Modeling solute transport and sulfate reduction in marsh sediments

Geochmicn et Cosmochimica Ada Vol. 51, pp. 1109-l 0 Pergamon Journals Ltd. 1987. Printed in U.S.A.

I20

0016.7037/87/$3.00

+ .oO

Modeling solute transport and sulfate reduction in marsh sediments* WILLIAM H. CASEY’ and ANTONIO C. LASAGA~ Department of Geosciences, The Pennsylvania State University, University Park, PA 16802, U.S.A. (Received May 2 1, 1985; accepted in revised form January 30, 1987)

Abstract-The seasonal oscillation in sulfate and chloride concentration profiles in some salt marsh sediments is due to exchange of solutes with water on the surface of the marsh, and to the desiccation of the sediment in summer. Desiccation is manifested by disappearance of surface waters, fluctuations in the location of the water table, and by removal of water from the sediment above the water table. The loss of water from the pore space is commonly accompanied by entry of air into the soil, which oxidizes sulfide. The oxidation causes titratable alkalinity to decrease and results in CO2 degassing. Diffusion models of salinity can account for the observed profiles but only as long as the marsh is maintained inundated. The complexities introduced to the solute transport equations by sediment desiccation invalidate steady-state modeling of solute transport and diagenesis. The concentration profiles of dissolved products of sulfate reduction, such as bicarbonate, require months to reestablish a steady state after being disrupted. If the profiles of dissolved products of sulfate reduction are disrupted seasonally, such as by a seasonal fluctuation in the water table, they may remain transient throughout the year. INTRODUCTION THE HIGH BIOLOGIC productivity of salt marshes has attracted biologists and geochemists interested in the flux of nutrients and chemical energy to the surrounding bays and estuaries. This productivity is closely tied to the transport of the dissolved byproducts of sulfate reduction out of the marsh sediment via pore fluid (cJ WIEGERT et al., 1983; MENDELSOHN and SENECA, 1980; KING et al., 1982; PETERSEN et al., 1983). This article addresses the nature of solute transport in salt marsh sediments where the pore water profiles of chloride and sulfate change seasonally (CHAPMAN, 1960; LORD and CHURCH, 1983). LORD and CHURCH (1983) made an important contribution when they treated the seasonal variation in pore fluid chemistry as resulting from diffusion, burial, and reaction of ions. The upper boundary condition for their model is a seasonal oscillation in salinity which, although not directly measured, was inferred from the shape of observed solute concentration profiles in the O-25 cm depth interval at various times during a year. Their model requires inundated sediment, no vertical groundwater flow, and that time-averaged profiles of solutes which are produced during sulfate reduction, such as ammonia, sulfide and bicarbonate, are at steady state. In this article, we address the cause of these transient concentration profiles, and the suitability of steady-state models to describe this chemistry. STUDY SITE AND TECHNIQUES Field work and analytical chemistry

The Palus Crisium study site is a peninsula covering about 0.1 km* of a marsh island in Chincoteague Bay, Virginia (Fig.

* This work was supported in part by the U.S. Department of Energy under contract number DE-AC04-76DPO0789. ’ Present address: Sandia National Laboratories, Geochemistry Division 1543, Albuquerque, NM 87 185, U.S.A. ‘Present address: Kline Geology Laboratory, Yale University, P.O. Box 6666, New Haven, CT 065 11, U.S.A.

1). Vegetation at the site is dominated by Spartina a/fernifIoru with lesser amounts of Spartina patens. Roughly 20 percent of the area at the site is unvegetated, or is very sparsely vegetated by Salicornia virginiana. The positions of these sites (pannes) were surveyed, and a single, large panne with an area of 4300 m* was chosen for detailed study of the pore water profiles. This panne will be labeled by LP in what follows. We chose to study the pore fluid chemistry in an unvegetated panne because: 1) the unvegetated pannes are topographically 5 to 15 cm lower than vegetated locations and retain surface water for a large fraction of the year; 2) solute transport in vegetated locations is complicated by plant uptake of solutes and water (e.g., CARLSON and FORREST,1982; DACEYand HOWE&1984); and 3) oxygen may be transported along plant roots (e.g., TEALand RANWISHER,1966; LUTHER~~al., 1982). While processes of solute transport are relatively simple in pannes, rates of sulfate reduction are nearly as high as in vegetated portions of the marsh (HOWARTHand MARINO, 1984; MARTENSand BERNER,1974; NEDWELLand ABRAM,1978). Thus, pannes provide a relatively simple subenvironment of the marsh for studying sediment chemistry and for modeling solute transport. Shallow cores were taken periodically in a ten-meter-wide area at the LP site (Fig. I), and pore waters sampled for sulfate and chloride concentration. Protiles of bicarbonate and sulfide were measured using Hesslein in situ samplers (HESSLEIN, 1976) to minimize loss of volatile components. The polycarbonate sample cells were covered with a Gelman Versapor filter membrane (0.2 micron pore diameter) which was glued into place with polycarbonate shavings dissolved into methylene chloride. The samplers are biologically inert (G. R. HOLDREN,pets. commun.). The in situ samplers were allowed to equilibrate for three months before sampling. We estimate that equilibration between pore fluid and the sampler occurs within one month, based upon compared profiles of salinity. The analytical techniques used to measure sulfate concentration are those outlined by PRESLEY(197 1). Alkalinity was measured by potentiometric titration using a standardized acid. The pH of the sediment was measured with a combination electrode which was inserted into the sediment. The electrode was calibrated with potassium biphthalate (pH = 4.0) and potassium phosphate-sodium hydroxide (CLINE, 1973). The precision of this technique is + 10 percent and was determined by analysis of standard sulfide solutions. The sulfide concentration of the standards was determined by potentiometric titration using a sulfide-specific ion electrode and lead perchlorate. Salinities of surface waters were measured in the field with an American Optical refractometer; which is accurate to within

1109

W. H. Casey and A. C. Lasaga

1110

C HINCOTEACUE

water content was determined by weighing the sample before. and after drying at 50°C. Uncertainty in these measurements are generally less than 5 percent.

BAY

RESULTS Surface salinities

(/

PALUS

CRISIUM

) / 38O

In large part. the solute profiles observed in marsh pore waters are generated in response to dramatic changes occurring on the surface of the marsh. ‘The profound effect of these surface processes necessitates measurement of surface concentration as a function of time. In particular, unlike normal marine sediments. the salinity can change quite drastically in marsh pore waters. The composition of surface salinities from four representative locations on the Palus Crisium site are plotted in Fig. 2. These four pannes were chosen to illustrate the seasonal change in surface water com00’ N positions on the marsh, and are representative of the other sites of roughly equal elevation where salinities

FIG. 1. The Palus Crisium study site. Large, semi-permanent bodies of standing water are outlined with a solid line. The symbols correspond to the ponds used to compile Fig. 2.

0.5%. The salinity of pore fluid was measured by titration for chloride (PRESLEY,197 I), except for the October core, which was measured using the refractometer. Accuracy and precision of the titration is within 0.2%. On selected profiles, the concentration of major cations was determined by atomic absorption spectroscopy using the techniques outlined by PRESLEY (1971). The analyses indicate that, to within analytical uncertainty (+5 percent), chloride concentration in pore fluid has the same proportionality to salinity as in normal seawater

were also measured through the year. Note that the data at all four sites generate essentially one common curve. Therefore, the surface salinities reflect processes that are not localized within a marsh. It is important to note also that the seasonal variation of the common surface salinity on the marsh defines a crude sinusoid with a mean salinity of about 20%. The portion of the year excluded from the solid curve in Fig. 2 represents that part of the year when the marsh surface is dry. Standing water disappears in late spring and does not reappear until early fall. In summer, water is removed from the marsh surface in hours to days after complete inundation by infiltration, evaporation. and transpiration. Surface water disappears earlier from the vegetated areas than from the pannes on the marsh, because vegetated locations are at a high elevation, and water is removed by plant transpiration in addition to evaporation. Transpiration is in the range of 0. I to 0.9 ml/cm*/day (e.g., MORRIS and WHITING, 1982; TEAL. and KANWISHER, 1970; DACEY and How& 1984; HOWES et ul., 1987), which is roughly equal to the pan

(CASEY, 1985).

Hydrologic properties The location of the water table was monitored with nested piezometers at ten locations in vegetated and unvegetated areas throughout the marsh (CASEY, 1985; Appendix E). PVC pipe (1.5 cm i.d.) was inserted into the peat to four depths (30,60, 130, and 260 cm) at each nest. An internal mandrel was used to prevent the entry of sediment into the piezometer during emplacement. The internal mandrel was removed following emplacement and the piezometers were allowed to equilibrate for five months. The location of the water table was measured using a calibrated dipstick with accuracy of rtl.0 cm. The shallowest piezometers (30 cm depth) require several hours or days to respond to a 15 cm change in the elevation of the water table. The term ‘water table’ is used to mean the saturated piezometric surface (BEAR, 1972), and does not imply any relation to a freshwater aquifer. Water-saturated porosity in peat was measured volumetrically. A core of peat was extruded into aluminum rings of known volume (90 cm3), and cut with a straight razor. The

o-LP

I

-

IJ

JY

A

S

0

SITE

N

D

J

F

M

A

M

J

Jy

FIG. 2. Seasonal variation in the composition of surface water in the four representative ponds identified in Fig. 1.The solid symbols indicate periods when the surface was dry, and are not intended to indicate a salinity. The stippled area identifies that portion of the year that the marsh surface is intermittently wet and dry. The line is a least-squares fit to the data from the LP site.

1111

Marsh sediment water modeling evaporation rate we measured above the marsh surface in summer (about 0.5 ml/cm’/day; CASEY, 1985). The salinity of surface water increases very rapidly as the marsh dries in spring, reaching values higher than 40%0 under virtually dry conditions. Periods of dryness, a relatively small fraction (0.3) of the year, are represented by solid symbols in Fig. 2. Although the marsh surface is generally dry during late spring and summer, the marsh continues to be flooded at infrequent intervals throughout the summer by rain and tides. These periodic inundations are illustrated by the few points in the dashed and stippled area of Fig. 2, which are not on the solid curve. The solid curve in Fig. 2 represents a least-squares (see Appendix) fitting curve through the surface saiinities for the “wet” portion of the year at the LP site. This least-squares curve will be used later as a boundary condition to model solute transport in sediment at the LP site.

below the marsh surface. Figure 3a gives the measured water table at a vegetated area immediately adjacent to the LP site based on the piezometer nest discussed earlier. Prior to late spring, the water table is at, or slightly above, the sediment surface. Figure 3a also shows that from late spring through early fall, the water table was consistently located 25-35 cm below the surFace in both vegetated and unvegetated portions of the marsh (see CASEY, 1985; Appendix E). The drop in the water table is accompanied by a dramatic decrease in the water content of the sediment. The water content of cores from the LP site in August are presented in Fig. 3b. Note that, before tidal inundation, water content decreases toward the marsh surface in the upper 10 cm of sediment. After inundation the water content over the same interval is nearly constant with depth, and mixing of pore fluid can be detected over the depth interval O-10 cm (CASEY, 1985). Mute

concenfration prqfles

Water table

Disappearance of surface water from the marsh coincides with a drop in the elevation of the water table

The concentration profiles of solutes such as bicarbonate and sulfide generally increase with depth in normal marine sediments (BERNER, 1980). However,

PIEZOMETER

MaI.

Apr.

May

June

July

Aug.

Sept.

Oct.

Nov.

DEPTk

Dec.

IO. AFTER

0

0.50 VOLUMETRIC

FLOODING

1.00 WATER

FRACTION

PIG. 3a. The location of the groundwater surface drops to below the marsh surface (0.0 cm on the figure) during late spring, summer, and early fall. FIG. 3b. Water content of the marsh soil at the LP site in August just before, and 24 hours after, tidal flooding.

W. H. Casey and A. C. Lasaga

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August

0 cm -

October

a3 ocm

83

.

August

Ocm

a3

October

Ocm

83

(dry)

(dry) 20cm-

2Ocm

20 cm

0

20 cm

P P .

.

. 40 cm -I

40 cm

l

.

40 cm

i’

40 cm

.

. m

l--2IO 0 Ocm

.

500

20

January 0

20cm.

April

84

84

April

January 84

Ocm

84

. : ;

0 z

20 cm

.n . ‘3 90 .O

. .

0 .

40 cm

0 .

40cm-

c

-00

.

0

0

OIb

0

Suiflde

Dissolved Titratable

Alkalinity

meq

500

,kmoles

i-l

I_’

FIG. 4. Seasonal variation in the titratable alkalinity at the LP site. The two symbols correspond to samples from the

FIG. 5. Seasonal variation in dissolved sulfide concentration in pore fluid. The symbols represent samples from the two sides of the in situ sampler.

sides of the in situ sample, about 30 cm apart.

in the upper 15 cm of marsh soil the concentrations vary considerably throughout the year (Figs. 4 and 5). Titratable alkalinity is very low in the uppermost O15 cm during August, when the soil is desiccated and pore water salinities are very high. Because of its effect on COZ speciation, the pH of the pore waters was also measured seasonally. The pH of the shallow marsh sediment is found to decrease slightly in summer. PH values range from about 5.5 to 6.5 when the sediment is inundated tobetween 4.0 and 6.0 in summer (CASEY,

company unsaturation for vegetated locations as well, we report piezometric data and pore fluid salinities from a site vegetated by Spartina alterniflora in Fig. 7. Note that during summer, when the water table lies below the marsh surface (see also Fig. 3b), pore fluid nearest the marsh surface is extremely saline (Fig. 71. Pore fluid a few millimeters or centimeters deeper is relatively less saline (40%0). Thus, sediment unsaturation accompanies hypersaline pore fluid salinities throughout the Palus Crisium site.

1985).

Both sulfate concentration and salinity exhibit strong seasonal oscillations in the depth interval O-25 cm and increase to very high values at the marsh surface in summer (Figs. 6 and 7). At sufficient depths in the sediment, the salinity of pore fluids is largely unaffected by the seasonal changes taking place near the sediment surface (e.g. see Fig. 6). In muddy coastal sediments of low permeability, this depth is reached below 25 cm (HOLDREN et al., 1975; LORD and CHURCH, 1983). The composition of pore fluids at this depth is a rough measure of the time-averaged composition nearer to the sediment surface. We refer to this value as the seasonally averaged composition. Figure 6 shows that the seasonally averaged salinity of pore fluid at 25 cm depth is about 70%0, which is significantly higher than the mean salinity of surface water on the marsh (about 20?&), or of the adjacent bay (about 3 1%o). To demonstrate that hypersalinity conditions ac-

.

4

,

0

.

,

,

250 Salinity

(%.J

&-_--_----.

0 Sulfate

300 (mmoles/I

FIG. 6. Depth profiles of salinity and sulfate concentration

at the LP site throughout a year.

!

Marsh sediment water modeling

$a

gv-30;

A* . I 9 I 1 I I 1 I I Mar.Apr. May Jun. Jul. Aug. Sap. Oct. Nov. Dec.

i.

SALINITY -z u

100

cl0

I

%o 200

--

/--June

12,

1984

-

B. FIG. 7a. The location of the water table at a site vegetated by Spartina alternijlora.The piezometer depth is 30 cm. Note that the water table lies below the marsh surface in summer. FIG. 7b. Depth profile of salinity at the vegetated site during June.

DISCUSSION

Interpretation of the pore fluid chemistry Four processes can affect the salinity of the surface waters on the marsh: 1) flooding by seawater; 2) dilution by rainfall; 3) exchange of solutes with the sediment; and 4) evapotranspiration. The net result of these processes is a salinity of surface waters which changes in a regular fashion through that part of the year when the marsh is not dry (e.g. Fig. 2). The average salinity of surface marsh water (20%0) is less than the salinity of seawater (3 1%o) in the adjacent bay and indicates that rainfall is as important as tidal flooding in controlling the composition of surface waters for much of the year. Two processes largely control the seasonal variation in salinity profiles with depth: 1) diffusion of Na+ and Cl- between the surface waters of varying salinity and the sediment pore water and 2) hypersaline conditions caused by desiccation of the sediment in summer. Drying of the near-surface sediment in summer causes both rates and mechanisms of solute transport to differ from the corresponding processes in winter, when the sediment is inundated. As marsh sediment dries, the decrease in water increases the salinity near the sediment surface at a faster rate than Na+ and Cl- can be transported to deeper pore fluids. Thus, salinity profiles are extremely steep in the O-5 cm depth interval, except for short periods immediately after tidal flooding or rain storms (see Table 1.1 Of CASEY, 1985). This sharp gradient is present in both vegetated and unvegetated

1113

locations in the marsh (cf LORD and CHURCH, 1983), and is typical of unsaturated, saline soils undergoing evaporation at the surface (GARDNER, 1959; BRESLER et al., 1982). The evaporation and transpiration of water from the surface of the marsh continues as the underlying sediment becomes water-unsaturated, leading to an increase in the salinity of marsh pore fluids. DACEY and HOWES ( 1984) and MORRIS and WHITING ( 1985) find that marsh soil may become unsaturated due to transpiration of Spartina alone following diurnal tidal inundation. Our results indicate that unsaturation for at least part of the year is probably required for all marsh soils which have average pore fluid salinities significantly higher than normal seawater. Although the salinity of pore fluid at 25 cm is relatively constant (at =70%o--see Fig. 6) this value does not represent the mean variation in salinity of surface waters. The salinity at 25 cm depth reflects several factors including: 1) the frequency of inundation by tides and rain; 2) the depth of standing water at the site following inundation; 3) the relative proportion of the year that the site is wet or dry; 4) the rate of evaporation and transpiration of water from the sediment; 5) the rate that solutes are exchanged between sediment and the overlying water; and, 6) the rate of solute transport to deeper sediment. Bicarbonate and sulfide are strongly affected by fluctuations in the water table (Figs. 4 and 5). These species are produced primarily as byproducts of sulfate reduction (BERNER, 1980) and are not present in appreciable quantities in normal seawater. Therefore, concentration of seawater by evaporation causes only minor increases in bicarbonate and sulfide concentration. However, the entry of oxygen into the marsh soil during unsaturated conditions leads to large decreases in the local concentration of bicarbonate and sulfide. Therefore, in the summer there is almost no titratable alkalinity and sulfide in the top 15 cm (unsaturated zone-Fig. 5) while the amount of both species increases sharply below 15 cm. The oxygen decreases the sulfide content by causing dissolved or solid sulfide to oxidize and produces sulfuric acid. The acid, in turn, neutralizes bicarbonate and reduces titratable alkalinity in the pore fluid. As there are no live plant rhizomes or roots at the LP site, we conclude that the decrease in alkalinity results from the oxidation of reduced sulfur species, production of sulfuric acid and the concomitant CO2 formation induced by the drop in pH. This interpretation is supported by HOWES et al. (1987), DACEY and HOWES ( 1984) and MORRIS and WHITING (1985) who report the entry of air into vegetated sediment during a fluctuation in the elevation of the water table. In their studies, the fluctuation was caused by plant transpiration of sediment pore water, and induced the oxidation of reduced sulfur compounds in the sediment. The decrease in alkalinity which we observe may also be partly attributable to the precipitation of carbonate minerals in the first few centimeters of sedi-

W. H. Casey and A.C.Lasaga

1114

ment. We find, however, that the ratio ofthe dissolved calcium to sodium in pore fluid (which has high HCO;) is in~stinguishable from normal seawater ratios in summer when the sediment is desiccated (CASEY,1985; Appendix A). We thus conclude that the most important cause of the decrease in alkalinity is titration by the sulfuric acid generated during oxidation of reduced sulfur compounds in sediment following sediment un~turation. MODELING SOLUTE TRANSPORT AND SULFATE

REDUCTION

One of the aims of this paper is to use models to explore some of the complex phenomena discussed in the earlier sections. In particular, the effects of periodic unsaturation at the surface of the marsh and the variations in surface composition during the year must be addressed by these models. In general, we must start with some form of a transport-chemical reaction equation. Ignoring biologic mixing, irrigation, and osmotic effects, a one-dimensional equation describing the chemical reaction and transport of solutes in sediment is (BERNER, 1980; BRESLERet al., 1982):

where: e &

D, GR

V

U’

Z

(2,

is concentration of the ion (moles/cm3 of SOlution), is the ion diffusion coefficient in sediment, which varies with depth in the soil and with time (cm2/yr), is a coefficient of mechanical dispersion that arises if u is not zero (cm2/yr), is the sum of rates of reactions that change the concentration of the solute. Changes in concentration resulting from the evapotranspiration of pore water are included in this term (moles/cm3 of ~lution/year~, is the velocity of fluid past the sediment grains. This may be due to flow caused by compaction, or by groundwater flow (cm/y@, is the velocity of solids relative to the sediment surface (cm/yr). For steady state sediment accumulation in the absence of compaction, this is constant and equal to the sedimentation rate, is the distance (depth) below the sediment surface (cm); z = 0 is defined as the sediment surface, and is the effective porosity of the sediment which varies with time and depth.

Equation (1) must be simplified further to explore the various processes discussed earlier because we lack detailed data on the important parameters. One central theme in this paper is the crucial role played by un~tumtion in the chemistry of marshes. Sediment unsaturation affects several of the terms in

Eqn. ( 1). The loss of wetted porosity increases the tortuosity of the remaining pore space, thereby decreasing the sediment diffusion coeflicient. Partial saturation of marsh soil further affects the sediment diffusion coefficients by creating gradients in ionic strength in the pore fluids. Changes in the ionic strength affects the sediment diffusion coefficients by increasing the solution viscosity and by altering the ion activity gradient (LASAGA,i 979). In Eqn. (2) we present an expression which accounts for these el%cts (OLSENeial., 1965; PORTERet al., 1960; ONSAGERand RJOSS, 1932; LASAGA. 1979):

where: is the tracer diffusion coefficient of the ion in normal seawater (cm2/yr), ~/~~ is the ratio of the viscosity of the pore fluid to that of normal seawater+ is the individual ion activity coetlicient. Y h is an empirical constant, is the volumetric water content of the soil, which cb we consider equal to the effective porosity, and !Y is an empirical constant (OLSEN and KEMPER, 1968).

00

By far, the most important term in the above equation describes the variation in the sediment diffusion

coefficient with porosity and tortuosity of the sediment (in the far right bracket of Eqn. 2). This term is whoily empirical and must be evaluated for different soils and conditions. For the conditions presented by OLSENand KEMPER (1968)(0.1 < 4 -c0.6, K = 001 to ,005, b = 10,soil suction pressures of 0.33 to 15.0 bars), diffusion is virtually halted as the water content decreases from values of 0.6 to 0.3 and the pore lluid loses connectivity, The effect is more pronounced for clay-rich soils than for sandy loams, but no data presently exist for marsh soils (see HARRELLand SAEED, 1980). Although gradients in ionic strength, and hence viscosity and activity, are secondary in importance relative to variations in the porosity and tortuosity, these effects alone can lead to a 40% vacation in diffusion coethcients in shallow marsh sediment (LASAGA,1979). We find that pore fluid salinities in both vegetated and unvegetated portions of the marsh approach saturation with sodium chloride in summer, and that the gradients in composition are extremely steep (Fig. 7). The advection term in Eqn. (1) may include a contribution from groundwater flow. When the marsh soil is unsaturated, recharge comes slowly from below, as pore fluid is absorbed upward along a gradient in cap illary pressure, and quickly from above, as the marsh is inundated. In a combined field and model study, HEMONDand FIFELD(1982) found #at steady upward flow of pore fluid can be induced by plant ~nspimtion alone, provided that there is a source of recharge at

Marsh sediment water modeling depth. A their site, recharge was through a sand aquifer beneath the marsh that contained fresh water. There is no identifiable source of recharge in the deeper sediment beneath the Palus Crisium site, which was chosen for its hydrologic isolation in the middle of Chincoteague Bay (Fig. 1). Permeabilities at depths of about 300 cm are estimated to be much less than lo-” to IO-‘* cm2 (CASEY, 1985; Appendix E). These permeabilities are too low to permit velocities of pore water flow greater than a few centimeters per year to be induced by the topographic head of the marsh relative to the adjacent bay. The piezometric surface of the Palus Crisium site lies above mean sea level at all times of the year, suggesting that the marsh soil is recharged by the infiltration of surface waters and not by upward flow from the adjacent bay. Note that a term is included in Eqn. (1) to account for mechanical dispersion of solutes. Mechanical dispersion may be a significant effect if unsaturation of the soil induces flow of pore water (e.g., BEAR, 1972; BRESLERet al., 1982), or if tidal influences cause regular variations in pore pressure. In this case, mechanical dispersion can arise even if there is no net flow of pore fluid (HEMOND and FIFIELD, 1982). In the subsequent discussion, we solve a simpler form of Eqn. (l), when necessary, by finite difference methods on a Prime 7200 computer. A Crank-Nicolson implicit scheme was employed and solved by recursive Gaussian elimination of the resulting t&diagonal matrix. The space step and time step were 1.0 cm and 0.01 years, respectively. Accuracy was checked by comparison with analytic solutions, and the error in the approximation is less than 0.1% of the true value for the test cases. Chloride concentration profiles If properly constrained, Eqn. (1) can be used to calculate reaction rates from the concentration of solutes in sediment pore fluids. Many of the parameters needed to evaluate the equation, however, are difficult to measure in marsh sediments as a function of time. As a result, highly simplified models are often used in modeling marshes. Therefore, in this section, we examine whether profiles of conservative solutes can be predicted using a simple equation and measured salinities of surface waters. Most marine sediments are treated with constant diffusion coefficients, no dispersion, no flow of pore fluid, and with negligible gradients in porosity. In such cases, Eqn. (1) reduces to (BERNER, 1980):

1115

For this reason, all of the coefficients were taken from independent sources, and were not adjusted to improve the fit of the observed and predicted concentration profiles. We employed a porosity of 0.8, a sediment diffusion coefficient of 200 cm2/yr (LI and GREGORY, 1974), and a reasonable burial rate of 0.31 cm/yr (CASEY, 1985; BARTBERGER,1976). These results differ from the LORD and CHURCH (1983) treatment in that the concentration of surface water is explicitly measured, not inferred. The transport of solutes in sediment during early diagenesis is typically modeled by employing the concentration at the sediment surface as a boundary condition for the transport equations. Because the marsh is not overlain by surface water for roughly 30 percent of the year, the nature of the boundary condition where the concentration of solutes is known, but varying through time (cJ, HOLDRENet al., 1975), is clearly appropriate when the composition of surface waters can be explicitly measured, but loses physical significance when the sediment is unsaturated. A better boundary condition for modeling ion transport would employ the measured flux of water out of the marsh when the surface is dry and the soil is unsaturated. Measuring the flux takes the loss of porosity, and the resulting increase in salinity of the pore fluid into consideration. In this paper, we will have to simplify the boundary condition to the surface concentration measured during the year. We used a least-squares fit to the measured composition of surface water as the boundary condition for periods when the marsh was inundated (Fig. 2, Appendix). For periods when the marsh surface was dry, the salinity of the surface water was set equal to 200%0, a salinity of pore fluid which we commonly observed in the O1 cm depth interval of the soil. The lower boundary condition was a constant concentration at great depth (750 cm) in the sediment. The results are not sensitive to the location of the lower boundary. Predicted and observed salinity profiles are presented in Fig. 8. Note that the profiles compare well for periods when the soil is inundated and a surface composition could be explicitly measured. That the predicted and

Predicted from surface composilion D= 200 cI+ yi’

If the solute transport is at a steady state, the right side of Eqn. (3) is set to zero. We first solved Eqn. (3) numerically to predict salinity profiles at the LP site through the year. The purpose of this analysis is to test the sensitivity of the transport model to periodic unsaturation of the soil.

!+----Sollnlty

%.

250

0

Sollnlty %.

250

FIG. 8. The predicted and calculated salinity profiles for pore fluids in marsh soil at the LP site. The solid circles on the measured profile identify data points.

W. H. Casey and A. C. Lasaga

Ill6

observed concentration profiles compare well when the marsh surface was inundated indicates that solute transport at these times can be treated in a similar fashion as subtidal environments with transient concentration profiles ($ HOLDRENd al., 1975). We conclude that when the marsh is inundated and there is no gradient in pore pressure to induce flow, ion transport is dominated by diffusion. There is disagreement near the sediment surface between the predicted and measured profiles for the May and July cores. This disagreement stems from the choice of a boundary salinity for the period from late spring to July, and from the changing nature of solute transport as the sediment becomes unsaturated. It is conceivable that agreement could be improved if the pore fluid in the O-l cm interval were sampled more frequently. It is much more likely, however, that the poor fit between predicted and observed concentration profiles is caused by complexity in solute transport arising from the loss of soil water. For that part of the year when the marsh is not overlain by surface water. which includes much of the growing season of Spartinu alterniflora, ion transport cannot be explained without addressing the complexities which arise when the water table drops below the marsh surface, and water is removed from the soil. The steady-state model for sulfate reductiort In this section we examine the validity of the steady state approximation for marsh sediment undergoing periodic disruption of the solute concentration profiles. The most conspicuous diagenetic process is the decomposition of organic matter and sulfate reduction by heterotrophic bacteria. In this case. the stoichiometry of reaction is known to be one sulfate ion destroyed per two carbon dioxide ions produced (WESTRICH, 1983): 2CHzO(NH,), + SO;’ + 2@H’ = 2HCO7 + 2PNH: + HS + H’

(4)

where fi is a stoichiometric constant. Other products of sulfate reduction, such as phosphate, are excluded from the reaction for the sake of simplicity. If the profiles can be treated as steady state profiles and simple linear kinetics are valid for the chemical reactions, first order rate constants for organic matter decay (k) can be calculated from the distribution with depth in the sediment pore water of sulfate and the dissolved products of sulfate reduction. The solution to Eqn. (3) for a constant influx of reactive organic matter to the sediment surface, a constant concentration of reactive organic matter at depth, and a constant concentration of the reactive solute at the sediment surface, is:

--$$$+Co s

C(z) = [ 1 - e&kz’w)] [

I

G(Z) = G”e-(kZ’v) organic matter

solutes

(5) (6)

where: F G(z)

i;

is a stoichiometric factor relating the reactants and products of sulfate reduction, is the concentration of reactive organic matter at depth, z in the sediment, expressed in units of moles of carbon per cm3 of solution. The concentration at the surface is Go, and is the first order rate constant (yr’),

One of the central aspects of marsh diagenesis is the periodic oscillation in near-surface sediment chemistry due to seasonal changes and dessication at the surface of the marsh. Therefore, we must solve Eqn. (3) for periodic boundary conditions. Thus, the assumptions implicit in Eqn. (5) are not strictly valid for modeling marsh soil which becomes periodically unsaturated. It is possible, however, to solve Eqn. (3) by numerical approximation and evaluate the influence of the nonsteady-state variables on the profiles. The results are discussed in what follows. In the steady-state approach to modeling sediment diagenesis, rate coefficients calculated from sulfate profiles using Eqns. (5) and (6) (steady-state solutions) are compared with coefficients calculated by modeling products of sulfate reduction such as bicarbonate, ammonia, or phosphorus using a modified form of Eqn. (5). In the LORD and CHURCH (1983) adaptation of the steady-state model, the concentration profiles of the dissolved products of sulfate reduction are assumed to be at steady state beneath a permanent zone where they are consumed. This is illustrated in Figs. 9a and 9b, where hypothetical profiles of bicarbonate are shown. In Fig. 9a, profiles are shown which extend to the sediment surface. This corresponds to the case of no recycling of solutes in the sediment. The transposed profile is shown in Fig. 9b below a zone where bicarbonate is consumed, either by root uptake or reaction with sulfuric acid in the pore fluids. Rate coefficients (k) as well as Go are calculated by fitting to Eqn. (5) the concentration profiles of products below this zone of consumption. Requiring examination is the assumption that these concentration profiles are at steady state. We fmd that the most important cause of the seasonal variations in the concentration profiles of sulfide in marsh soil is oxidation resulting from fluctuations in the water table. This oxidation of sulfide produces sulfuric acid which neutralizes bicarbonate. Recent work (DACEY and HOWE& 1984; MORRISand WHITING, 1985; HEMOND and FIFIELD, 1982) indicates that periodic fluctuations in the elevation of the water table are a general feature of Spartina altemifora marshes. If solute concentration profiles are to be used to calculate reaction rates in marsh soil, they must return quickly to a steady-state profile after being disrupted by a fluctuation in the water table. We numerically simulated the evolution of bicarbonate profiles exposed to unsaturation in marsh soil to examine the speed at which they return to steady state. We chose to model bicarbonate because the stoi-

1117

Marsh sediment water modeling

FIG. 9a. A simulated steady-state profile of bicarbonate that extends from the sediment surface. FIG. 9b. The steady-state profiles of bicarbonate beneath a 15-cm thick zone in the sediment where bicarbonate is consumed. The coefficients used to generate Figs. 9a and 9b are: G” = 1600 pmoles/cm3, k = 0.03 yr-‘, 0, = 110 cm2/yr, w = 0.31 cm/yr.

of reaction is immediately known from Eqn. (4). The results, however, are also appropriate for more complicated solutes, such as phosphorus and ammonia. Two types of simulations were treated, and are shown in Fig. 10.

chiometty

( 1) The initial condition is a steady-state (i.e., Eqn. (5)) profile of bicarbonate concentration except that it is altered by setting the concentration of bicarbonate equal to 2.2 nmoles/l in the 0 to 15 cm interval of the sediment (Fig. 9b). We can then examine the time necessary to reestablish a steady-state profile for these new boundary conditions. This is intended to model a natural system where bicarbonate ion is neutralized by acid as rapidly as it is produced by the bacterial oxidation of organic matter. Below 15 cm, the evolution in the bicarbonate profile was modeled by solving the transport equations (3, 4, and 5) (Fig. 10a). This simulation assumes that the concentration at 15 cm is constant, i.e., that products of sulfate reduction, such as bicarbonate, are permanently consumed in the O15 cm interval of the soil (z I 15 cm, t > 0) = 2.2 mmole/l. (2) In the second simulation, (Fig. lob), we used the same steady-state profile of bicarbonate concentration as the initial condition, but assumed that it is consumed in the 0.15 cm interval for only 0.3 years.

During this time, the concentration of bicarbonate in the 0- 15 cm interval was 2.2 mmole/l. After 0.3 years, the profile was allowed to return to the initial steadystate profile in a manner described by the transport equations (3, 4 and 5). We intended this case to simulate the seasonal disruption of solute profiles for about four months. We used a constant concentration (steady state) of bicarbonate at 750 cm for the lower boundary condition in both simulations. Altering this lower boundary concentration did not affect the results, as 750 cm is far from the surface zone where our interest lies. The first simulation (Fig. 10a) indicates that several decades or centuries are required to re-establish a steady-state, once the profile is disturbed in shallow marsh soil (Table 1). This length period when the profiles are transient is caused by a large upward, diffusive flux of bicarbonate from deeper pore fluids. In BERNER'S (1980) solution to the transport equation, the concentration of solutes produced during sulfate reduction approaches a constant value with depth in the sediment. This concentration, however, differs for steady-state profiles which extend from different locations in the sediment column. For example, for the conditions used to generate Fig. 10a (Go = 1600 mmoles/l, k = 0.03 yr-‘), the concentration of bicar-

8.

A

I = 0 yrs.

steady

state

steady

stole

J-.

0

mm&s 1.' 50

FIG. 10a. Formation of a permanent zone of nutrient consumption above the anoxic sediments. The steady-state profiles, which can be used to calculate reaction rates via Eqn. (4) are shown as dotted lines. The calculated transient profiles are solid. FIG. lob. Bicarbonate concentration was consumed in the near-surface zone only for 0.3 years. The solid lines represent the transient concentration profiles. The dotted line represents the steady-state profile used as the initial condition. The concentration profile at 0.3 years is the same in diagrams A and B.

1118

W. H. Casey and A. C. Lasaga

bonate at depths greater than w/k 5 10 cm is 45 mmoles/l of pore fluid. If this same profile begins at 15 cm, where G = 350 mmoles/l, rather than at the sediment surface, the concentration at depth is only about 10 mmoles/l. Rate coefficients vary considerably with time if they are incorrectly calculated by applying the steady-state model to the transient profiles. For the example presented in Fig. 10, the rate constant, k, and the concentration of reactive organic matter, G”, are within 50 percent of the true value after 25 years (Table 1). The time required to reestablish 95 percent of a steady-state profile is controlled by the reactivity of the sediment and the thickness of the zone of consumption. Much less time is required to establish steady state beneath a IO-cm thick zone than one extends to 15 cm. The thickness of the zone where solutes are consumed, however, also limits the accuracy with which rate coefficients can be calculated, even if there is adequate time to attain the new steady state. The curved portion of the solution profile, where concentration changes rapidly with depth, provides the optimum conditions for calculating the rate coefficients. This portion of the concentration profile lies at depths shallower than about (w/k). There is little resolution in calculating rate coefficients if the curved portion of the concentration profiles is destroyed by the production of acid in the unsaturated part of the soil. The resulting profile is nearly linear, and cannot be uniquely described by an exponential function. In the second simulation, the destruction of bicarbonate ion in the 0- 15 cm interval was halted after 0.3 years. After 0.3 years, bicarbonate is allowed to be replenished in the shallow soil by reaction diffusive transport upward from deeper pore fluids. For the example in Fig. lob, more than six years are required to reach 95 percent of the original steady-state profile. Rate coefficients calculated by fitting the transient profiles to Eqn. (5) would be within 50 percent of the true values only within about a year following the disturbance. The simulations provide a best-case analysis, in that we define the boundary to be sharp between zones of solute production and consumption in the soil. In the

marsh soil, the observed transition zone may vary, as the depth of the water table changes rapidly with time, and fluctuates to 25 cm depths and lower (CASEY.1985: MORRIS and WHITING, 1985; DACEY and HOWE& 1984; HOWESet al., 1987). One can see in Fig. 10 that fluctuations in the water table at shallow depths can affect the concentration profiles over a large interval in the sediment. These calculations show that it would be very difficult to obtain the true rate constant from the measured concentration profiles in cases where the term (w/k) in Eqns. (5) and (6) has dimensions similar to that of the fluctuating water table. The distribution of measured sulfate reduction rates indicate that the normalized length for reaction (w/k) in most marsh sediments is on the order of 10 cm or less (HOWARTH and TEAL, 1979; HOWESet al., 1984; HOWARTH and MERKEL, 1984; HOWARTH and GIBLIN, 1983; HoWARTHand MARINO, 1984). Because the concentration profiles are slow to return to steady state after being destroyed by the production of sulfuric acid in the unsaturated soil, using average profiles over time does not guarantee that the profiles are near steady state. For cases where bicarbonate is destroyed seasonally by fluctuations in the water table, the concentration profile may be transient for the entire year down to depths of 50 cm (Fig. lob). Rate coefficients modeled from these profiles are unlikely to resemble the coefficients needed to calculate sulfate reduction rates from the steady-state model (Eqns. (6) and (7)). CONCLUSIONS We combine our study of pore fluid chemistry with marsh-wide data on surface salinities and sediment hydrology, as well as numerical calculations which are appropriate for both vegetated and unvegetated marshes. Thus our conclusions can be generalized to salt marshes characterized by seasonal oscillations of pore fluid chemistry and average salinities which are higher than normal sea water. (1) Two processes account for the seasonal oscillations in chloride concentrations reported in some salt

Marsh sediment water modeling

marsh sediments: A) loss of water from within sediment pores; and B) the exchange of ions with overlying water when the marsh is inundated. The presence of average salinities, significantly higher than normal seawater values, which oscillate with the seasons, probably requires loss of water from the soil pores for at least part of the year. (2) Unsaturated conditions imply air entry into the soil to replace vacated pore space. This disrupts the nutrient concentration profiles so that they cannot be modeled in the same manner for all seasons. The periodic disruption of these profiles in this way distinguishes them from the profiles in subtidal sediments which are always saturated. The marsh soil is unsaturated for roughly 30% of the year at the LP site. During this time the water table (piezometric surface) lies at about 25 cm depth beneath the marsh surface. (3) The equations which exactly describe solute transport in marsh soil are non-steady state, and also non-linear. Pore water chemistry at shallow depths in the sediment is affected by varying diffusion coefficients, groundwater flow, and the loss of effective porosity. These combined effects give rise to an accumulation of solutes in the near-surface soil during summer. Models to explain solute transport in subtidal environments are not always appropriate for salt marshes. These models require inundated sediment, no groundwater flow, and are not accurate when the sediment becomes unsaturated. (4) The assumption that solute profiles are at a steady state, which may be necessary to calculate reaction rates, is not justified for many marsh soils. This assumption may be justified for marsh soils where the relative position of the zones of solute production and consumption have been stable for several decades, and where the zone of solute consumption is thin. In some cases, the concentration profile of those solutes produced during sulfate reduction, such as bicarbonate and ammonia, requires months to reestablish steady state after being disrupted. Acknowledgements-The authors would like to thank Drs. J. Dacey, H. Hemond, T. Church, A. Giblin, G. Holdren, R. Howarth, and C. Lord for their suggestions and comments. We particularly would like to thank Dr. H. L. Barnes whose comments improved both the clarity of this manuscript and our thoughts about the problem, as well as Drs. A. Guber and R. Slingerland. This project was supported by a Shell Fellowship to W. H. Casey. This work was supported in part by the U.S. Department of Energy under contract number DEAC04-76DPO0789.

Editorial handling: R. C. Aller

REFERENCES BARTBERGERC. E. (1976) Sediment sources and sedimentation rates: Chincoteague Bay, Maryland and Virginia. J. Sediment. Petrol. 44, 326-336. BEARJ. (1972) Dynamics ofFluids in Porous Media. Elsevier, New York. 764~. BERNERR. k. (lb80) Early diagenesis: A Theoretical Appreach. Princeton Univ. Press, 241~. BRESLERE., MCNEALB. L. and CARTERD. L. (1982) Saline and Sodic Soils. Springer-Verlag, New York, 129~.

Ill9

CARLSONP. R. JR. and FORESTJ. (1982) Uptake of dissolved

sulfide by Spartina altemiflora: evidence from natural sulfur isotope abundance ratios. Science 216,633-635. CASEYW. H. (1985) Modeling solute transport, sulfate reduction, and radionuclide migration in salt marsh sediments. Ph.D. thesis, Pennsylvania State University, 28 lp. CHAPMANW. J. (1960) Salt Marshes and Salt Deserts of the World. Plant Science Monographs, Interscience, New York. CLINEJ. D. (1973) Spectrophotometric determination of hydrogen sulfide in natural waters. Limnol. Oceanogr. 14,454457. DACEYJ. W. H. and HOWESB. L. (1984) Water uptake by roots controls water table movement and sediment oxidation in short Spartina marsh. Science 224,487-490. GARDNERW. R. (1959) Solution of flow equations for the drying of soils and other porous media. Soil Sci. Sot. Amer. Proc. 23, 183-187. HARRELLJ. W. and SAEEDM. (1980) The effect of moisture on phosphorous diffusion in coal mine spoils. Soil Sci. 129, 261-265.

HEMONDH. F. and FIFIELDJ. L. (1982) Subsurface flow in salt marsh peat: A mode1 and field study. Limnol. Oceanogr. 27, 126-136. HESSLEINR. H. (1976) An in-situ sampler for close interval pore water studies. Limnol. Oceanogr. 21, 9 12-9 14. HOLDRENG. R., BRICKER0. P. and MATISOFFG. ( 1975) A model for the control of dissolved manganese in the interstitial waters of Chesapeake Bay. In Marine Chemistry in the Coastal Environment (ed. T. M. CHURCH), pp. 36438 1. Amer. Chem. Sot. series no. 18. HOWARTHR. W. and TEAL J. M. (1979) Sulfate reduction in a New England salt marsh. Limnol. Ocennogr. 21,908914. HOWARTHR. W. and GIBLINA. E. (1983) Sulfate reduction in the salt marshes of Sapelo Island, Georgia. Limnol. Oceanogr. 28,10-82. HOWARTHR. W. and MARINOR. (1984) Sulfate reduction in salt marshes, with some comparisons to sulfate reduction in microbial mats. In Microbial Mats: Stomatolites. pp. 245-263. Alan R. Liss Inc., New York. HOWARTHR. W. and MERKEL S. (1984) Pyrite formation and the measurement of sulfate reduction in salt marsh sediments. Limnol. Oceanogr. 29, 598-608. Howls B. L., DACEYJ. W. H. and KINGG. M. (1984) Carbon flow through oxygen and sulfate reduction pathways in salt marsh sedime& Limnol. Oceanogr. 29, 1037- Id5 1. HOWESB. L.. DACEYJ. W. H. and GOEHRINGERD. t 1987) Factors co&rolling the growth form of Spurtina alte&7ora~~ Feedbacks between aboveground production, sediment oxidation, nitrogen and salinity. Limnol. Oceanogr. (in press). KING G. M., KLUG M. J., WIEGERTR. G. and CHALMERS A. G. (1982) Relation of soil water movement and sulfide concentration to Spartina alterniflora production in Georgia salt marsh. Science 218, 61-63. LASAGA,A. C. (1979) The treatment of multi-component diffusion and ion pairs in diagenetic fluxes. Amer. J. Sci. 279, 324-346. LI Y. H. and GREGORYS. (1974) Diffusion of ions in sea water and in deep-sea sediments. Geochim. Cosmochim. Acta 38, 703-7 14. LORD C. J. and CHURCHT. M. (1983) The geochemistry of salt marshes: sedimentary ion diffusion, sulfate reduction, and pyritization. Geochim. Cosmochim. Acta 47, 13811392. LUTHERG. S. III, GIBLINA. E., HOWARTHR. W. and RYANS R. A. (I 982) Pyrite and oxidized iron mineral phases formed from pyrite oxidation in salt marsh and estuarine sediments. Geochim. Cosmochim. Acta 46, 267 l-2676. MARTENSC. S. and BERNERR. A. (1974) Methane production in the interstitial waters of sulfate depleted marine sediments. Science 185, 1167-1169. MENDELSOHNI. A. and SENECAE. D. (1980) The influence of soil drainage on the growth of the salt marsh cordgrass

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TEAL .I. M. and KANWISHER J. W. (1970) Total energy balance in salt marsh grasses. Ecofonv 51,690-695.

WESTRICHJ. Tr (1983) The consequences and controls of bacterial sulfate reduction in marine sediments. Ph.D. thesis. Yale University, 53Op. WIEGERTR. G., CHALMERSA. G. and RANDERS~NP. F. (1983) Productivity gradients in salt marshes: The response of Spartina alterniflora to experimentally manipulated soil water movement. Oikos 41. 1-h APPENDIX The variation in salinity of surface water at the LP site was approximated by a series of the form:

J is the fraction of the year. A is the coefficient. Values of A are given below:

i

.1

I 2 3 4 5 6

8.49 23.44 -80.6 1 116.20 -59.74 Il.13