elements, and contaminants which have been complexed or associated with it, when it is solubilized. All of the preceding forms of respiration also produce carbon dioxide (CO2). This increase in CO2 levels is accompanied by a decrease in pH due to the formation
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of carbonic acid (H2CO3) and bicarbonate (HCO 3 ). That decrease can, in turn, lead to the mobilization of cations adsorbed to the sediments as an increasing number of hydrogen ions compete for negatively charged adsorption sites, through cation exchange. Phosphorous, principally in the form of phosphate, and nitrogen, principally in the form of nitrate or ammonia, are liberated directly by the mineralization of organic matter, are liberated directly by the mineralization of organic matter, as well as indirectly by the reduction of iron and manganese oxyhydroxides. As a result, those two macronutrients (N and P) are often depleted in surface waters where they are commonly the limiting nutrients for primary productivity and are enriched in subsurface and porewaters, where they are released with the decomposition of organic matter. This nutrient-type profile is also paralleled by those of other trace- and micronutrients, and primary productivity is often limited by the subsequent flux of those remobilized nutrients to the euphotic zone in overlying waters. Flux rates out of anoxic lacustrine and estuarine sediments reported in the literature are generally of the order of 0–20 mg m 2 day1 for phosphate, 0–10 mg m 2 day1 for nitrate, and 0–30 mg m 2 day1 for ammonium. However, these flux rates vary substantially by environmental setting, as well as spatially and temporally. In addition, these nutrient fluxes may be reversed, from the overlying waters to sediments, under some conditions. Other Redox Reactions The redox state of sediments and associated porewaters influences the retention and release of trace elements and other compounds, as illustrated by chromium. It commonly exists in the environment in two oxidation states: trivalent chromium, Cr(III), which is an essential trace element, and hexavalent chromium, Cr(VI), which is carcinogenic. The latter species is also much more labile in most aquatic systems, and may be found at potentially toxic levels (>100 mg l 1) in some aquifers due to either natural processes or industrial pollution. While Cr(III) is more thermodynamically stable under reducing conditions, it is generally only found at low levels because it exists in solution primarily as cations (Cr3þ, Cr(OH)þ 2, Cr(OH)2þ) that have high affinities for sediments with a net negative surface charge (e.g., iron oxides and clay minerals) and Cr(III) readily precipitates as Cr(OH)3 and Cr2O3 in the pH ranges commonly encountered in the environment. In contrast, Cr(VI) forms relatively soluble oxyanions (e.g., CrO2 4 and HCrO4 ) that have less affinity for sediment surfaces and, therefore, may occur at higher concentrations in many aquifers. Although Cr(VI) may be naturally converted to Cr(III) in reducing groundwater systems, the microbial oxidation of Cr(III) to Cr(VI) at the aerobic/anaerobic interface in sediments has been
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associated with anomalously high levels of the potentially toxic species in some aquifers – as has the discharge of industrial Cr(VI) in other aquifers. Sediment redox processes also strongly affect metalsulfide minerals, which can be solubilized under oxidizing conditions. The oxidation of metal-sulfides such as cinnabar (HgS), cuprite (CuS), galena (PbS), sphalerite (ZnS), chalcopyrite (CuFeS2), and chalcocite (Cu2S) by oxygen is similar to that of pyrite (FeS2): 4 FeS2 þ 15 O2 þ 14 H2 O ! 4 FeðOHÞ3 þ 8 H2 SO4
The oxidation of the reduced Fe(II) and sulfide in pyrite results in the formation of Fe(III) in the form of iron oxide (Fe(OH)3) and sulfate in the form of sulfuric acid (H2SO4), respectively. The H2SO4 produced by that oxidation is responsible for much of the acid mine drainage (pH < 5) generated from mines and mine tailings. While the abiotic oxidation of sulfidic minerals is thermodynamically favorable in aerated waters, the kinetics are generally very slow at ambient environmental conditions because of the large activation energy required. However, the oxidation of sulfidic minerals is markedly accelerated by bacteria (e.g., Thiobacillus ferrooxidans) which obtain energy from the oxidation of pyrite. Acid mine drainage of sulfidic deposits is accompanied by the leaching and mobilization of other heavy metals. These may include arsenic, cadmium, copper, lead, manganese, mercury, selenium, and zinc, which are all relatively toxic to most aquatic organisms. Consequently, surface waters downstream from acid mine drainage may be nearly devoid of all but microbial life until pH levels increase and iron and manganese precipitate out as oxyhydroxides that scavenge those toxic elements. However, sediment burial following the deposition of iron and manganese oxyhydroxides leads to reducing conditions in subsurface sediments, which thermodynamically favor the reductive dissolution of those compounds. This process releases not only iron and manganese back into solution, but also metals and metalloids that had been adsorbed onto the oxyhydroxides. As a result, sediments with relatively high levels of contaminant metals that have been transported downstream from acid mine drainage constitute a potential source of contamination to porewaters and overlying waters. Similarly, other sediments that have relatively large amounts of metals (e.g., Cd, Cu, Pb) and metalloids (e.g., As, Se) – from natural or industrial sources – scavenged onto iron and manganese oxyhydroxides represent a potential source of pollution under reducing conditions. Again, those reducing conditions may be catalyzed by the deposition of organic matter, either following an algal bloom or the discharge of organic industrial, agricultural, and municipal wastes.
Heavy metals are not the only class of compounds which are readily adsorbed by iron and manganese oxyhydroxides. Nutrients such as phosphate and nitrate also exhibit this same behavior under oxidizing conditions, and are thus released when the sediments become reducing and the Fe(III) and Mn(IV) precipitates are solubilized to dissolved forms of Fe(II) and Mn(II). Thus, a change in the redox state of sediments to more reducing conditions can result in the release of phosphate and nitrate to overlying waters where they can promote eutrophication. As a result, lakes which receive nutrient-rich runoff from urban or agricultural areas are often mechanically aerated in order to maintain oxygenated bottom waters and surface sediments to prevent the reductive dissolution of iron and manganese oxyhydroxides and the subsequent release of nutrients. This method has also been employed successfully to promote the reformation of iron and manganese oxyhydroxides and recovery of lakes following eutrophication. Mineral Formation and Weathering The formation and weathering of the mineral matrix of sediments plays a primary role in the retention and release of ions and compounds. The formation of sediments by authigenic precipitation removes ions from solution and creates fresh surfaces for the adsorption of other inorganic and organic compounds as a secondary retention mechanism. The weathering of sediments, conversely, releases both adsorbed components and parent material into solution. The overall chemical processes of mineral precipitation and dissolution can be generally described in terms of nucleation, crystal growth, and weathering. Nucleation is often the rate-limiting step of mineral precipitation and can occur by homogeneous or heterogeneous mechanisms. During homogeneous precipitation, crystal nuclei are formed in a saturated solution by the random collision of ions. Heterogeneous nucleation involves the formation of mineral nuclei on the surfaces of reactive solids already present. Regardless of the nucleation pathway, once stable nuclei are formed, mineral growth proceeds spontaneously until the solution is no longer saturated with respect to that mineral. The rate of nucleation is controlled by the degree of supersaturation, temperature, and the geometry of initial nuclei formed or heterogeneous materials available for nuclei seeding. Nucleation rates are also dependent upon the specific interfacial free energy, which is the difference in free energy between an ion bound within the mineral matrix and an equivalent ion bound to the mineral surface. Crystal growth involves transport of reacting ions to the mineral surface, transport of reaction products away from the mineral surface, and surface interactions
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including adsorption, surface complexation, dehydration, and cation exchange. Mineral growth can be transport controlled, where growth is limited by the rate at which ions or complexes migrate to the surface via diffusion and advection, or surface controlled, where the rate-limiting step is the surface interaction involved. Most mineral formation reactions are surface-controlled processes. The nature of sediment weathering products and their subsequent release depends upon the mineralogical composition of the parent material, the chemical composition of the aqueous phase, and the nature of fluid flow. Chemical weathering reactions are often classified by the attacking agent type and the manner in which the mineral is altered. Dissolution and hydrolysis reactions are mediated by potential attacking substances including acids, oxygen-containing ligands, and water. Since many weathering reactions can be treated as acid–base reactions, the pH of sediment porewaters exerts a primary control on sediment dissolution and weathering rates. The source of acidity in sediments can be organic acids exuded by plant roots or from the decay of organic matter; sulfuric acid from sulfide oxidation; nitric, sulfuric, and hydrochloric acids in acid rain; and carbonic acid, with CO2 derived from the atmosphere or the respiration of organic matter. If simple dissolution is the primary mechanism of weathering, the process is termed congruent dissolution, as illustrated by the dissolution of quartz (SiO2(qtz)): SiO2ðqtzÞ þ 2 H2 O $ H4 SiO4ðaqÞ
This reaction is reversible, and dissolved silica in the form of silicic acid, H4SiO4(aq), can precipitate to form quartz. But the kinetics of quartz precipitation are quite slow below 70 C, and as a result silicic acid concentrations are often measured in excess of its predicted solubility of about 180 mM in low-temperature soil and sediment porewaters. Therefore, while thermodynamics predict the precipitation of quartz from many soil and sediment porewaters, the slow kinetics involved often result in the supersaturation of dissolved silica. The dissolution of calcite, CaCO3(calc), is another congruent weathering reaction: CaCO3ðcalcÞ þ H2 CO3 ! Ca2þ þ 2 HCO3–
In this process, calcium carbonate and carbonic acid react to form calcium cations, Ca2þ, and bicarbonate anions, HCO3. This reaction yields no solid phase, because the carbonate mineral undergoes complete dissolution. The equilibrium of this reaction, as well as that of the preceding reaction, is pH dependent, with higher acidity resulting in greater dissolution of the product ions into solution. If a secondary mineral forms during a chemical weathering reaction the process is referred to as incongruent dissolution. A common form this type of process takes
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is the weathering of aluminosilicates minerals, which are stable at high temperatures and pressures characteristic of the Earth’s interior, to clay minerals, which are more stable at the low temperatures and pressures characteristic of the Earth’s surface. An example of incongruent dissolution is the hydrolysis of potassium feldspar, KAlSi3O8(Kspar), to kaolinite, Al2Si2O5(OH)4(kaol): 2 KAlSi3 O8ðKsparÞ þ 2Hþ þ 9 H2 O ! Al2 Si2 O5 ðOHÞ4ðkaolÞ þ 4 H4 SiO4ðaqÞ þ 2 Kþ
The incongruent dissolution of potassium feldspar not only forms the solid-phase kaolinite, but also releases silicic acid, H4SiO4(aq), and potassium cations, Kþ, into solution. The stability of the different mineral phases is, again, dependent upon pH and fluid chemistry, with increased acidity and decreased silicic acid and potassium cation concentrations promoting feldspar weathering. Determination of which mineral phase is most stable under prevailing sediment conditions can be made with thermodynamic models and stability (Eh:pH) diagrams. Weathering reactions generally neutralize acids and release base cations (Ca2þ, Mg2þ, Naþ, Kþ) into solution. Weathering of aluminosilicates releases dissolved silica, but not comparable levels of aluminum due to its relatively low solubility around neutral pH (7). But in freshwater lakes with sediments and soils low in carbonates and low acid-neutralizing capacities that are impacted by acid rain, the pH of porewaters eventually drops to levels where the dissolution of aluminum (Al3þ) is sufficient to cause asphyxiation in fish and invertebrates. Nevertheless, aluminum is often treated as a conservative element and its concentration is compared to that of other elements in sediments to estimate levels of contamination. These qualitative estimates are based on the assumption that the concentration of aluminum in contaminated sediments is not enriched relative to its average crustal abundance ( 8.1 mg g 1), while concentrations of other elements may be enriched in contaminated sediments relative to their average crustal abundance. The resulting enrichment factor (EF) is then derived with a simple normalization of those ratios: EF ¼
½X sediment =½A1sediment ½X crustal abundance =½A1crustal abundance
where[X] is the concentration of the element of interest and [A1] is the concentration of aluminum. The presence of ligands and organic matter in sediment porewaters can increase mineral weathering by complexing cations involved in the weathering reaction. This decreases the cation’s free ion concentration and causes a thermodynamic shift which favors further mineral dissolution, as predicted by Le Chatelier’s principle. Complexation by organic matter also increases the solubility of metals via this mechanism, which acts to
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facilitate the dissolution of minerals containing metals exhibiting low solubilities at near neutral pHs, such as Fe3þ, Al3þ, and Hg2þ. Finally, microbes commonly play an important role in the rates of both mineral precipitation and weathering, and thus the sequestering or release of nutrients and contaminants in sediments. Microbes often mediate the underlying chemical reactions involved or take advantage of the energy released by the associated reactions. For example, microbes can facilitate nucleation by their cell membranes when they act as nucleation sites or when they produce an organic molecule that serves as a template for an inorganic crystal, a phenomenon known as biomineralization. Similarly, microbes often effectively control the rate of sulfide mineral oxidation, a thermodynamically favorable but kinetically slow inorganic reaction, by their production of enzymes which catalyze the reaction as a means of harnessing the energy released. Adsorption Adsorption is the result of chemical bonding between a gas, nonaqueous liquid, or dissolved compound and the solid sediment mineral matrix or organic matter. The strength and reversibility of this interaction can vary substantially depending upon the nature of the bonding. The most important substrates for adsorption in sediments are generally clay minerals, organic matter, and iron and manganese oxyhydroxides. Organic compounds with hydrophobic regions often undergo adsorption onto sediment organic matter or mineral faces by Van der Waals or hydrophobic interactions. Adsorption of metals and other ions is generally via ionic or electrostatic interactions, with bonding sometimes being more covalent in nature. Metals that are complexed by DOM will also be retained by sediments if the organic matter with which they are associated is adsorbed onto sediment particles. The adsorptive behavior of sediment is controlled by its organic content, surface area, and thus particle size, pH, and the type and density of adsorption sites. Sediments with more organic coatings generally absorb relatively large amounts of organic and hydrophobic compounds. Fine-grained sediments with relatively small particle sizes and large total surface areas have a relatively high number of adsorption sites, and thus relatively greater adsorption capacity than that of largergrained sediments. Since clay particles exhibit both a high surface area as well as a net negative charge density, they are generally more effective at adsorbing dissolved metals and other cations than coarser-grained sediments. The mineralogy of sediments also influences the density and type of adsorption and exchange sites present, as does their moisture content, which can play an important role in adsorption, with increasing water content being associated with decreasing retention of nonpolar compounds.
c
d
b e a
f h
g
Figure 1 Stylized representation of a clay particle with surface coatings: (a) iron oxyhydroxides, (b) manganese oxyhydroxides, (c) aluminosilicates, (d) other inorganics (e.g., calcite, apatite), (e) organic matter, (f) microorganisms (e.g., bacteria), (g) cross section of the clay particle showing net negative charge on its surface, (h) clay surface with a net negative charge resulting in the subsequent adsorption of cations and other coatings as detailed above. Although each of these coatings is shown as being discrete, they generally overlap and overlay each other.
A stylized sequencing of some of those processes is illustrated in Figure 1. Chemical adsorption of cations can also be treated as a surface complexation reaction in which lone electron pairs of primarily oxygen, nitrogen, and sulfur atoms at the solid surface are donated to metals and other cations to form surface complexes. In this model, surface complexing sites compete with dissolved complexing agents for cations, both being capable of forming inner or outer sphere complexes. As previously noted, surface hydroxyl groups of iron and manganese oxyhydroxide solids exhibit strong affinities for many trace metals that are scavenged by sediments. But, the presence of dissolved ligands capable of outcompeting sediment surface complexing sites can lead to metal desorption and mineral dissolution. The pH of sediment porewaters is one of the primary controls on the adsorption of compounds by sediments. The number of negatively charged surface sites decreases with pH as they are filled and neutralized by hydrogen ions. Thus, metal adsorption is generally low at low pH when the ratio of free adsorption sites to metal concentration is low. At higher pH, metals are much more effectively scavenged as more acidic functional groups on organic matter become deprotonated and available for complexation and less mineral surface adsorption sites are filled by hydrogen ions. The pH of sediment porewaters also exerts a control on adsorption of many compounds by its influence on the solubility of minerals, especially iron and manganese oxyhydroxides, which are important adsorption substrates. The partitioning of a compound between the dissolved and solid phase can be described by the ratio of its equilibrium concentrations in the sorbed phase, Cs, and in solution, Cw. This ratio is referred to as a partition coefficient, Kd, where Cs is in mol kg1, and Cw is in mol l1: Kd ¼
Cs Cw
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However, for nutrients and contaminants sorbed on suspended particulates in aquatic systems that partitioning is often defined differently: Kd ¼
mg l – 1 in particular phase mg l – 1 in dissolved phase
Because surface sorption of trace metals is often by surface complexation-like interactions with oxygen donor atoms at the solid surface in the form of Si-O, Fe-O, FeOH, Al-O, Al-OH, or Mn-O, or oxygen or nitrogen atoms in organic matter, the affinity of trace metals for binding sites on solids, and thus their Kd, has been found to often follow the thermodynamic stability of metal complexes described by the Irving-Williams series. The series predicts that, for an oxygen-containing ligand, the general affinity for trace metals, under equivalent conditions, will be Pb > Cu > Ni > Co > Zn > Cd > Fe > Mn > Mg. The relationship between the concentration of a compound in the dissolved phase and the adsorbed phase varies over any range of the compound’s total concentration, and is known as a sorption isotherm. The shape of a sorption isotherm is compound and sorbent dependent. However, experimental data often exhibit behavior similar to mathematically derived isotherms such as the Freundlich or Langmuir isotherms, allowing the sorption behavior of compounds to be simplified and modeled or predicted under certain conditions. Cation Exchange Many different components of sediments are capable of adsorbing cations from solution and releasing equivalent amounts of previous retained cations back into solution by ion exchange. The ability of a given soil or sediment to retain cations can be measured and is referred to as its cation exchange capacity (CEC). Cation exchange occurs via electrostatic interactions between cations in the sediment porewaters and negatively charged sites on sediment particles or organic matter. A cation will be retained by the negatively charged adsorption site in the sediment if it outcompetes and replaces one of the preexisting cations. Since cations of trace metals (e.g., Cd, Cr, Cu, Hg, Pb, and Zn) generally displace hydrogen ions (Hþ) and other major cations (Ca2þ, Mg2þ, Naþ, and Kþ) from exchange sites on sediments, trace metal concentrations are often enriched in sediments relative to their average crustal abundance (i.e., EF > 1). Conversely, increasing ionic strength and concentrations of competing cations in solution decreases the adsorption of metals by neutralizing the negative surface charge and increasing competition for exchange sites. As previously noted, that process occurs in lakes with low acid-neutralizing capacity which are impacted by acid deposition. In those areas, base cations and metals (e.g., Al3þ) are solubilized as the lakes become acidified and
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cation exchange sites are filled with hydrogen ions (Hþ), causing toxic conditions for the aquatic biota. Processes at the Freshwater–Seawater Interface Since estuaries are unique environments, it is not surprising that processes unique to this boundary between freshwater and seawater impact the retention and release of sedimentand particle-bound compounds. One of these processes is the flocculation of DOM, which occurs with the dramatic increase in ionic strength from freshwater (0.002 M) to seawater (0.7 M). The flocculation includes both humic and fulvic acids, which remained dissolved due in part to charged functional groups that interact with polar water molecules in freshwater. However, as DOM transitions from freshwater to seawater the increase in ionic strength and major cation (Ca2þ, Mg2þ, Naþ, and Kþ) concentrations neutralizes the charges responsible for keeping the DOM dissolved and in solution, and thus the organic material flocculates and precipitates out. Colloids, very small particles (0.001–1 mm) that are not truly dissolved but are kept in solution by electrostatic forces and turbulent fluid flow, also flocculate due to the increase in ionic strength in estuarine waters. At low ionic strength these small particles with similar surface charges tend to repel one another, preventing the close physical interaction necessary to form large particles; but at higher ionic strengths these electrostatic forces are destabilized. The freshwater–seawater interface is also generally accompanied by an increase in pH and a reduction in flow velocity, which further contribute to colloidal destabilization, flocculation, and precipitation at this interface. Flocculation and sedimentation of organic matter and colloidal particles are responsible for a substantial decrease not only in the concentration of DOM in the water column, but also in the total water concentration of trace metals and other compounds which are commonly particle bound or complexed by riverine organic matter. This often causes a concurrent increase in sediment organic matter and trace metal levels in estuarine sediments, which may exceed those of freshwater sediments upstream or marine sediments downstream. Consequently, estuarine sediments often act as a trap for nutrients and contaminants, which then tend to be recycled between sediments and overlying waters within estuaries.
Transport Processes Molecular and Ionic Diffusion Diffusion is a continuous process of species migration that tends to decrease concentration gradients both within sediment porewaters and between those waters and overlying waters. Ionic diffusion refers to the diffusion of
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charged species that interact electrostatically, while molecular diffusion is usually used to describe the migration of neutral species. The flux, Fd (mol cm2 s1), produced by the molecular/ionic diffusion in porewaters follows Fick’s law of diffusion, which can be described in a one-dimensional (1D) model: Fd ¼ – Ds
qC qx
where Ds is the effective diffusion coefficient of the solute inside the pores in cm2 s1, the porosity, C the solute concentration in mol l1, and x the position, or distance, in cm. The diffusion of solutes in water can be described with a simplified version of Fick’s law using the molecular diffusion coefficient in free water, Dw, which accounts for temperature and mass effects, with diffusion rates increasing with increasing temperature and decreasing mass of the solute. However, due to the porosity of sediments and the added tortuosity of the path taken by an ion or molecule around sediment grains, to apply Fick’s law to solute diffusion in sediments requires a correction be made to the diffusion coefficient. This is done by the use of Ds, which for sediments with high porosity is commonly derived from Dw via Archie’s law (Ds ¼ Dw). Advection Although diffusion of compounds is important in sediment porewaters and on small spatial scales of less than a meter, rates of diffusion alone do not account for all solute transport in most systems. Advection, the movement of the fluid itself, accounts for much of the solute transport over larger distances and in overlying waters. Specifically, advection represents a flow of sediment or water. For a
solute, the advective flux Fa (mol cm2 s1) in a 1D model can be defined as Fa ¼ uC
with u the speed of water flow (cm s1). Here, advection can be due to burial, compaction, and/or external hydrological flow. The movement of solutes in sediments and porewaters is often primarily via diffusion, but principally via advection in aquatic and atmospheric systems. Both processes, diffusion and advection, are important at the sediment–water or sediment–air interface. Rates of solute transport within sediments by diffusion to the sediment–water interface are largely influenced by the concentration gradient between the sediments and the overlying waters. Rates of diffusion out of sediments are increased when the solutes are then advected away into the overlying waters, as this transport prevents their accumulation at the interface and maintains a large concentration gradient across the sediment–water interface, thus facilitating higher diffusive fluxes than would occur without advection. Advection can also act to replenish solutes (e.g., O2) in overlying waters at this interface and facilitate faster diffusion rates in the reverse direction into sediments. The direction of the concentration gradient of nutrients and contaminants, relative to overlying waters, will determine whether sediments will act as a source or sink for these compounds. A conceptual diagram of transport processes occurring at the sediment–water interface is shown in Figure 2. Bioturbation Bioturbation is the transport of solutes and solids by the activities (e.g., feeding and movement) of macrobenthos.
Colloids Desorption
Flocculation Sorption Disintegration
Particles
Desorption
Water Sorption Sedimentation Resuspension Sediment
Advection
Dissolved in overlying water Advective and diffusive exchange
Water
Desorption
Particles
Sorption
Dissolved in porewater
Advection Sediment
Diffusion Burial
Figure 2 Conceptual diagram of transport and exchange processes near the sediment–water interface.
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Figure 3 Bioturbation of subsurface sediments by benthic organisms (e.g., mollusks, arthropods, and annelids). The structures and activities of those organisms ventilate and extend the depth of oxic conditions in adjacent sediments, illustrated by the light area highlighted by the dashed line, above anoxic sediments, indicated by their darker coloration.
These include arthropods, annelids, and mollusks, which live in biogenic structures buried in the sediment. For example, chironomids dig and ventilate semipermanent U-shaped tubes; tubificid oligochetes live upside down, feeding on bottom sediment; and bivalves burrow into sediments with only their siphon linked to the surface – as illustrated in Figure 3. This biologically mediated flux of molecules and ions Fb (mol cm2 s1) is generally considered as a random phenomenon similar to diffusion. For a solute, the flux is defined as Fb ¼ – DB
qC qx 2 1
with DB the biodiffusion coefficient (cm s ). For a solid, the flux, Fbs, is defined as Fbs ¼ – ð1 – ÞDBs
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interface, a phenomenon that has been the subject of various modeling attempts. This process can be viewed as a nonlocal transport of solutes, with an exchange rate characterized by the irrigation coefficient (s1), that can be estimated from the size and density of the tubes as well as the solute gradients around a tube. In general, decreases to zero with depth, in order to reflect the decreasing density of active fauna. Although bioirrigation is a 3D process, a 1D equation has been used successfully in various models. This equation considers that the speed of intake or input of a solute by bioirrigation at a depth x can be described as q C ¼ ðCtube – C Þ qt x
with Ctube the solute concentration in the tube (mol l1) and t the time (s). In the calculation, Ctube is usually considered to be similar to the concentration measured in water overlying the sediment. Hydrodynamic Dispersion This process corresponds to the mixing of solutes resulting from hydrodynamic flows through permeable sediments and from the movement of waves at the sediment surface. It is usually approximated by a diffusion process function of the magnitude of the porewater velocity and the grain size. The flux is defined as Fd ¼ – Dvd
qC qx
with Dvd the vertical dispersion coefficient (cm2 s1).
qCs qx
with Fbs in g cm2 s1, DBs the bioturbation coefficient of the solid, and Cs the concentration of the solid compound in mol g1. In addition to the slight difference in formulations, it has been found that the effect of bioturbation on solutes can be 20 times more important than on solids.
Bioirrigation Since many organisms live below the oxic layer in surface sediments, they need to irrigate their tubes with oxygenated water from the water column. In addition to advecting oxygen into anoxic sediments, the renewal of water in the sediment structure fulfills other needs of the benthic organisms. These include the transport of food, metabolic wastes, gametes, and environmental stimuli. Bioirrigation of the tubes induces an increase of the fluxes of compounds in solution at the sediment–water
Measuring Nutrient and Contaminant Fluxes in Sediments Porewater Sampling: Squeezing, Peepers, DGT, Microelectrodes Since most nutrients and contaminants contained in sediments are predominantly associated with the solid phase, small changes in their sediment concentrations can translate into relatively large changes in their porewater concentrations. As a consequence, measurements of porewater concentrations are usually the most sensitive indicator of processes occurring in sediments. Similarly, measurements of concentration gradients in porewaters enable calculations of fluxes within sediments and across the sediment–water interface. Consequently, several methods have been developed for sampling sediment porewaters. (1) Interstitial water can be extracted by squeezing or centrifugation of successive slices of a sediment core, after it has been
3180 Ecological Processes | Sediment Retention and Release
collected. (2) Interstitial water may also be collected in situ, using peepers. These are typically plastic devices with a vertical series of small chambers initially filled with high-purity water, which are enclosed within semipermeable membranes. The peepers are inserted into the sediments for an extended period (weeks) to allow the chambers to equilibrate with porewater, after which they are collected and the water in the peepers is extracted. (3) Interstitial water may also be collected in thin gels (DET–DGT) placed between two plates, which are inserted into sediments to allow equilibration of the gels with porewaters before they are collected and sampled. (4) Finally, microelectrodes can be either inserted directly in the sediment or in sediment cores immediately after sampling. Since there are advantages, disadvantages, and limitations for each of these methods, they are often used in combination to obtain a suite of complementary measurements, along with concurrent measurements of particulate sediment concentrations. Benthic Flux Chamber Measurements An alternative to sampling porewaters to evaluate fluxes of nutrients and contaminants at the sediment–water interface is the use of a benthic flux chamber. This type of device is essentially a container enclosing sediment and a small volume of overlying water that is incubated for periods lasting from a few hours to several weeks. During this time, concentrations of nutrients and contaminants in the overlying water are monitored at regular intervals, and changes indicate the direction and proportion of their fluxes; an increase in concentrations indicates a flux coming out of the sediment, while a decrease shows a flux into the sediment. Although flux chambers provide a direct measure of nutrient and contaminant fluxes, they are not without their shortcomings. Two of the more salient being that (1) the enclosed water will over time become anoxic with contact with the sediment, and (2) the enclosed water is secluded from surrounding benthic currents that mix water
and sediment. Unless artificially regulated, both of these phenomena will induce anomalous changes in the redox chemistry that will influence the measured fluxes. Sequential Extraction Methods While total concentrations of nutrients and contaminants in sediments are measured to quantify spatial and temporal gradients, concentrations of specific fractions on the surface and within sediments may be further characterized using selective extraction techniques. These extraction techniques provide more detailed information on how those constituents are bound in sediments, as well as their bioavailability and conditions for their remobilization from the sediments. As noted at the beginning of this article, bioavailability and mobility are especially important in studies of nutrients and contaminants in sediments because they are dynamic systems which are subject to chemical and physical changes on the timescales of seconds, days, seasons, years, decades, and longer. To address these concerns, various sequential extraction techniques have been developed to determine semiquantitatively how elements and compounds, including nutrients and contaminants, are associated with sediments. The extractions are designed to sequentially extract from the most weakly bound (e.g., ion exchangeable) to the most strongly bound (e.g., refractory crystal lattice) fractions. While the fractions are operationally defined, rather than definitive and truly specific, they have proved to be relatively useful in characterizing their relative bioavailability and potential for diagenic remobilization from sediments, as well as their biogeochemical cycling between sediments and water. Table 1 provides an example of a sequential extraction for sediments. It shows the extractants and conditions used in five sequential steps and the operationally defined fraction solubilized in each of those steps. There are numerous variations of these techniques (i.e., number of steps, extractants and conditions, and operationally defined fractions) with similar applicabilities and
Table 1 Sequential extraction steps and operationally defined fractions used to characterize the phase distribution of constituents, including nutrients and contaminants, in sediments Step no.
Extractants and conditions
Operationally defined fraction
1 2 3 4
1 M NaOAc, pH 8.2, 25 C 1 M NaOAc, pH 5 (HOAc), 25 C 0.04 M NH2OH-HCl in 25% (v/v) HOAc, 100 C 0.02 M HNO3 þ 30% H2O2, pH 2, 85 C; then 3.2 M NH4OAc in 20% (v/v) HNO3 Concentrated HNO3 þ HF, 110 C
Exchangeable Bound to carbonates Reducible, bound to Fe–Mn oxides Oxidizable, strongly bound to organics and sulfide Residual, recalcitrant
5
Modified from Tessier A, Campbell P G C, and Bisson M (1979) Sequential extraction procedure for the speciation of particulate trace-metals. Analytical Chemistry 51: 844–851.
Ecotoxicology | Sediments: Setting, Transport, Mineralization, and Modeling
limitations. Those limitations, along with previously noted limitations in measuring the speciation and fluxes of nutrients and contaminants in sediments, attest to their truly complex biogeochemical cycling and bioavailability. See also: Adsorption; Microbial Ecological Processes: Aerobic/Anaerobic; Physical Transport Processes in Ecology: Advection, Diffusion, and Dispersion; Transport in Porous Media.
Further Reading Allen HE (ed.) (1995) Metal Contaminated Aquatic Sediments. Chelsea, MI: Ann Arbor Press. Allen HE, Huang CP, Bailey GW, and Bowers AR (eds.) (1995) Metal Speciation and Contamination of Soils. Boca Raton, FL: Lewis Publishers. Baudo R, Giesy J, and Mantau H (eds.) (1990) Sediments: Chemistry and Toxicity of In-Place Pollutants. Ann Arbor, MI: Lewis Publishers. Berg P, Risgaard-Petersen N, and Rysgaard S (1998) Interpretation of measured concentration profiles in sediment pore water. Limnology and Oceanography 43: 1500–1510. Berg P, Rysgaard S, Funch P, and Sejr MK (2001) Effects of bioturbation on solutes and solids in marine sediments. Aquatic Microbiology and Ecology 26: 81–94. Berg P, Rysgaard S, and Thamdrup B (2003) Dynamic modelling of early diagenesis and nutrient cycling. A case study in an arctic marine sediment. American Journal of Science 303: 905–955. Berner RA (1971) Principles of Chemical Sedimentology. New York: McGraw-Hill.
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Berner RA (1980) Early Diagenesis. Princeton, NJ: Princeton University Press. Boudreau BP (1997) Diagenetic Models and their Implementation: Modelling Transport and Reactions in Aquatic Seduments. Berlin: Springer. Committee on Bioavailability of Contaminants in Soils and Sediments, Water Science and Technology Board, Division on Earth and Life Studies, National Research Council (2003) Bioavailability of Contaminants in Soils and Sediments: Processes, Tools, and Applications. Washington, DC: National Academies Press. DePinto JV, Lick W, and Paul JF (eds.) (1994) Transport and Transformation of Contaminants Near the Sediment–Water Interface. Boca Raton, FL: Lewis Publishers. Morel FMM and Hering JG (1993) Principles and Applications of Aquatic Chemistry. New York, NY: Wiley-Interscience. Schwarzenbach RP, Gschwend PM, and Imbode DM (eds.) (2002) Environmental Organic Chemistry, 2nd edn. New York: WileyInterscience. Song Y and Mu¨ller G (1999) Sediment–Water Interactions in Anoxic Freshwater Sediments: Mobility of Heavy Metals and Nutrients. Berlin: Springer. Stumm W and Morgan JJ (eds.) (1996) Aquatic Chemistry, Chemical Equilibria and Rates in Natural Waters, 3rd edn. New York: Wiley. Tessier A, Campbell P G C, and Bisson M (1979) Sequential extraction procedure for the speciation of particulate trace-metals. Analytical Chemistry 51: 844–851. Van Cappellen P and Gaillard J-F (1996) Biogeochemical dynamics in aquatic sediments. In: Lichtner PC, Steefel CI, and Oelkers EH (eds.) Reviews in Mineralogy, Vol. 34: Reactive Transport in Porous Media, pp. 336–376. Van Cappellen P and Wang Y (1996) Cycling of iron and manganese in surface sediments: A general theory for the coupled transport and reaction of carbon, oxygen, nitrogen, sulfur, iron, and manganese. American Journal of Science 296: 197–243.
Sediments: Setting, Transport, Mineralization, and Modeling L Kamp-Nielsen, University of Copenhagen, Hillerød, Denmark ª 2008 Elsevier B.V. All rights reserved.
Net Deposition as an Aggregated Approach Theoretical Models
Further Reading
Sediments are particulate matter that can be or have been transported by fluid, wind, and glaciers and which might have been deposited as a layer of solid particles in dense suspension at the bottom of water bodies. The parts of sediments which have their origin outside the water bodies are called allochthonous sediments and have been transported by runoff from the drainage basin of the water body or by wet or dry deposition on the surface of the water body. Particles formed within the water body and the sediment are the autochthonous parts of the sediments and are transformations of dissolved elements to particles by chemical and biological processes. In the photic zone of rivers, lakes, coastal waters, and oceans, inorganic carbon as carbon dioxide or bicarbonate is fixed as particulate organic carbon by photosynthetic
organisms, or fixed as inorganic carbonate in corals, foramifers, and coccolithophorids. The dead and living organic particles can be processed through the aquatic food chain and sink to the bottom as living or dead particulate organic material. Due to increased pH, as a result of photosynthesis, the solubility product of calcium and magnesium carbonate can be exceeded and precipitation of carbonates may occur. Both the allochthonous and the autochthonous parts of the sediments are subject to further biogenetical processes in the sediment environment and other autogenic fractions are generated. On a geological timescale, surface sediments are young structures, but they play an important role in the global and local cycling of elements like carbon, nutrients, and metals – all of which are important for the productivity in
Sediment Retention and Release F J Black, C Gallon, and A R Flegal, University of California Santa Cruz, Santa Cruz, CA, USA ª 2008 Elsevier B.V. All rights reserved.
Introduction Chemical Reactions Transport Processes
Measuring Nutrient and Contaminant Fluxes in Sediments Further Reading
Introduction
Chemical Reactions
Both the retention and release of nutrients and contaminants by sediments are controlled by complex physical and biogeochemical interactions. Sediments, porewaters, and overlying waters are dynamic matrices that alternatively serve as sources and sinks for elements and compounds, as physical and biogeochemical conditions vary. Some of those conditions markedly change over small spatial scales, from nanometers to millimeters, and over time, from nanoseconds to millennia. Consequently, rates of scavenging and mobilization of elements and compounds to and from sediments, respectively, may also exhibit pronounced spatial and temporal variability. This summary, therefore, briefly describes some of the principal factors involved in the uptake and release of nutrients and contaminants from sediments, as well as the diffusive and advective dispersion of those constituents in sediment porewaters and overlying waters. This includes a synopsis of chemical reactions that affect the speciation, cycling, and bioavailability of constituents in sediments that can result from the degradation of organic matter, other redox transformations, the dissolution and precipitation of mineral phases, and sorption processes, including cation exchange. It also includes brief comments on transport processes related to geochemical and physical characteristics of the media, as well as the influence of benthic organisms on the distribution and dispersion of those constituents in sediments and associated waters. The complex integration of all the preceding factors determines the physical and biogeochemical cycling of nutrients and contaminants in sediments. Directions and rates of those cycles are governed by prevailing chemical and physical conditions, and are often mediated or driven by microbial communities. While those constituents may initially be entrained in sediments by geochemical sorption, authigenic precipitation, and biological scavenging, they may subsequently be remobilized from the sediments into dissolved and colloidal phases by mineral weathering, desorption, and decomposition.
Organic Matter Decomposition When suspended particulate, colloidal, and dissolved organic matter (DOM) are scavenged from surface waters and deposited to benthic sediments they encounter the sediment–water interface, an important region in the biogeochemical cycling of many elements and compounds. Most biologically mediated oxidation and mineralization of organic matter, referred to as early diagenesis, occurs within that interface, which typically extends a few centimeters into the surface sediments. This early diagenesis is carried out by microbes using the most energetically favorable oxidant available, which in aerobic environments is molecular oxygen, O2. This is illustrated by the simplified oxidation of organic matter, based on Redfield’s ratio of C:N:P of 106:16:1: ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 138 O2 ! 106 CO2 þ 16 HNO3 þ H3 PO4 þ 122 H2 O
The G of the above reaction is 3190 kJ mol1, which is thermodynamically favorable and yields substantial free energy. The oxidation of organic material via this reaction will be dominated by aerobic bacteria as long as sufficient O2 is available. However, the rate of O2 consumption will exceed the rate of O2 diffusion from overlying waters to sediments at some sediment depth. In highly productive aquatic systems and those impacted by relatively large discharges of organic waste, the depletion of O2 may occur at or above the sediment–water interface. Microbial depletion of O2 leads to the formation of suboxic conditions in deeper sediments, where further oxidation of organic material continues with microbes using the next most energetically favorable terminal electron acceptor. As one oxidant is consumed, the next most energetically favorable oxidant is utilized. MnO2 and NO 3 reduction occur following the loss of O2, sequentially followed by iron reduction, sulfate reduction, and methanogenesis. This sequence is illustrated by the following balanced equations for the oxidation of organic material and the Gibbs
Ecological Processes | Sediment Retention and Release
free energy associated with each terminal electron acceptor: Manganese oxide reduction: G ¼ 3090 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 236 MnO2 þ 472 Hþ ! 236 Mn2þ þ 106 CO2 þ 8 N2 þ H3 PO4 þ 366 H2 O
Nitrate reduction: G ¼ 3030 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 94:4 HNO3 ! 106 CO2 þ 55:2 N2 þ H3 PO4 þ 177:2 H2 O
Iron reduction: G ¼ 1330 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 424 FeOOH þ 848 Hþ ! 424 Fe2þ þ 106 CO2 þ 16 NH3 þ H3 PO4 þ 742 H2 O
Sulfate reduction: G ¼ 380 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ þ 53 SO24 – ! 53 S2 – þ 106 CO2 þ 16 NH3 þ H3 PO4 þ 106 H2 O
Methanogenesis: G ¼ 350 kJ mol1 ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þ ! 53 CH4 þ 53 CO2 þ 16 NH3 þ H3 PO4
The distribution of redox zones with depth below the sediment–water interface where each of the above respiration pathways occurs varies with physical and biogeochemical conditions. In naturally eutrophic and contaminated sediments, where inputs of organic matter are relatively high, O2 penetration by diffusion from overlying waters may be limited to a few millimeters. In oligotrophic freshwater and deep-sea sediments, where organic matter inputs are much lower, O2 may diffuse a few centimeters below the sediment–water interface. However, the advection of surface waters into the sediments due to bioturbation commonly accounts for greater and uneven penetration of O2 and other terminal electron acceptors in benthic sediments. Consequently, the distribution of redox zones in sediments may be highly heterogeneous. Multiple microenvironments often exist within millimeters of each other where aerobic respiration, nitrate reduction, and sulfate reduction are carried out simultaneously at comparable depths. Since the rate of organic matter oxidation is temperature dependent due to its effect on microbial respiration rates, early diagenesis in shallow water sediments also displays seasonal variability. The oxidation of organic material in sediments during early diagenesis causes a number of chemical changes beyond those in the organic matter itself. The decomposition of the organic matter also releases other nutrients, trace elements, and contaminants which have been complexed or associated with it, when it is solubilized. All of the preceding forms of respiration also produce carbon dioxide (CO2). This increase in CO2 levels is accompanied by a decrease in pH due to the formation
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of carbonic acid (H2CO3) and bicarbonate (HCO 3 ). That decrease can, in turn, lead to the mobilization of cations adsorbed to the sediments as an increasing number of hydrogen ions compete for negatively charged adsorption sites, through cation exchange. Phosphorous, principally in the form of phosphate, and nitrogen, principally in the form of nitrate or ammonia, are liberated directly by the mineralization of organic matter, are liberated directly by the mineralization of organic matter, as well as indirectly by the reduction of iron and manganese oxyhydroxides. As a result, those two macronutrients (N and P) are often depleted in surface waters where they are commonly the limiting nutrients for primary productivity and are enriched in subsurface and porewaters, where they are released with the decomposition of organic matter. This nutrient-type profile is also paralleled by those of other trace- and micronutrients, and primary productivity is often limited by the subsequent flux of those remobilized nutrients to the euphotic zone in overlying waters. Flux rates out of anoxic lacustrine and estuarine sediments reported in the literature are generally of the order of 0–20 mg m 2 day1 for phosphate, 0–10 mg m 2 day1 for nitrate, and 0–30 mg m 2 day1 for ammonium. However, these flux rates vary substantially by environmental setting, as well as spatially and temporally. In addition, these nutrient fluxes may be reversed, from the overlying waters to sediments, under some conditions. Other Redox Reactions The redox state of sediments and associated porewaters influences the retention and release of trace elements and other compounds, as illustrated by chromium. It commonly exists in the environment in two oxidation states: trivalent chromium, Cr(III), which is an essential trace element, and hexavalent chromium, Cr(VI), which is carcinogenic. The latter species is also much more labile in most aquatic systems, and may be found at potentially toxic levels (>100 mg l 1) in some aquifers due to either natural processes or industrial pollution. While Cr(III) is more thermodynamically stable under reducing conditions, it is generally only found at low levels because it exists in solution primarily as cations (Cr3þ, Cr(OH)þ 2, Cr(OH)2þ) that have high affinities for sediments with a net negative surface charge (e.g., iron oxides and clay minerals) and Cr(III) readily precipitates as Cr(OH)3 and Cr2O3 in the pH ranges commonly encountered in the environment. In contrast, Cr(VI) forms relatively soluble oxyanions (e.g., CrO2 4 and HCrO4 ) that have less affinity for sediment surfaces and, therefore, may occur at higher concentrations in many aquifers. Although Cr(VI) may be naturally converted to Cr(III) in reducing groundwater systems, the microbial oxidation of Cr(III) to Cr(VI) at the aerobic/anaerobic interface in sediments has been
3174 Ecological Processes | Sediment Retention and Release
associated with anomalously high levels of the potentially toxic species in some aquifers – as has the discharge of industrial Cr(VI) in other aquifers. Sediment redox processes also strongly affect metalsulfide minerals, which can be solubilized under oxidizing conditions. The oxidation of metal-sulfides such as cinnabar (HgS), cuprite (CuS), galena (PbS), sphalerite (ZnS), chalcopyrite (CuFeS2), and chalcocite (Cu2S) by oxygen is similar to that of pyrite (FeS2): 4 FeS2 þ 15 O2 þ 14 H2 O ! 4 FeðOHÞ3 þ 8 H2 SO4
The oxidation of the reduced Fe(II) and sulfide in pyrite results in the formation of Fe(III) in the form of iron oxide (Fe(OH)3) and sulfate in the form of sulfuric acid (H2SO4), respectively. The H2SO4 produced by that oxidation is responsible for much of the acid mine drainage (pH < 5) generated from mines and mine tailings. While the abiotic oxidation of sulfidic minerals is thermodynamically favorable in aerated waters, the kinetics are generally very slow at ambient environmental conditions because of the large activation energy required. However, the oxidation of sulfidic minerals is markedly accelerated by bacteria (e.g., Thiobacillus ferrooxidans) which obtain energy from the oxidation of pyrite. Acid mine drainage of sulfidic deposits is accompanied by the leaching and mobilization of other heavy metals. These may include arsenic, cadmium, copper, lead, manganese, mercury, selenium, and zinc, which are all relatively toxic to most aquatic organisms. Consequently, surface waters downstream from acid mine drainage may be nearly devoid of all but microbial life until pH levels increase and iron and manganese precipitate out as oxyhydroxides that scavenge those toxic elements. However, sediment burial following the deposition of iron and manganese oxyhydroxides leads to reducing conditions in subsurface sediments, which thermodynamically favor the reductive dissolution of those compounds. This process releases not only iron and manganese back into solution, but also metals and metalloids that had been adsorbed onto the oxyhydroxides. As a result, sediments with relatively high levels of contaminant metals that have been transported downstream from acid mine drainage constitute a potential source of contamination to porewaters and overlying waters. Similarly, other sediments that have relatively large amounts of metals (e.g., Cd, Cu, Pb) and metalloids (e.g., As, Se) – from natural or industrial sources – scavenged onto iron and manganese oxyhydroxides represent a potential source of pollution under reducing conditions. Again, those reducing conditions may be catalyzed by the deposition of organic matter, either following an algal bloom or the discharge of organic industrial, agricultural, and municipal wastes.
Heavy metals are not the only class of compounds which are readily adsorbed by iron and manganese oxyhydroxides. Nutrients such as phosphate and nitrate also exhibit this same behavior under oxidizing conditions, and are thus released when the sediments become reducing and the Fe(III) and Mn(IV) precipitates are solubilized to dissolved forms of Fe(II) and Mn(II). Thus, a change in the redox state of sediments to more reducing conditions can result in the release of phosphate and nitrate to overlying waters where they can promote eutrophication. As a result, lakes which receive nutrient-rich runoff from urban or agricultural areas are often mechanically aerated in order to maintain oxygenated bottom waters and surface sediments to prevent the reductive dissolution of iron and manganese oxyhydroxides and the subsequent release of nutrients. This method has also been employed successfully to promote the reformation of iron and manganese oxyhydroxides and recovery of lakes following eutrophication. Mineral Formation and Weathering The formation and weathering of the mineral matrix of sediments plays a primary role in the retention and release of ions and compounds. The formation of sediments by authigenic precipitation removes ions from solution and creates fresh surfaces for the adsorption of other inorganic and organic compounds as a secondary retention mechanism. The weathering of sediments, conversely, releases both adsorbed components and parent material into solution. The overall chemical processes of mineral precipitation and dissolution can be generally described in terms of nucleation, crystal growth, and weathering. Nucleation is often the rate-limiting step of mineral precipitation and can occur by homogeneous or heterogeneous mechanisms. During homogeneous precipitation, crystal nuclei are formed in a saturated solution by the random collision of ions. Heterogeneous nucleation involves the formation of mineral nuclei on the surfaces of reactive solids already present. Regardless of the nucleation pathway, once stable nuclei are formed, mineral growth proceeds spontaneously until the solution is no longer saturated with respect to that mineral. The rate of nucleation is controlled by the degree of supersaturation, temperature, and the geometry of initial nuclei formed or heterogeneous materials available for nuclei seeding. Nucleation rates are also dependent upon the specific interfacial free energy, which is the difference in free energy between an ion bound within the mineral matrix and an equivalent ion bound to the mineral surface. Crystal growth involves transport of reacting ions to the mineral surface, transport of reaction products away from the mineral surface, and surface interactions
Ecological Processes | Sediment Retention and Release
including adsorption, surface complexation, dehydration, and cation exchange. Mineral growth can be transport controlled, where growth is limited by the rate at which ions or complexes migrate to the surface via diffusion and advection, or surface controlled, where the rate-limiting step is the surface interaction involved. Most mineral formation reactions are surface-controlled processes. The nature of sediment weathering products and their subsequent release depends upon the mineralogical composition of the parent material, the chemical composition of the aqueous phase, and the nature of fluid flow. Chemical weathering reactions are often classified by the attacking agent type and the manner in which the mineral is altered. Dissolution and hydrolysis reactions are mediated by potential attacking substances including acids, oxygen-containing ligands, and water. Since many weathering reactions can be treated as acid–base reactions, the pH of sediment porewaters exerts a primary control on sediment dissolution and weathering rates. The source of acidity in sediments can be organic acids exuded by plant roots or from the decay of organic matter; sulfuric acid from sulfide oxidation; nitric, sulfuric, and hydrochloric acids in acid rain; and carbonic acid, with CO2 derived from the atmosphere or the respiration of organic matter. If simple dissolution is the primary mechanism of weathering, the process is termed congruent dissolution, as illustrated by the dissolution of quartz (SiO2(qtz)): SiO2ðqtzÞ þ 2 H2 O $ H4 SiO4ðaqÞ
This reaction is reversible, and dissolved silica in the form of silicic acid, H4SiO4(aq), can precipitate to form quartz. But the kinetics of quartz precipitation are quite slow below 70 C, and as a result silicic acid concentrations are often measured in excess of its predicted solubility of about 180 mM in low-temperature soil and sediment porewaters. Therefore, while thermodynamics predict the precipitation of quartz from many soil and sediment porewaters, the slow kinetics involved often result in the supersaturation of dissolved silica. The dissolution of calcite, CaCO3(calc), is another congruent weathering reaction: CaCO3ðcalcÞ þ H2 CO3 ! Ca2þ þ 2 HCO3–
In this process, calcium carbonate and carbonic acid react to form calcium cations, Ca2þ, and bicarbonate anions, HCO3. This reaction yields no solid phase, because the carbonate mineral undergoes complete dissolution. The equilibrium of this reaction, as well as that of the preceding reaction, is pH dependent, with higher acidity resulting in greater dissolution of the product ions into solution. If a secondary mineral forms during a chemical weathering reaction the process is referred to as incongruent dissolution. A common form this type of process takes
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is the weathering of aluminosilicates minerals, which are stable at high temperatures and pressures characteristic of the Earth’s interior, to clay minerals, which are more stable at the low temperatures and pressures characteristic of the Earth’s surface. An example of incongruent dissolution is the hydrolysis of potassium feldspar, KAlSi3O8(Kspar), to kaolinite, Al2Si2O5(OH)4(kaol): 2 KAlSi3 O8ðKsparÞ þ 2Hþ þ 9 H2 O ! Al2 Si2 O5 ðOHÞ4ðkaolÞ þ 4 H4 SiO4ðaqÞ þ 2 Kþ
The incongruent dissolution of potassium feldspar not only forms the solid-phase kaolinite, but also releases silicic acid, H4SiO4(aq), and potassium cations, Kþ, into solution. The stability of the different mineral phases is, again, dependent upon pH and fluid chemistry, with increased acidity and decreased silicic acid and potassium cation concentrations promoting feldspar weathering. Determination of which mineral phase is most stable under prevailing sediment conditions can be made with thermodynamic models and stability (Eh:pH) diagrams. Weathering reactions generally neutralize acids and release base cations (Ca2þ, Mg2þ, Naþ, Kþ) into solution. Weathering of aluminosilicates releases dissolved silica, but not comparable levels of aluminum due to its relatively low solubility around neutral pH (7). But in freshwater lakes with sediments and soils low in carbonates and low acid-neutralizing capacities that are impacted by acid rain, the pH of porewaters eventually drops to levels where the dissolution of aluminum (Al3þ) is sufficient to cause asphyxiation in fish and invertebrates. Nevertheless, aluminum is often treated as a conservative element and its concentration is compared to that of other elements in sediments to estimate levels of contamination. These qualitative estimates are based on the assumption that the concentration of aluminum in contaminated sediments is not enriched relative to its average crustal abundance ( 8.1 mg g 1), while concentrations of other elements may be enriched in contaminated sediments relative to their average crustal abundance. The resulting enrichment factor (EF) is then derived with a simple normalization of those ratios: EF ¼
½X sediment =½A1sediment ½X crustal abundance =½A1crustal abundance
where[X] is the concentration of the element of interest and [A1] is the concentration of aluminum. The presence of ligands and organic matter in sediment porewaters can increase mineral weathering by complexing cations involved in the weathering reaction. This decreases the cation’s free ion concentration and causes a thermodynamic shift which favors further mineral dissolution, as predicted by Le Chatelier’s principle. Complexation by organic matter also increases the solubility of metals via this mechanism, which acts to
3176 Ecological Processes | Sediment Retention and Release
facilitate the dissolution of minerals containing metals exhibiting low solubilities at near neutral pHs, such as Fe3þ, Al3þ, and Hg2þ. Finally, microbes commonly play an important role in the rates of both mineral precipitation and weathering, and thus the sequestering or release of nutrients and contaminants in sediments. Microbes often mediate the underlying chemical reactions involved or take advantage of the energy released by the associated reactions. For example, microbes can facilitate nucleation by their cell membranes when they act as nucleation sites or when they produce an organic molecule that serves as a template for an inorganic crystal, a phenomenon known as biomineralization. Similarly, microbes often effectively control the rate of sulfide mineral oxidation, a thermodynamically favorable but kinetically slow inorganic reaction, by their production of enzymes which catalyze the reaction as a means of harnessing the energy released. Adsorption Adsorption is the result of chemical bonding between a gas, nonaqueous liquid, or dissolved compound and the solid sediment mineral matrix or organic matter. The strength and reversibility of this interaction can vary substantially depending upon the nature of the bonding. The most important substrates for adsorption in sediments are generally clay minerals, organic matter, and iron and manganese oxyhydroxides. Organic compounds with hydrophobic regions often undergo adsorption onto sediment organic matter or mineral faces by Van der Waals or hydrophobic interactions. Adsorption of metals and other ions is generally via ionic or electrostatic interactions, with bonding sometimes being more covalent in nature. Metals that are complexed by DOM will also be retained by sediments if the organic matter with which they are associated is adsorbed onto sediment particles. The adsorptive behavior of sediment is controlled by its organic content, surface area, and thus particle size, pH, and the type and density of adsorption sites. Sediments with more organic coatings generally absorb relatively large amounts of organic and hydrophobic compounds. Fine-grained sediments with relatively small particle sizes and large total surface areas have a relatively high number of adsorption sites, and thus relatively greater adsorption capacity than that of largergrained sediments. Since clay particles exhibit both a high surface area as well as a net negative charge density, they are generally more effective at adsorbing dissolved metals and other cations than coarser-grained sediments. The mineralogy of sediments also influences the density and type of adsorption and exchange sites present, as does their moisture content, which can play an important role in adsorption, with increasing water content being associated with decreasing retention of nonpolar compounds.
c
d
b e a
f h
g
Figure 1 Stylized representation of a clay particle with surface coatings: (a) iron oxyhydroxides, (b) manganese oxyhydroxides, (c) aluminosilicates, (d) other inorganics (e.g., calcite, apatite), (e) organic matter, (f) microorganisms (e.g., bacteria), (g) cross section of the clay particle showing net negative charge on its surface, (h) clay surface with a net negative charge resulting in the subsequent adsorption of cations and other coatings as detailed above. Although each of these coatings is shown as being discrete, they generally overlap and overlay each other.
A stylized sequencing of some of those processes is illustrated in Figure 1. Chemical adsorption of cations can also be treated as a surface complexation reaction in which lone electron pairs of primarily oxygen, nitrogen, and sulfur atoms at the solid surface are donated to metals and other cations to form surface complexes. In this model, surface complexing sites compete with dissolved complexing agents for cations, both being capable of forming inner or outer sphere complexes. As previously noted, surface hydroxyl groups of iron and manganese oxyhydroxide solids exhibit strong affinities for many trace metals that are scavenged by sediments. But, the presence of dissolved ligands capable of outcompeting sediment surface complexing sites can lead to metal desorption and mineral dissolution. The pH of sediment porewaters is one of the primary controls on the adsorption of compounds by sediments. The number of negatively charged surface sites decreases with pH as they are filled and neutralized by hydrogen ions. Thus, metal adsorption is generally low at low pH when the ratio of free adsorption sites to metal concentration is low. At higher pH, metals are much more effectively scavenged as more acidic functional groups on organic matter become deprotonated and available for complexation and less mineral surface adsorption sites are filled by hydrogen ions. The pH of sediment porewaters also exerts a control on adsorption of many compounds by its influence on the solubility of minerals, especially iron and manganese oxyhydroxides, which are important adsorption substrates. The partitioning of a compound between the dissolved and solid phase can be described by the ratio of its equilibrium concentrations in the sorbed phase, Cs, and in solution, Cw. This ratio is referred to as a partition coefficient, Kd, where Cs is in mol kg1, and Cw is in mol l1: Kd ¼
Cs Cw
Ecological Processes | Sediment Retention and Release
However, for nutrients and contaminants sorbed on suspended particulates in aquatic systems that partitioning is often defined differently: Kd ¼
mg l – 1 in particular phase mg l – 1 in dissolved phase
Because surface sorption of trace metals is often by surface complexation-like interactions with oxygen donor atoms at the solid surface in the form of Si-O, Fe-O, FeOH, Al-O, Al-OH, or Mn-O, or oxygen or nitrogen atoms in organic matter, the affinity of trace metals for binding sites on solids, and thus their Kd, has been found to often follow the thermodynamic stability of metal complexes described by the Irving-Williams series. The series predicts that, for an oxygen-containing ligand, the general affinity for trace metals, under equivalent conditions, will be Pb > Cu > Ni > Co > Zn > Cd > Fe > Mn > Mg. The relationship between the concentration of a compound in the dissolved phase and the adsorbed phase varies over any range of the compound’s total concentration, and is known as a sorption isotherm. The shape of a sorption isotherm is compound and sorbent dependent. However, experimental data often exhibit behavior similar to mathematically derived isotherms such as the Freundlich or Langmuir isotherms, allowing the sorption behavior of compounds to be simplified and modeled or predicted under certain conditions. Cation Exchange Many different components of sediments are capable of adsorbing cations from solution and releasing equivalent amounts of previous retained cations back into solution by ion exchange. The ability of a given soil or sediment to retain cations can be measured and is referred to as its cation exchange capacity (CEC). Cation exchange occurs via electrostatic interactions between cations in the sediment porewaters and negatively charged sites on sediment particles or organic matter. A cation will be retained by the negatively charged adsorption site in the sediment if it outcompetes and replaces one of the preexisting cations. Since cations of trace metals (e.g., Cd, Cr, Cu, Hg, Pb, and Zn) generally displace hydrogen ions (Hþ) and other major cations (Ca2þ, Mg2þ, Naþ, and Kþ) from exchange sites on sediments, trace metal concentrations are often enriched in sediments relative to their average crustal abundance (i.e., EF > 1). Conversely, increasing ionic strength and concentrations of competing cations in solution decreases the adsorption of metals by neutralizing the negative surface charge and increasing competition for exchange sites. As previously noted, that process occurs in lakes with low acid-neutralizing capacity which are impacted by acid deposition. In those areas, base cations and metals (e.g., Al3þ) are solubilized as the lakes become acidified and
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cation exchange sites are filled with hydrogen ions (Hþ), causing toxic conditions for the aquatic biota. Processes at the Freshwater–Seawater Interface Since estuaries are unique environments, it is not surprising that processes unique to this boundary between freshwater and seawater impact the retention and release of sedimentand particle-bound compounds. One of these processes is the flocculation of DOM, which occurs with the dramatic increase in ionic strength from freshwater (0.002 M) to seawater (0.7 M). The flocculation includes both humic and fulvic acids, which remained dissolved due in part to charged functional groups that interact with polar water molecules in freshwater. However, as DOM transitions from freshwater to seawater the increase in ionic strength and major cation (Ca2þ, Mg2þ, Naþ, and Kþ) concentrations neutralizes the charges responsible for keeping the DOM dissolved and in solution, and thus the organic material flocculates and precipitates out. Colloids, very small particles (0.001–1 mm) that are not truly dissolved but are kept in solution by electrostatic forces and turbulent fluid flow, also flocculate due to the increase in ionic strength in estuarine waters. At low ionic strength these small particles with similar surface charges tend to repel one another, preventing the close physical interaction necessary to form large particles; but at higher ionic strengths these electrostatic forces are destabilized. The freshwater–seawater interface is also generally accompanied by an increase in pH and a reduction in flow velocity, which further contribute to colloidal destabilization, flocculation, and precipitation at this interface. Flocculation and sedimentation of organic matter and colloidal particles are responsible for a substantial decrease not only in the concentration of DOM in the water column, but also in the total water concentration of trace metals and other compounds which are commonly particle bound or complexed by riverine organic matter. This often causes a concurrent increase in sediment organic matter and trace metal levels in estuarine sediments, which may exceed those of freshwater sediments upstream or marine sediments downstream. Consequently, estuarine sediments often act as a trap for nutrients and contaminants, which then tend to be recycled between sediments and overlying waters within estuaries.
Transport Processes Molecular and Ionic Diffusion Diffusion is a continuous process of species migration that tends to decrease concentration gradients both within sediment porewaters and between those waters and overlying waters. Ionic diffusion refers to the diffusion of
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charged species that interact electrostatically, while molecular diffusion is usually used to describe the migration of neutral species. The flux, Fd (mol cm2 s1), produced by the molecular/ionic diffusion in porewaters follows Fick’s law of diffusion, which can be described in a one-dimensional (1D) model: Fd ¼ – Ds
qC qx
where Ds is the effective diffusion coefficient of the solute inside the pores in cm2 s1, the porosity, C the solute concentration in mol l1, and x the position, or distance, in cm. The diffusion of solutes in water can be described with a simplified version of Fick’s law using the molecular diffusion coefficient in free water, Dw, which accounts for temperature and mass effects, with diffusion rates increasing with increasing temperature and decreasing mass of the solute. However, due to the porosity of sediments and the added tortuosity of the path taken by an ion or molecule around sediment grains, to apply Fick’s law to solute diffusion in sediments requires a correction be made to the diffusion coefficient. This is done by the use of Ds, which for sediments with high porosity is commonly derived from Dw via Archie’s law (Ds ¼ Dw). Advection Although diffusion of compounds is important in sediment porewaters and on small spatial scales of less than a meter, rates of diffusion alone do not account for all solute transport in most systems. Advection, the movement of the fluid itself, accounts for much of the solute transport over larger distances and in overlying waters. Specifically, advection represents a flow of sediment or water. For a
solute, the advective flux Fa (mol cm2 s1) in a 1D model can be defined as Fa ¼ uC
with u the speed of water flow (cm s1). Here, advection can be due to burial, compaction, and/or external hydrological flow. The movement of solutes in sediments and porewaters is often primarily via diffusion, but principally via advection in aquatic and atmospheric systems. Both processes, diffusion and advection, are important at the sediment–water or sediment–air interface. Rates of solute transport within sediments by diffusion to the sediment–water interface are largely influenced by the concentration gradient between the sediments and the overlying waters. Rates of diffusion out of sediments are increased when the solutes are then advected away into the overlying waters, as this transport prevents their accumulation at the interface and maintains a large concentration gradient across the sediment–water interface, thus facilitating higher diffusive fluxes than would occur without advection. Advection can also act to replenish solutes (e.g., O2) in overlying waters at this interface and facilitate faster diffusion rates in the reverse direction into sediments. The direction of the concentration gradient of nutrients and contaminants, relative to overlying waters, will determine whether sediments will act as a source or sink for these compounds. A conceptual diagram of transport processes occurring at the sediment–water interface is shown in Figure 2. Bioturbation Bioturbation is the transport of solutes and solids by the activities (e.g., feeding and movement) of macrobenthos.
Colloids Desorption
Flocculation Sorption Disintegration
Particles
Desorption
Water Sorption Sedimentation Resuspension Sediment
Advection
Dissolved in overlying water Advective and diffusive exchange
Water
Desorption
Particles
Sorption
Dissolved in porewater
Advection Sediment
Diffusion Burial
Figure 2 Conceptual diagram of transport and exchange processes near the sediment–water interface.
Ecological Processes | Sediment Retention and Release
Figure 3 Bioturbation of subsurface sediments by benthic organisms (e.g., mollusks, arthropods, and annelids). The structures and activities of those organisms ventilate and extend the depth of oxic conditions in adjacent sediments, illustrated by the light area highlighted by the dashed line, above anoxic sediments, indicated by their darker coloration.
These include arthropods, annelids, and mollusks, which live in biogenic structures buried in the sediment. For example, chironomids dig and ventilate semipermanent U-shaped tubes; tubificid oligochetes live upside down, feeding on bottom sediment; and bivalves burrow into sediments with only their siphon linked to the surface – as illustrated in Figure 3. This biologically mediated flux of molecules and ions Fb (mol cm2 s1) is generally considered as a random phenomenon similar to diffusion. For a solute, the flux is defined as Fb ¼ – DB
qC qx 2 1
with DB the biodiffusion coefficient (cm s ). For a solid, the flux, Fbs, is defined as Fbs ¼ – ð1 – ÞDBs
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interface, a phenomenon that has been the subject of various modeling attempts. This process can be viewed as a nonlocal transport of solutes, with an exchange rate characterized by the irrigation coefficient (s1), that can be estimated from the size and density of the tubes as well as the solute gradients around a tube. In general, decreases to zero with depth, in order to reflect the decreasing density of active fauna. Although bioirrigation is a 3D process, a 1D equation has been used successfully in various models. This equation considers that the speed of intake or input of a solute by bioirrigation at a depth x can be described as q C ¼ ðCtube – C Þ qt x
with Ctube the solute concentration in the tube (mol l1) and t the time (s). In the calculation, Ctube is usually considered to be similar to the concentration measured in water overlying the sediment. Hydrodynamic Dispersion This process corresponds to the mixing of solutes resulting from hydrodynamic flows through permeable sediments and from the movement of waves at the sediment surface. It is usually approximated by a diffusion process function of the magnitude of the porewater velocity and the grain size. The flux is defined as Fd ¼ – Dvd
qC qx
with Dvd the vertical dispersion coefficient (cm2 s1).
qCs qx
with Fbs in g cm2 s1, DBs the bioturbation coefficient of the solid, and Cs the concentration of the solid compound in mol g1. In addition to the slight difference in formulations, it has been found that the effect of bioturbation on solutes can be 20 times more important than on solids.
Bioirrigation Since many organisms live below the oxic layer in surface sediments, they need to irrigate their tubes with oxygenated water from the water column. In addition to advecting oxygen into anoxic sediments, the renewal of water in the sediment structure fulfills other needs of the benthic organisms. These include the transport of food, metabolic wastes, gametes, and environmental stimuli. Bioirrigation of the tubes induces an increase of the fluxes of compounds in solution at the sediment–water
Measuring Nutrient and Contaminant Fluxes in Sediments Porewater Sampling: Squeezing, Peepers, DGT, Microelectrodes Since most nutrients and contaminants contained in sediments are predominantly associated with the solid phase, small changes in their sediment concentrations can translate into relatively large changes in their porewater concentrations. As a consequence, measurements of porewater concentrations are usually the most sensitive indicator of processes occurring in sediments. Similarly, measurements of concentration gradients in porewaters enable calculations of fluxes within sediments and across the sediment–water interface. Consequently, several methods have been developed for sampling sediment porewaters. (1) Interstitial water can be extracted by squeezing or centrifugation of successive slices of a sediment core, after it has been
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collected. (2) Interstitial water may also be collected in situ, using peepers. These are typically plastic devices with a vertical series of small chambers initially filled with high-purity water, which are enclosed within semipermeable membranes. The peepers are inserted into the sediments for an extended period (weeks) to allow the chambers to equilibrate with porewater, after which they are collected and the water in the peepers is extracted. (3) Interstitial water may also be collected in thin gels (DET–DGT) placed between two plates, which are inserted into sediments to allow equilibration of the gels with porewaters before they are collected and sampled. (4) Finally, microelectrodes can be either inserted directly in the sediment or in sediment cores immediately after sampling. Since there are advantages, disadvantages, and limitations for each of these methods, they are often used in combination to obtain a suite of complementary measurements, along with concurrent measurements of particulate sediment concentrations. Benthic Flux Chamber Measurements An alternative to sampling porewaters to evaluate fluxes of nutrients and contaminants at the sediment–water interface is the use of a benthic flux chamber. This type of device is essentially a container enclosing sediment and a small volume of overlying water that is incubated for periods lasting from a few hours to several weeks. During this time, concentrations of nutrients and contaminants in the overlying water are monitored at regular intervals, and changes indicate the direction and proportion of their fluxes; an increase in concentrations indicates a flux coming out of the sediment, while a decrease shows a flux into the sediment. Although flux chambers provide a direct measure of nutrient and contaminant fluxes, they are not without their shortcomings. Two of the more salient being that (1) the enclosed water will over time become anoxic with contact with the sediment, and (2) the enclosed water is secluded from surrounding benthic currents that mix water
and sediment. Unless artificially regulated, both of these phenomena will induce anomalous changes in the redox chemistry that will influence the measured fluxes. Sequential Extraction Methods While total concentrations of nutrients and contaminants in sediments are measured to quantify spatial and temporal gradients, concentrations of specific fractions on the surface and within sediments may be further characterized using selective extraction techniques. These extraction techniques provide more detailed information on how those constituents are bound in sediments, as well as their bioavailability and conditions for their remobilization from the sediments. As noted at the beginning of this article, bioavailability and mobility are especially important in studies of nutrients and contaminants in sediments because they are dynamic systems which are subject to chemical and physical changes on the timescales of seconds, days, seasons, years, decades, and longer. To address these concerns, various sequential extraction techniques have been developed to determine semiquantitatively how elements and compounds, including nutrients and contaminants, are associated with sediments. The extractions are designed to sequentially extract from the most weakly bound (e.g., ion exchangeable) to the most strongly bound (e.g., refractory crystal lattice) fractions. While the fractions are operationally defined, rather than definitive and truly specific, they have proved to be relatively useful in characterizing their relative bioavailability and potential for diagenic remobilization from sediments, as well as their biogeochemical cycling between sediments and water. Table 1 provides an example of a sequential extraction for sediments. It shows the extractants and conditions used in five sequential steps and the operationally defined fraction solubilized in each of those steps. There are numerous variations of these techniques (i.e., number of steps, extractants and conditions, and operationally defined fractions) with similar applicabilities and
Table 1 Sequential extraction steps and operationally defined fractions used to characterize the phase distribution of constituents, including nutrients and contaminants, in sediments Step no.
Extractants and conditions
Operationally defined fraction
1 2 3 4
1 M NaOAc, pH 8.2, 25 C 1 M NaOAc, pH 5 (HOAc), 25 C 0.04 M NH2OH-HCl in 25% (v/v) HOAc, 100 C 0.02 M HNO3 þ 30% H2O2, pH 2, 85 C; then 3.2 M NH4OAc in 20% (v/v) HNO3 Concentrated HNO3 þ HF, 110 C
Exchangeable Bound to carbonates Reducible, bound to Fe–Mn oxides Oxidizable, strongly bound to organics and sulfide Residual, recalcitrant
5
Modified from Tessier A, Campbell P G C, and Bisson M (1979) Sequential extraction procedure for the speciation of particulate trace-metals. Analytical Chemistry 51: 844–851.
Ecotoxicology | Sediments: Setting, Transport, Mineralization, and Modeling
limitations. Those limitations, along with previously noted limitations in measuring the speciation and fluxes of nutrients and contaminants in sediments, attest to their truly complex biogeochemical cycling and bioavailability. See also: Adsorption; Microbial Ecological Processes: Aerobic/Anaerobic; Physical Transport Processes in Ecology: Advection, Diffusion, and Dispersion; Transport in Porous Media.
Further Reading Allen HE (ed.) (1995) Metal Contaminated Aquatic Sediments. Chelsea, MI: Ann Arbor Press. Allen HE, Huang CP, Bailey GW, and Bowers AR (eds.) (1995) Metal Speciation and Contamination of Soils. Boca Raton, FL: Lewis Publishers. Baudo R, Giesy J, and Mantau H (eds.) (1990) Sediments: Chemistry and Toxicity of In-Place Pollutants. Ann Arbor, MI: Lewis Publishers. Berg P, Risgaard-Petersen N, and Rysgaard S (1998) Interpretation of measured concentration profiles in sediment pore water. Limnology and Oceanography 43: 1500–1510. Berg P, Rysgaard S, Funch P, and Sejr MK (2001) Effects of bioturbation on solutes and solids in marine sediments. Aquatic Microbiology and Ecology 26: 81–94. Berg P, Rysgaard S, and Thamdrup B (2003) Dynamic modelling of early diagenesis and nutrient cycling. A case study in an arctic marine sediment. American Journal of Science 303: 905–955. Berner RA (1971) Principles of Chemical Sedimentology. New York: McGraw-Hill.
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Berner RA (1980) Early Diagenesis. Princeton, NJ: Princeton University Press. Boudreau BP (1997) Diagenetic Models and their Implementation: Modelling Transport and Reactions in Aquatic Seduments. Berlin: Springer. Committee on Bioavailability of Contaminants in Soils and Sediments, Water Science and Technology Board, Division on Earth and Life Studies, National Research Council (2003) Bioavailability of Contaminants in Soils and Sediments: Processes, Tools, and Applications. Washington, DC: National Academies Press. DePinto JV, Lick W, and Paul JF (eds.) (1994) Transport and Transformation of Contaminants Near the Sediment–Water Interface. Boca Raton, FL: Lewis Publishers. Morel FMM and Hering JG (1993) Principles and Applications of Aquatic Chemistry. New York, NY: Wiley-Interscience. Schwarzenbach RP, Gschwend PM, and Imbode DM (eds.) (2002) Environmental Organic Chemistry, 2nd edn. New York: WileyInterscience. Song Y and Mu¨ller G (1999) Sediment–Water Interactions in Anoxic Freshwater Sediments: Mobility of Heavy Metals and Nutrients. Berlin: Springer. Stumm W and Morgan JJ (eds.) (1996) Aquatic Chemistry, Chemical Equilibria and Rates in Natural Waters, 3rd edn. New York: Wiley. Tessier A, Campbell P G C, and Bisson M (1979) Sequential extraction procedure for the speciation of particulate trace-metals. Analytical Chemistry 51: 844–851. Van Cappellen P and Gaillard J-F (1996) Biogeochemical dynamics in aquatic sediments. In: Lichtner PC, Steefel CI, and Oelkers EH (eds.) Reviews in Mineralogy, Vol. 34: Reactive Transport in Porous Media, pp. 336–376. Van Cappellen P and Wang Y (1996) Cycling of iron and manganese in surface sediments: A general theory for the coupled transport and reaction of carbon, oxygen, nitrogen, sulfur, iron, and manganese. American Journal of Science 296: 197–243.
Sediments: Setting, Transport, Mineralization, and Modeling L Kamp-Nielsen, University of Copenhagen, Hillerød, Denmark ª 2008 Elsevier B.V. All rights reserved.
Net Deposition as an Aggregated Approach Theoretical Models
Further Reading
Sediments are particulate matter that can be or have been transported by fluid, wind, and glaciers and which might have been deposited as a layer of solid particles in dense suspension at the bottom of water bodies. The parts of sediments which have their origin outside the water bodies are called allochthonous sediments and have been transported by runoff from the drainage basin of the water body or by wet or dry deposition on the surface of the water body. Particles formed within the water body and the sediment are the autochthonous parts of the sediments and are transformations of dissolved elements to particles by chemical and biological processes. In the photic zone of rivers, lakes, coastal waters, and oceans, inorganic carbon as carbon dioxide or bicarbonate is fixed as particulate organic carbon by photosynthetic
organisms, or fixed as inorganic carbonate in corals, foramifers, and coccolithophorids. The dead and living organic particles can be processed through the aquatic food chain and sink to the bottom as living or dead particulate organic material. Due to increased pH, as a result of photosynthesis, the solubility product of calcium and magnesium carbonate can be exceeded and precipitation of carbonates may occur. Both the allochthonous and the autochthonous parts of the sediments are subject to further biogenetical processes in the sediment environment and other autogenic fractions are generated. On a geological timescale, surface sediments are young structures, but they play an important role in the global and local cycling of elements like carbon, nutrients, and metals – all of which are important for the productivity in