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Wetland Biogeochemistry Lab., Soil and Water Science Dep., Univ. of Florida, 106 Newell Hall, P.O. Box 110510, Gainesville, FL 32611
Corresponding author (Krr{at}gnv.ifas.ufl.edu)
Received for publication January 4, 2000.
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONMATERIALS AND METHODSRESULTSDISCUSSIONCONCLUSIONSREFERENCES |
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Abbreviations: BOD, biochemical oxygen demand • DRP, dissolved reactive phosphorus • SOD, soil oxygen demand
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTSDISCUSSIONCONCLUSIONSREFERENCES |
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Wetland soils are characterized by slow turnover of organic material due to a limited supply of terminal electron acceptors. This results in accumulation of organic matter and is one of the principal, long-term removal mechanisms of P and N in treatment wetlands (Kadlec, 1999). The mineralization of organic matter and subsequent nutrient release is predominately controlled by quality of organic material (Godshalk and Wetzel, 1978) and the supply of electron acceptors (D'Angelo and Reddy, 1994a,b). Thus, the supply of electron acceptors is one of the key determinants of organic matter decomposition and release of organically bound nutrients (McLatchey and Reddy, 1998). Thus, the rate of utilization of electron acceptors can indicate the rate of cycling of key nutrients, such as N and P.
Nutrient loading increases net accumulation of organic matter and associated nutrients in wetlands through accelerated primary productivity (Craft and Richardson, 1993; Reddy et al., 1993). However, increased loading not only increases total nutrient content of soils but also increases soluble forms, which can potentially be released into the water column (Reddy et al., 1998). This release of nutrients from soils becomes very critical and can maintain eutrophic conditions in wetlands, even when external loads are curtailed.
Several methods have been used to determine the magnitude and direction of nutrient flux from soils to the overlying water column. Deterministic models of solute flux, such as Fick's First Law of Diffusion, rely on close-interval estimation of porewater concentration gradients (Berner, 1980). These gradients are typically measured using finely sliced soil and sediment samples or by using porewater equilibration (dialysis) devices (D'Angelo and Reddy, 1994a; Moore et al., 1991, 1998). Phosphorus flux can also be determined using more direct approaches, such as by observing changes in water column concentrations inside in situ benthic chambers (Gomez-Parra and Forja 1993; Callender and Hammond, 1982; McCaffrey et al., 1980), or by measuring changes in floodwater concentrations in intact, incubated soil and sediment cores (Moore et al., 1998). The results from these approaches to flux measurement are often compared to derive some estimate of the variability between the techniques, and presumably attempt to determine a "best" measurement technique for a given system. Fickian diffusion has been compared with flux determined with incubated cores and benthic chambers in lakes (Sinke et al., 1990), open ocean (Devol, 1987), nearshore marine (Hopkinson, 1987), and estuarine (Callender and Hammond, 1982) ecosystems, but not in wetlands.
There has been a great deal of research on nutrient exchange dynamics between lake sediments and the overlying water column, but comparatively little research on the influence of wetland soils on water quality. Accordingly, we investigated several factors that are known to influence nutrient exchange processes in order to test the hypothesis that the supply of electron acceptors exerts a control on the rate of flux of P from soils. Additionally, we examined several techniques commonly used to estimate exchange rates. This study will have broad-ranging implications in other wetlands where nutrient exchange rates are of interest.
The objectives of this study were to (i) determine P flux from soil to overlying water column as measured by porewater equilibrators, in situ benthic chambers, and intact soil cores; (ii) determine the role electron acceptors play in governing P flux; and (iii) estimate internal P load to the water column.
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MATERIALS AND METHODS |
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TOPABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTSDISCUSSIONCONCLUSIONSREFERENCES |
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Experiment I: Phosphorus Flux under Ambient Floodwater Conditions
Floodwater samples of approximately 10 mL were removed from each core at 0, 1, 3, 5, 7, and 10 d, immediately filtered, acidified, and analyzed for dissolved reactive phosphorus (DRP). This resulted in the removal of approximately 60 mL of floodwater over the course of the experiment. In this experiment and all of the following, evaporative and sampling losses were measured over the 10-d experimental period and were accounted for in the determination of mass flux. The average incubation temperature was 28.5°C (±2.6).
Experiment II: Consumption of Electron Acceptors and Phosphorus Release
Following the conclusion of Experiment I, the floodwater was mechanically aerated to obtain an initial dissolved O2 content of approximately 6 mg L-1. The cores were then sealed with polyethylene piston-type closures. Dissolved O2 was then measured in each replicate core by inserting a dissolved O2 bottle electrode through a port in the piston. The volume to surface area ratio of 100 (L to m2) was sufficient to allow the O2 measurements to be made over a short time interval (within 4 to 12 h) (Bowman and Delfino, 1980). Oxygen consumption measurements were made in groups of three due to equipment limitations; however, the whole experiment was completed within 3 d. Even though there was a slight offset in time between the dissolved O2 determinations on the first and last sets of cores, this effect was minimal, since all environmental conditions were held constant.
Immediately following O2 consumption measurements, consumption of O2, NO3–N, and SO4–S were measured in the same cores. The cores were opened, aerated, and spiked with 900 mg of KNO3 to obtain a final NO3–N concentration of 9 mg L-1. This is well above ambient concentrations and thus the measured rates represent a potential maximum estimate of NO3 reduction under NO3 nonlimiting conditions. Ambient water column SO4–S in the cores was 34 (±16) mg L-1 at the initiation of the experiment, so spiking was not necessary. Oxygen consumption rates determined in this experiment are not reported, as only two measurements were obtained and the data from the continuous measurements described previously were believed more accurate. Floodwater samples (27 mL) were withdrawn for analysis at 0, 0.3, 1, 2, 4, 7, and 10 d. During the experimental period, floodwater dissolved O2 concentrations were reduced to <0.1 mg L-1, thus measured P flux represents conditions of O2 depletion and influence of alternate electron acceptors. The average incubation temperature was 25.5°C (±1.5).
In a separate experiment, soil oxygen demand (SOD) was measured on surface soil (0–10 cm) collected from Field Stations 1 through 8 along a nutrient enrichment gradient in WCA-2A (DeBusk et al., 1994) (Fig. 1). Approximately 5 g (wet weight) of soil was placed into a 250-mL biochemical oxygen demand (BOD) bottle placed on a magnetic stirrer. The BOD bottle was filled with deionized water and aerated for 5 min. A dissolved O2 probe was inserted into the bottle (YSI 5905 BOD probe, Yellow Springs Instruments, Yellow Springs, OH); the contents of the BOD bottle were continuously stirred during the 8 h incubation. Dissolved O2 measurements were made at 0, 0.5, 1, 2, 4, 6, and 8 h after initiation of the experiment. Incubations were made in the dark at 25°C. All measurements were made using three-replicate samples. The SOD rate was calculated as follows:
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Experiment III: Phosphorus Flux under Aerobic Floodwater Conditions
At the end of the Experiment II, the overlying floodwater from each core was siphoned off and immediately replaced with water collected from a nutrient-unaffected area of WCA-2A. This reflood water was coarse-filtered through Whatman qualitative #4 filters. The DRP concentration of the reflood water was 6 µg L-1. A 10-mL sample was withdrawn for analysis at 0, 7, 14, 21, 30, and 36 d. The approximate incubation temperature was 20°C. The conditions of this experiment were similar to Experiment I, except that the period of anoxia imposed in Experiment II presumably lead to the death of invertebrate populations. Thus, flux in this experiment represents nutrient exchange without the augmentation of solute transport due to invertebrates.
Experiment IV: Influence of Water-Level Drawdown and Reflooding on Phosphorus Flux
The soil cores from the previous experiment were then uncovered, drained of all floodwater, and exposed to air-drying for approximately 2 mo. Since the soil cores were relatively short (approx. 30 cm), all drying that occurred was a result of evaporation from the soil surface (i.e., no subsurface drawdown was used to expedite the drying process). Platinum-tipped redox electrodes were installed at a depth of 5 cm in one replication from each of the eight stations to approximately determine the extent of drying and oxidation. On average, the redox potential (Eh) increased from -29 (±237) to 215 (±81) mV during the drying period. After drying, soil cores were reflooded with filtered (Whatman #42) water from the nutrient-unaffected area within WCA-2A. The floodwater volume was adjusted such that each core contained 1 L (6 cm) of overlying water. The average incubation temperature was 26.0°C (±2.2).
In addition to the above described study, several intact soil cores measuring 10 cm in diameter were collected from Stations F1, F4, and U3 on 13 Mar. 1997 after a prolonged dry period in the field (Fig. 1). At the time of collection, the cores were devoid of floodwater. The cores were reflooded with 0.95 L of water collected from the field site on 20 Mar. 1997, resulting in a floodwater depth of 12 cm. The cores were incubated in the dark at 22°C. A 10-mL sample was withdrawn at 0, 1, 2, 4, 8, 15, 22, 29, and 39 d and analyzed for DRP, as described previously. Details concerning the sequence of experiments performed on the intact soil cores and the following field experiments are given in Fig. 2 .
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Three replicate chambers were placed in WCA-2A at Stations 1 (1.4 km from inflow), 3 (3.3 km from inflow), and 8 (10.2 km from the inflow) on 16 July 1996 and 4 Aug. 1997 (Fig. 1). The floodwater inside the chambers was spiked with 80 mg of N as KNO3 to obtain a floodwater concentration of approximately 1 mg L-1 NO3–N for the 1996 event. For the 1997 benthic chamber experiment, the chambers were spiked with 8 mg of N as KNO3 for a final water column concentration of 0.1 mg L-1 NO3–N. This concentration more closely approximated ambient NO3–N levels. Sulfate spiking was not needed, as the levels in the water column during the chamber experiment were in the range of 17 (±2) mg L-1. Dissolved O2, NO3–N, SO4–S, temperature, and pH were measured for a period of 24 h. The water surrounding the chambers was monitored for the same analytes as mentioned above. A Hydrolab (Austin, TX) Datasonde 3 was installed at each of the three stations to record ambient temperature, dissolved O2, and pH concurrent with the benthic chamber measurements. The average in situ temperature during the July 1996 deployment was 30.5°C (±1.1). During August 1997, the water column inside each chamber was mechanically aerated to obtain an initial dissolved O2 concentration of approximately 1.5 mg L-1 at Stations 1 and 3. Dissolved O2 levels at Station 8 were initially high, therefore aeration was not necessary. The average temperature during the August 1997 experiment was 28.7°C (±1.6). The recirculation pump was operated for approximately 1 min prior to each O2 determination in order to have sufficient water velocity for the oxygen electrode. Visual inspections were made to insure that sediment resuspension did not occur during pump operation.
Electron acceptor consumption and P flux measured in the soil cores and benthic chambers were calculated by determining the slope of the concentration vs. time curve through linear regression, then multiplying by the floodwater volume to soil surface area ratio of the soil core:
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Porewater Equilibrators
The porewater equilibrators used in this study were patterned after a device described by Hesslein (1976). Porewater equilibrators consist of 2- x 10- x 50-cm blocks of acrylic into which are milled 8-cm3 cells that are spaced vertically 1 cm apart. The sample cells were filled with deionized water and overlain with 0.2-µm pore diameter polyethersulfone membranes (Gelman Sciences, Ann Arbor, MI, Product no. XE 22061). The equilibrators were placed into acrylic containers, sealed, and purged of O2 with N2 gas. In the field, the equilibrators were removed from the containers and pushed 30 to 40 cm into the soil. They were then left in situ for 2 wk, allowing time for dissolved constituents in the sediment porewater to equilibrate with the deionized water inside the cells. Equilibrium studies have shown that 2 wk is sufficient time for porewater nutrients to equilibrate with the solution inside the equilibrator cells in sediments of high porosity (Carignan, 1984; Carignan et al., 1985). Diffusion modeling has also indicated that the time necessary to achieve 90% equilibration with porewater dissolved P, through molecular diffusion alone, is approximately 3 d (Webster et al., 1998). If the assumption is made that some resupply of soluble P occurs through exchange with the solid phase, 99% equilibration occurs in approximately 2 d (Harper et al., 1997). Since the effects of macrobenthos would likely shorten the equilibration period, 2 wk was deemed adequate for equilibration. After the equilibration period, the equilibrators were withdrawn from the soil and the cells were sampled by withdrawing the contents with a syringe. Samples were withdrawn from the cells at 1-cm intervals to a depth of 10 cm below the soil–water interface, and thereafter every 2 cm. The samples were stored at 4°C until analysis.
To determine temporal variability in porewater concentration gradients, porewater equilibrators were installed in WCA-2A at Stations 1 through 8 on 7 Feb., 8 May, and 11 Sept. 1996.
The concentration gradients were used to estimate the flux of P across the soil–water interface using Fick's First Law:
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= soil porosity (cm3 cm-3), Ds = bulk soil diffusion coefficient (cm2 s-1), 
C/Z = concentration gradient of component i with respect to depth, Z (cm), and 8.64 x 105 is a units conversion.
The average pH in the upper 10 cm of soil during the February 1996 equilibrator sampling at Stations 1 through 8 was 7.05 (±0.08). This was used to determine the diffusion coefficient of P in water. Since the speciation of soluble P is pH dependent, an interpolated value of Ds was used. Li and Gregory (1974) reported the diffusion coefficients of HPO2-4 and H2PO-4 in pure water as 7.34 x 10-6 and 8.46 x 10-6 cm2 s-1, respectively. An average of these two values was used in the calculation of diffusive flux, 7.9 x 10-6 cm2 s-1. The diffusion coefficient was modified for the restrictive effect of soil structure by dividing it by the square of the soil tortuosity. Soil tortuosity (
) was calculated from a relationship developed by Sweerts et al. (1991), or:
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is the soil porosity.
Analytical Methods
Dissolved O2 was measured with a YSI model 58 O2 meter (Yellow Springs Instrument Company) equipped with a YSI model 5730 stirring electrode using USEPA Method 360.1 (USEPA, 1979). Dissolved O2 and temperature were continuously recorded in the greenhouse experiments with a Campbell Scientific (Logan, UT) Model CR10 datalogger. Nitrate N and SO4–S were determined with a Dionex (Sunnyvale, CA) Series 4500i ion chromatograph, using EPA Method 300.0 (USEPA, 1979). Floodwater samples were analyzed for DRP using a Technicon (Tarrytown, NY) AutoAnalyzer, and EPA Method 365.1 (USEPA, 1979). The pH and Eh were measured with a Fisher Scientific (Pittsburgh, PA) Accumet Model 1002 pH meter.
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RESULTS |
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TOPABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTSDISCUSSIONCONCLUSIONSREFERENCES |
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Experiment III: Phosphorus Flux Under Aerobic Floodwater Conditions
Experiment III (aerobic P flux following anoxic electron acceptor experiment) was performed to determine if the initial DRP flux observed in Experiment I could be maintained at the same intensity and if asphyxiation of aerobic soil biota had any effect on P flux. The concentration of DRP in the floodwater of the Station 1 cores increased from approximately 172 to 332 µg L-1 over the 36-d experimental period, while cores taken from Station 2 increased from an average of 524 to 1196 µg L-1. The remaining stations showed very little or no increase in DRP. In general, all stations showed much lower P flux after reflooding, compared with the two previous experiments (Table 2).
Experiment IV: Influence of Water-Level Drawdown and Reflooding on Phosphorus Flux
At the end of the 60-d flooding period (Experiments I, II, and III), floodwater was removed and soil cores were allowed to dry for approximately 2 mo. This was followed by flooding for an additional period of 60 d. Dissolved P flux from these cores was lower in soil cores obtained from the affected area (2.3 km from inflow), as compared with flux measured in cores immediately after removal from the field (Experiments I and II) (Table 3).
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Field Experiments
Benthic Chambers
The mean initial water column DRP concentration at the initiation of the 1996 chamber experiment was 60 µg L-1, while the water column DRP concentration at the beginning of the 1997 experiment was considerably lower at approximately 15 µg L-1. The DRP concentration at Station 8 (10.1 km from inflow) generally remained below the limit of detection (ca. 2 µg L-1) throughout the July 1996 and the August 1997 experimental periods.
The average P flux from Stations 1 and 3 for the July 1996 experiment was 10 and 9 mg m-2 d-1, respectively, and was calculated using the first 4 h of data for Station 1 and the first 8 h of data for Station 3. Station 1 was the only station that showed any measurable P flux during the August 1997 experiment. Flux from this station was very similar to that measured in the 1996 experiment, 9.8 (±3.7) mg P m-2 d-1 (Table 4).
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The ambient dissolved O2 concentration was also low during the August 1997 experiment. However, an attempt was made to aerate the water column that succeeded in elevating the water column dissolved O2 to approximately 1.5 mg L-1. The consumption rates measured at all three stations were on the order of 100 mg m-2 d-1 using this approach.
Nitrate removal was rapid at Station 1 with almost complete disappearance within 8 h. The calculated NO3–N consumption rate was 274 mg m-2 d-1. Nitrate consumption at Station 3 was slower than at Station 1, with nearly all the NO3–N consumed in 24 h (Table 4). However, in the nutrient unaffected site (Station 8) approximately 39% of NO3–N added was still in the water column after the 24 h. Nitrate consumption in August 1997 was highest at Station 1 (124 mg m-2 d-1). Rates observed in the 1997 chamber experiment were approximately half those observed in 1996, probably due to the lower initial concentration used in the 1997 experiment.
There was very little change in SO4–S concentration inside the chambers during the 24 h period (data not shown). Ambient dissolved O2, pH, and temperature at each of the three stations were continuously recorded during the 24-h period that the benthic chambers were in place (Fig. 4c). Diel cycling in dissolved O2 and pH indicates increased primary productivity in the water column at Station 8, compared with Stations 1 and 3 (Fig. 6) .
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTSDISCUSSIONCONCLUSIONSREFERENCES |
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Dissolved P flux measured using porewater P concentration gradients was approximately an order of magnitude lower than P flux measured using benthic chambers and intact soil cores. Many researchers have performed comparative studies similar to this one in lakes and estuaries and have also noted disparities between calculated and observed flux (Callender and Hammond, 1982; Devol, 1987; Hopkinson, 1987); however, prior to this study, these comparisons had not been made in wetlands. For instance, Gomez-Parra and Forja (1993) compared flux based on sediment porewater concentrations gradients to benthic chamber P flux in the coastal waters of the southwest of Cadiz, Spain and found that in all cases the benthic chamber flux exceeded the gradient-derived flux, sometimes by as much as 29 times.
Though not directly measured in our study, the advection of porefluids out of the sediment due to the activity of invertebrates can result in a flux at least equal in magnitude to molecular diffusion alone (McCaffrey et al., 1980; Van Rees et al., 1996). In experiments conducted with labeled tracers, McCaffrey et al. (1980) were able to distinguish flux due to concentration gradients from flux caused by bioturbation. The combined flux compared favorably with flux measured with chambers alone. Callender and Hammond (1982) compared the calculated P flux, based on concentration gradients, with flux measurements in benthic chambers and found a significant "flux enhancement" or ratio of in situ to calculated flux. They concluded that nutrient fluxes measured in situ with benthic chambers may be 1 to 10 times higher than flux calculated from porewater nutrient profiles and the increased fluxes were attributed to irrigation of sediments by macrofauna. The authors found that field locations that had the lowest concentrations of macrofauna also had the best agreement between calculated and measured flux. Van Rees et al. (1996) investigated the effect of benthic invertebrates on diffusion coefficients in Lake Okeechobee, Florida and found that the diffusion coefficient varied by as much as 470% within a single sediment type. They attributed this range to the presence of benthic fauna. Several other factors, including temperature, dissolved O2 content of floodwater, rate of stirring, and presence of biota have also been shown to regulate P flux from sediments to the overlying water column (Holdren and Armstrong, 1980; Moore et al., 1998; Khalid and Patrick, 1974; Olila and Reddy, 1997).
Other factors that are difficult to estimate in deterministic models of P exchange with overlying floodwater include microgradients at the sediment–water interface (Berner, 1980; Krom et al., 1994; Davison et al., 1991) and variability in sediment physical characteristics. However, Sweerts et al. (1991) found that the ratio of the diffusion coefficient in water to the diffusion coefficient in sediment varied little between sediments of low and high porosity, indicating that this is a minor concern.
In Experiment II, the floodwater dissolved O2 was completely depleted, which probably decreased the activity of benthic invertebrates, thereby decreasing solute transport due to bioturbation. Rutgers et al. (1984) described a technique to discriminate between the flux of silica caused by molecular diffusion and the flux augmented by bioturbation. They found a dramatic decrease in silica flux as the dissolved oxygen content within the benthic chambers approached zero, suggesting a reduction in biotic activity. If the differences in the calculated and the observed P flux in the WCA-2A study were due to activity of the benthic organisms, then it should be possible to modify the diffusion coefficient to account for their activity.
This can be expressed as:
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Knowing the values for the aerobic and anaerobic flux and Ds allows the calculation of De for use in subsequent diffusion studies. Since the average ratio between the aerobic and anaerobic (asphyxiated) flux of P was approximately 7, a more realistic value for the soil diffusion coefficient (Ds) in these sediments may be closer to 3.78 x 10-5 cm2 s-1. This value considers both the tortuosity effects of soil structure and the increased solute transport presumably caused by benthic organisms. Further evidence of the contribution of benthic organisms to nutrient transport was found in Experiment IV. The anticipated result of soil drying was the increased mineralization of stored nutrients, resulting in greater P flux upon reflooding. That this was not observed possibly indicates a negative effect of soil drying on the local benthic community.
Porewater concentration of DRP, as measured with the equilibrators, was generally highest during February and lowest in September, with the exception of Station 5 (5.12 km from S-10C), which showed a reverse trend. This is in contrast to porewater profiles observed in WCA-2A in 1990 (Koch-Rose et al., 1994). In the 1990 study, porewater nutrient profiles were generally higher in July and lowest in May. The authors attributed this seasonal sequence to high plant uptake of porewater nutrients (and therefore depressed concentration gradients) in spring and high mineralization rates of organic matter in late summer, leading to elevated porewater nutrient profiles. Sinke et al. (1990) found a high correlation between porewater CH4, DRP, and NH4–N and also concluded that this is evidence that mineralization of native Netherlands peat soils determined porewater concentrations of these nutrients. Stations 6 through 8 showed no distinct seasonal trend in porewater P and remained at or near the detection limit.
Results presented in this study showed that the water column redox conditions and water-level drawdown can significantly influence P flux to the water column. This internal load of P becomes a critical factor in regulating eutrophication status of the wetland, once external loads are curtailed. Thus, identification of internal P loading is essential to determine the time required for recovery. High P flux (expressed as an average of all conditions) from soils affected by P loading suggests that these soils function as a source of P to the overlying water column (Fig. 8) . If we assume that approximately 25% of the total P in the top 30 cm of soil (Reddy et al., 1998) is potentially mobile and can diffuse at a rate of approximately 2 mg P m-2 d-1 into the overlying water column, the measured P flux would be sustained for a period of approximately 5 yr.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONMATERIALS AND METHODSRESULTSDISCUSSIONCONCLUSIONSREFERENCES |
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There was a gradient in O2 consumption with respect to distance from the source of P loading in the intact soil cores. Oxygen consumption at Station 1 (1.4 km from inflow) was 3.6 g m-2 d-1 and decreased to a constant 1.5 g m-2 d-1 approximately 4 km from the inflow. There was no apparent trend in NO3–N and SO4–S consumption as measured in intact soil cores with respect to distance from the inflow. Average NO3–N and SO4–S consumption rates were 120 (±81) and 130 (±52) mg m-2 d-1, respectively. Soil oxygen demand was correlated to P flux, while NO3 and SO4 consumption rates were not, indicating the dominant role that aerobic processes have on P mineralization in wetland soils. Phosphorus flux from intact soil cores after the electron acceptor experiment (Experiment II) was considerably lower than P flux in soil cores maintained aerobically, possibly due to the death of invertebrate organisms. Draining, drying, and reflooding of intact soil cores resulted in P flux that was intermediate between the flux observed in anaerobically (Experiment II) and aerobically (Experiment III) incubated soil cores. Intact soil cores retrieved from Station F1 (1.8 km) after field-drying showed P flux that was similar to P flux measured in intact soil cores from Station 1, or 5.2 mg m-2 d-1, on average. The remaining soil cores retrieved from Stations F4 (6.8 km) and U3 (11.0 km) showed little or no P flux. Results of this study suggest that the soils underlying WCA-2A will function as a source of P loading to the water column, even after external loads are reduced.
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ACKNOWLEDGMENTS |
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REFERENCES |
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