Chemical composition, pH, and redox state of sulfur and iron in complete vertical porewater profiles from two Sphagnum peat bogs, Jura Mountains, Switzerland

Geochimica et Cosmochimica Acta, Vol. 61, No. 6, pp. 1143-l 163, 1997 Copyright 0 1997 Elsevier Science Ltd

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Chemical composition, pH, and redox state of sulfur and iron in complete vertical porewater profiles from two Sphagnum peat bogs, Jura Mountains, Switzerland PHILIPP STEINMANN* and WILLIAM SHOTYK’ Geological Institute, University of Berne, Baltzerstrasse 1, CH-3012 Bern, Switzerland (Received July 6, 1995; accepted in revised form December

16, 1996)

Abstract-Complete porewater profiles from two peat bogs in the Jura Mountains were analysed for major and trace inorganic anions and cations. At La Tourbikre des Genevez (TGe) and Etang de la Grubre (EGr) , peat formation began approximately 5,000 and 10,000 years BP, respectively. The maximum depths of peat accumulation are 140 cm (TGe) and 650 cm (EGr) ; previous geochemical studies showed that the ombrogenic sections of the bogs extend to depths of approximately 20 cm (TGe) and 2.50 cm (EGr) Water samples were obtained using in situ diffusion equilibrium samplers (peepers), which allow filtered (0.2 pm) porewaters to be obtained while preventing degassing and oxidation. These samplers were found to be well suited to bog porewaters and allowed volatile (dissolved Con, acetate) and redox-sensitive species (HS -, Fe’+) to be quantified without further sample preparation or treatment. Aqueous species concentrations were determined immediately afterwards using ion chromatography with either conductivity (acetate, HCO;, Cl-, Br-, NO;, HPOi-, SO:-, Na+, NH:, K+, Mg*+, Ca*+), amperometry (HS-), or absorbance detection (Fe( III) and Fe(I1)). The comprehensive analyses of anions and cations allowed humic substances to be calculated by the difference in electrical charge balance (i.e., the anion deficit). Concentrations of total dissolved CO, (2- 12 mM) showed that carbonate equilibria play a significant role in the acid-base chemistry throughout the profiles. In near surface, ombrogenic porewaters with pH around 4, however, protons (approx. 160 peq/L) are contributed mainly by the dissociation of humic substances (2-7 mM DOC). In the deepest, minerogenic layers H2C03 is the predominant acid at both sites. At these depths, carbonate alkalinity (up to 3 meq/L at EGr, up to 8 meq/L at TGe) arises from reaction of the pore fluids with mineral matter in the underlying sediments. In the transition zone between the ombrogenic and minerogenic extremes, organic and inorganic acids are equal in importance. Unidentified organic S species accounted for 90-99% of total dissolved sulfur (ST ) in the porewaters at TGe, with SOi- and HS- the dominant inorganic species; S species with intermediate oxidation states such as SOi- and S@- were always less than the detection limit of approximately 0.4 PM. At TGe the sulfate concentrations exceeded those of sulfide, with 1.25 and 0.25 PM, respectively, being typical. At EGr, ST and SOi- were comparable to the waters at TGe, but HS- at EGr was always less than the detection limit of 0.15 PM. At both sites dissimilatory sulfate reduction is limited by the low concentrations of sulfate supplied to the bog surfaces (i.e., atmospheric deposition only), and the uptake of sulfate and its conversion to organic S compounds by the living plants. Despite the anoxic condition of the waters, the ratio of Fe(III), to Fe(II), was always high: at EGr this ratio was generally l:l, and even in the sulfidic waters at TGe the ratio was 1:3. PHREEQE was used to calculate the effect of complex-forming organic ligands on ( Fe3+ } and ( Fe’+ ] in these porewaters. The relatively high ratios of Fe(III), compared to Fe(II), can be explained in terms of the much greater thermodynamic stability of the organic complexes of Fe3* compared to those of Fe*+. Copyri@ 0 1997 Elsevier Science Ltd 1. INTRODUCTION

S (Carter, 1955) and the speciation of Al, Fe (Perdue et al., 1976), and other trace metals (Reuter and Perdue, 1977). There are four main peatland types: bogs, fens, swamps, and marshes (Gore, 1983). Fens, swamps, and marshes are minerogenic wetlands, that is, they receive their inorganic solids mainly from percolating groundwaters, and sometimes also, particularly in the case of marshes, from flooding surface waters. In contrast to minerogenic peatlands, the surface peat layers in ombrogenic bogs are hydrologically isolated from the influence of local groundwaters and surface waters, and inorganic solids are supplied exclusively by atmospheric deposition (Dau, 1823; Damman, 1986, 1987). As a result, bog surface waters are oligotrophic (Ramann, 1895; Kivinen, 1935; Malmer, 1986) with Mg and Ca concentrations in the porewaters up to ten times lower than rainwater values (Shotyk and Steinmann, 1994). Bog surface waters are also

The surface waters in peatlands exemplify highly coloured, organic rich natural waters. Globally, various kinds of peatlands occupy approximately 5% of the Earth’s land area (Kivinen and Pakarinen, 1981) Locally, however, they may be much more important, e.g., 30% of the land area of Finland and 17% of Canada are covered by peat. In many regions of the world, therefore, organic rich natural waters from peatlands may have a significant influence on the chemical composition of surface waters, especially the cycling of

*Present address: Institut F. A.-Fore], route de Suisse 10, CH 1290 Versoix, Switzerland (steinmanQsc2a.unige.ch). ‘Author to whom correspondence should be addressed (shotyk@ geo.unibe.ch).

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acidic (Kotilainen, 1927; Gorham et al., 1985) with pH values of 4 being typical. Below the peat surface, bog waters are anoxic (Hesselman, 1910; Malmstriim, 1923). These and other characteristics of bog waters are reviewed in more detail elsewhere (Gorham et al., 1985; Shotyk, 1987, 1989a,b). All peatland waters are characteristically rich in DOC, with typical concentrations in the range lo-50 mg/l; humic substances account for 70-90% of this DOC (Thurman, 1985). With approximately 15 beq of carboxylic and phenolit H per mg of C, dissolved humic substances are thought to control the pH of bog su$zce waters (Hemond, 1980; McKnight et al., 1985; Urban et al., 1986, 1989). However, peat bog porewaters are also rich in dissolved CO1, with concentrations commonly about one hundred times atmospheric pco, (Websky, 1864; Thompson et al., 1927; Shotyk, 1989b). Perhaps not surprisingly, there is a long history of controversy regarding the relative importance of organic vs. inorganic acids in these waters (see Shotyk, 1988, for a review). As peat accumulates over time, ombrogenic bogs may develop from the natural acidification of fens and swamps, especially in cool, humid climates which favour the rapid growth of Sphagnum and its subsequent accumulation as peat (Johnson, 1985; Crum, 1987; Succow, 1988; Gottlich, 1990). As a consequence, ombrogenic Sphagnum-dominated peats often overlie minerogenic Carex- and wooddominated fen and swamp peats, with resulting vertical geochemical gradients in pH (Waksman, 1930) and dissolved solids (Siegel and Glaser, 1987; Hill and Siegel, 1991; Shotyk et al., 1992). To date, however, the relative importance of organic acids vs. carbonic acid as pH controls in such profiles remains to be quantified. In addition, the redox state of these waters and changes in redox state with depth have not yet been investigated. One possible explanation for the paucity of reliable redox chemistry data for these waters may be the difficulties of collecting and analyzing the CO*charged, organic-rich, anoxic waters without altering their existing chemical state. One goal of this paper is to quantitatively evaluate the relative importance of dissolved humic substances and carbonic acid to the pH of peatland waters representing a vertical geochemical gradient from ombrogenic at the surface to minerogenic at depth. By studying complete vertical profiles, the contribution of organic acids and CO1 to the {H+ ) of the waters can be quantified from pH 4 at the surface of the bogs to pH 7 at the interface between peat and the underlying mineral sediment. To do this, porewater samples from complete vertical profiles were obtained from two contrasting Sphagnum bogs using in situ diffusion equilibrium samplers (peepers). These samplers allow filtered (0.2 pm) porewaters to be obtained while preventing degassing and oxidation and have been used successfully to collect porewaters in many kinds of wetlands and peatlands (King et al., 1982; Wheeler and Giller, 1984; Loiselle, 1986; Casey and Lasaga, 1987; Wieder and Lang, 1988; Tarutis et al., 1992; Marnette et al., 1993; Shannon and White, 1996). Peepers are well suited to bog porewaters, allowing volatile (dissolved CO*, acetate) and redox-sensitive species (HS _, Fe’+) to be quantified without further sample preparation or treatment

(Steinmann and Shotyk, 1996). A second goal of the paper is to evaluate the redox state of the porewaters with respect to S and Fe. The results given here represent the first quantitative analyses of the acid-base chemistry and redox state of porewaters from an ombrogenic peat bog. 2. MATERIALS

AND METHODS

2.1. Study Sites Both bogs are in the Franches Montagnes region of the Jura Mountains, in northwestern Switzerland (Fig. la). The Franches Montagnes consist of a calcareous plateau with low relief, characterized by the absence of rivers due to abundant karst phenomena. The average annual temperature in this area is 5.5”C, and precipitation exceeds 1300 mm per year. The major element geochemistry of the solid phase and the mineralogy of the two sites is discussed in Steinmann and Shotyk ( 1996).

a

b

/I

I-

NNW

100m

-’

SSE

TGe

z

1002

;

1000

ti E .P L

996 996 994 I

100m

’

Fig. 1. (a) Location (arrow) of the bogs at Etang de la Gruere (EGr) and La Tourbiere des Genevez (TGe) in the Franches Montagnes, Jura Mountains, Switzerland. (b) Vertical section TGe, after Welten ( 1964). Notice that the bog is separated from the calcareous bedrock by a layer of mineral sediments (clays and marl) which are elevated above the surrounding mineral soils. Not only is the bog elevated above the mineral soils, it is also cut off from contact with them by the dolines (karstic sink holes) which surround the bog. (c) Vertical section through the dome shaped peninsula at EGr constructed using data from Joray ( 1942). The small open triangles indicate the points where Joray ( 1942) drilled peat cores for macrofossil and pollen analysis. (b) and (c) : S indicates the position in which the samples were collected for the present study. The darker shading distinguishes the zone of strongly decomposed peat with low hydraulic conductivity from the overlying and underlying peat layers which are less decomposed (lighter shading). Arrows indicate the preferential directions of fluid flow.

Porewater 2.1.1.

profiles

from peat bogs in the Jura Mountains

La TourbiPre des Genevez

The peat bog La Tourbiere des Genevez (TGe) is at an elevation of 1020 m a.s.l., near the village of Tramelan. This bog has developed on an outcrop of Tertiary clays and marls (up to 1 m thick) which is elevated above and surrounded by a ring of karstic sink holes (Fig. lb). Penetration of calcareous water from the surrounding area into the bog, therefore, is prevented by ( 1) the generally flat topography of the area, (2) the elevation of the bog above the surrounding terrain, and (3) the presence of the dolines surrounding the bog (Gerber and Monbaron, 1990). This 7.2 ha elliptical bog extends no more than 590 m from west to east and 170 m from north to south. TGe has a pronounced domed structure: from north to south the increase in elevation from the edge of the peatland to the centre of the dome is 2.5 m (Fig. lb). Porewaters at this site were sampled at the apex in the center of the dome. The stratigraphy at the sampling site is as follows: low ash (2% by peats down to 40 cm; 40-70 cm weight) Sphagnum dominated Sphagnum Eriuphorum peats with 3-5% ash; 70-140 cm Curex dominated peat. The ash content increases steadily from 40 cm to reach approximately 7% at a depth of 120 cm. The degree of decomposition of the peat as expressed on the von Post scale (von Post and Granlund, 1925; Clymo. 1983; Landva et al., 1986; Hobbs, 1986) varies as follows through the profile: O-25 cm (H3), 25-50 cm (H4), 50-70 cm (H5), 70-100 cm (H4), 100-140 cm (H5H6). At 140 cm the mineral sediment is reached, consisting of essentially of 60% clay, 30% quartz, <5% each of plagioclase feldspar and potassium feldspar, ~2% calcite and dolomite. A brownish to greenish peaty clay (at 140-160 cm) with ash content ranging from 75% to 85% is followed by a greyish to greenish clayey silt containing roots and yellowish spots (160-200 cm).

2.1.2.

Etung de la GruPre

The peat bog Etang de la Gruere ( EGr) is located just 4 km west of TGe and occupies 22.5 ha. The samples were taken in the central part of the bog where it is strongly domed (Fig. lc) and peat accumulation is more than 6 m (Joray, 1942). The stratigraphy at the sampling site is: Sphagnum dominated peats down to 60 cm; Sphagnum Eriophorum peats from 60-250 cm; Sphagnum dominated peats again from 250-420 cm; Carex dominated peat from 370-640 cm; at 640 cm there is a sharp boundary with the underlying gray silty clay. The ash content of the peats is typically less than 2% down to a depth of about 3 m. From there downward the ash content increases to reach 9% at a depth of 6 m. The profile has a strongly decomposed middle part with less decomposed bottom and top; O-25 cm (H2), 25-60 cm (H4), 60-130 cm (H5), 130-420 cm (H8), 420-550 cm (H4). 550-640 cm (H3). Like TGe. this bog is also strongly domed; based on the elevations measured by Joray (1942), the site where the bog was cored for the present study is 4 m higher than the edge of the bog, a distance of approximately 300 m. EGr has developed on Oxfordian clays and marls consisting essentially of 40% clay and 40% quartz, with <5% of each of plagioclase feldspar and potassium feldspar, and approximately 3% calcite and 2% dolomite (Steinmann, 1995). The topography of the site combined with the presence of a 2 m thick layer of highly decomposed ( H8) peat suggests that the upper 2-3 m of this profile is beyond the influence of laterally penetrating waters. 2.1.3.

Hydrological construints 011jluidflow

There have been no quantitative hydrological studies undertaken at either TGe or EGr, and in general little is known about the hydrology of raised bogs (Ingram, 1982; Armstrong, 1995). Nevertheless, the peculiar morphology of these two sites and their development within a karst plateau allow a general description of the hydrologic situation to be drawn. The two bogs represent independent, isolated peatlands developed on small islands of siliceous sediments surrounded by a karstic plateau. Due to the low relief and abundant karst phenomena in the surrounding areas, there is no opportunity for outside surface waters to penetrate those sections of the bogs which are above ground level. In addition, the surface peat layers in each bog are clearly elevated heyond the surrounding mineral soils, up to 3 m in the case of TGe

1145

and 4 m at EGr. Upward movement of water within these sections of the peat profiles is not possible because the flat topography provides no hydraulic head for artesian type of water flow. The water level in the artificial lake surrounding EGr and the water level in the karstic sink holes around TGe show that maximum possible upward extent of groundwaters; the surface layers of the bogs are approximately 3 m (TGe) and 4 m (EGr) above these water levels. As Ingram ( 1982) has shown, upward movement of waters driven by capillary forces is insignificant in raised bogs. The porewater profiles taken in these sections of the bogs (i.e., samples which are elevated above the surrounding mineral soils and above the local water table), therefore. represent fluids whose inorganic solids were supplied exclusively by atmospheric precipitation (rainwater and dry deposition). The interpretation of the ombrogenic status of the surface layers in these two bogs has been confirmed by independent geochemical studies of the major and trace elements in the peats and porewaters, relative to average composition of rainwater in this area (Shotyk and Steinmann, 1994; Shotyk, 1996a,b). In the uppermost layers, the dry bulk density varies considerably with depth in both peat cores, first increasing with depth, then decreasing (Fig. 2a,b). Boelter (1969) showed that there is a close relationship between peat bulk density and hydraulic conductivity. Using Boelter’s relationship hydraulic conductivities between lo-* and IO-’ cm s-’ are obtained for the uppermost peat layers, whereas in the denser subsurface peat the hydraulic conductivity is only lO-4 cm s-r. Thus, in each of these bogs there is a thick zone (20-30 cm) of well decomposed peat characterized by a hydraulic conductivity which is up to one hundred times lower than the surface layer. As a result, in each of these bogs these zones of reduced permeability represent an impediment to vertical fluid flow, and most of the water which is supplied to the surface of the bogs is expected to leave either by evapotranspiration, or by migrating horizontally through the more permeable surface layer; fluid flow vertically downward is expected to be negligible. In the uppermost, poorly decomposed layers at EGr (first 30 cm or so) where the hydraulic conductivity is relatively high (log K = -2.5 cm/s), water may percolate vertically downward only to depths of 30-40 cm; from this depth, most of the fluid flow will be laterally away from the center of the dome. Below 130 cm at EGr there is a layer (ca. 130-420 cm) of highly decomposed (H8) peat. According to Baden and Eggelsmann ( 1963), the hydraulic conductivity of H8 Sphagnum is ten to one hundred times less than poorly decomposed (H3) peat. This thick layer of well decomposed (H8) peat will, therefore, act as a barrier to vertical fluid flow. Hoag and Price ( 1995) showed that in such well decomposed peats the hydraulic conductivities may be as low as IO-’ cm SK’, and that diffusion may become the dominant transport mechanism. The very large differences in the hydraulic properties of these peats can be used as a first approximation to constrain possible water flow patterns (Fig. 1b,c). In the lower part of the profile at EGr, below this hydraulic barrier ( 130-420 cm) of very well decomposed (H8) peat, the peat is again less decomposed (H4) and is, therefore, much more permeable. In this basal section of the bog lateral infiltration of water into the sampled profile may be important because of the gentle slope of the underlying basal clays and marls. These fluids, however, are not derived from outside of the peatland. Instead, they are derived from reaction of rainwater with the shallow peat layers surrounding the bog (i.e., the lagg zone). These basal, minerogenic porewaters can reasonably be expected to be influenced by reaction with the underlying clays and marls. The situation in the deeper layers at TGe is similar to that described for EGr, but on a smaller scale. Downward and lateral flow is most likely in the less decomposed surface layers. Flow in the more decomposed intermediate layer (below 40 cm) will once again be restricted. In the basal peat layers with somewhat higher hydraulic conductivity (H4 from 70 to 100 cm), lateral flow may be important. However, because this basal layer is elevated above the surrounding soils and cut off by the karstic sink holes, penetration of groundwater or surface water from outside the bog is impossible. Thus, the chemical composition of the pore fluids in the deeper layers at TGe result from the interaction of the rainwater, peat bog porewater, and the porewaters of the mineral sediments underlying the peat. 2.2. Sample

Collection

The porewater samples analyzed in this study were obtained using peepers. A detailed discussion of the sampling procedures is given

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and W. Shotyk

a

60-

0

20 40 60

Stratigraphy

0.1

0.2

80 100

ash (%)

P (g cmm3)

b

t

a ’ B t

100

5

Stratigraphy

10

15

20

ash (%)

0

3

6

ash (%)

9

0

0.1

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P (g cm")

Fig. 2. Stratigraphy (dominant botanical composition: S, Sphagnum; E, Eriophorum; C, Cm-a; and degree of humification, von Post scale), ash content, bulk density TGe (a), and EGr (b). The bulk density profile can be used to infer an estimated hydraulic conductivity (log K) as discussed in the text.

elsewhere (Steinmann and Shotyk, 1996) but briefly summarized here. Peepers consist of a Plexiglas housing made up of individual 30 mL chambers that were filled in the laboratory with deionized, deaerated water (18 MR) and covered with a 0.2 pm membrane filter (HT-Tuffryn). The chambers were inserted into the bogs at varying depths and were allowed to equilibrate for 4-6 weeks with the porewaters. Two types of peepers were used: one allowing porewaters to be sampled in the shallow, near surface zone with a resolution of 2.5 cm, the other one for collecting waters from deeper layers at 60 cm intervals. To prevent oxidation during sample collection and handling, the peepers were pulled directly from the bogs into N2 filled glove bags. Individual chambers were then sampled directly through the glove bag using syringes. The samples were brought to the laboratory in the closed syringes which were kept in a cold storage bag and then analyzed immediately.

Membrane type and equilibration times as a function of membrane type and peeper design have been evaluated in a detailed study by Brand1 and Hanselmann ( 1991). The equilibration time chosen for the present study was based on both their results and from own equilibration experiments in the lab. Equilibration experiments in the lab using deionized water with inorganic ions in concentrations similar to those found in the porewaters the peepers equilibrated within days. A longer equilibration time and membranes with a relatively large pore size (0.2 pm instead of a dialysis membrane) were used in order to allow also larger organic molecules and associated metal ions to enter the sample chambers. Leaching experiments (using either deionized water or dilute HCl solution with pH 4) showed that no contamination occurred from the peepers, membranes, and syringes used. As noted earlier, peepers have been used successfully to collect

Porewater profiles from peat bogs in the Jura Mountains porewaters in many kinds of wetlands and peatlands (King et al., 1982; Wheeler and Giller, 1984; Loiselle, 1986; Casey and Lasaga, 1987; Wieder and Lang, 1988; Tarutis et al., 1992; Mamette et al., 1993; Shannon and White, 1996). In order to test whether the peepers worked well in the Jura bogs, we compared samples obtained using peepers with samples obtained by squeezing slices of peat (Steinmann and Shotyk, 1996). In general, the concentrations of conservative species such as Cl and the major element cations were comparable with both methods, indicating that the equilibration time chosen was adequate. However, a great advantage of the porewaters obtained using peepers was that they were filtered to 0.2 pm in situ, and that no further preparation or treatment was needed prior to analyzing them. More importantly, the porewaters which had been expressed from the peat cores and subsequently filtered in the lab yielded SOi- concentrations which were a factor of 10 or more times higher, and PO:- concentrations which were five to ten times lower, probably because of oxidation and degassing of the squeezed porewaters (Steinmann and Shotyk, 1996). The squeezed porewaters also had much higher concentrations of Br- and NO;, and lower concentrations of Fe, than the peeper samples. To quantify redoxsensitive species such as S and Fe, therefore, the peepers have a clear advantage over other sampling techniques. Thus, the peepers were chosen for sampling the bog porewaters while preventing the risk of degassing or oxidizing the samples. Using peepers, a total of eleven porewater profiles were sampled at TGe and 6 at EGr. The profiles considered in this paper were taken on the following dates: at TGe, T12, October 1994; Tll, June 1994; T9, July 1993; T8, July 1993; T3, December 1992; T2, August 1991; at EGr, E9, October 1994; E8, July 1994; E7, June 1994; E6, June 1994; E3, June 1993. 2.3. Chemical Analyses Total dissolved Na, K, Mg, Ca, Fe, Al, P, S, and Si were determined by ICP-AES in a set of subsamples which had been acidified and frozen upon returning from the field to the lab. Independent studies have used ICP spectrometry successfully for measuring total S (Hordijk et al., 1989). Major anions, major cations, sulfide, and Fe( II)/Fe( III) were analysed using ion chromatography (IC). Complete details of the IC methods are given elsewhere (Steinmann and Shotyk, 1995a-c) but are briefly summarized here. The anions F- , acetate, formate, Cl-, NO;, Br-, NO;, PO:-, SO:-, S,O:- , and HCO, were analyzed using a Dionex AG4A-SC guard column followed by an AS4A-SC separator column, an ASRS I suppressor, and a Dionex conductivity detector. The samples were injected with the same syringes in which they were collected in the field. This approach ensured that the samples were always kept in closed containers, except for a very short time when they were open to the NZ atmosphere, and degassing was, therefore, minimal. One advantage of the procedures used for sample collection, handling, and analysis is that the concentration of total dissolved CO, (Cr ) can be determined as HCO; together with the other anions in less than 12 min, in a single run (Steinmann and Shotyk, 1995a). The IC method allowed HCO; to be measured in the concentration range 0.7-14 mM. All injections were made through an in line Dionex OnGuard P cartridge to remove humic substances which otherwise foul the analytical columns. Subsamples for sulfide measurements were collected in the field in separate syringes containing base in order to preserve sulfide in the nonvolatile HS form. The sulfide measurements were performed using a Dionex Carbopac PA-l guard column as the separator column and a Dionex amperometric detector with a Ag working electrode (Steinmann and Shotyk, 1995b). The retention time for HS was approximately 1.2min, and the detection limit 0.15 PM. Measurements of sulfide in the porewaters included both the free, dissolved sulfide (HS-), as well as acid volatile sulfide (AVS). AVS in the porewaters was obtained by acidifying the samples and trapping the released H,S gas (Steinmann and Shotyk, 1995b). From these results, the possible importance of sulfide complexed by metals was found to be quantitatively unimportant in the bog porewaters. With respect to ferric and ferrous Fe, the traditional approach to measuring Fe?+ and Fe’+ m geological samples is to measure Fe’+ calorimetrically, total Fe by atomic spectroscopy, and to calculate Fe3+ by difference (Fritz and Popp, 1985; Stucki and Komadel,

1147

1988). Here, ferric and ferrous Fe were separated on a Dionex CS5 separator column using a pyridine-2,6-dicarboxylic acid (PDCA) eluant; the retention times were approximately 5 and 10 min, respectively. After leaving the column the metals were complexed by 4( Z-pyridylazo)resorcinol (PAR) and the absorbance of the coloured complexes measured at 520 nm using a Dionex UV/Vis detector. Detection limits were approximately 90 and 180 nM, respectively. The samples for Fe analyses were injected as collected in the field without any further treatment. This prevented sample alteration and gave precisions (RSD) for both Fe(II1) and Fe(I1) measurements of approximately 10% (Steinmann and Shotyk, 1995~). The IC data were compared with ICP analyses of the same samples and the sum of Fe( III) and Fe(I1) measured by IC on average was 90% of ICP results (r’ = 0.999). Cations (Na+, NH;, K’, Mg*+, and Ca*+) were measured using a Dionex CS-12 column with CSRS-I suppressor and conductivity detector (Steinmann and Shotyk, 1995a). Subsamples for cation measurements were acidified prior to injection through an OnGuardP cartridge in order to liberate any of these metals which may have been complexed by organic matter. Comparison of these measurements with ICP analyses of the same samples (Steinmann and Shotyk, 1995a) showed that the IC method yielded complete recoveries. For Mg and Ca, for example, [Mg-ICP] = 0.97 [Mg-IC] - 0.03 ppm, r2 = 0.998 (n = 21); [Ca-ICP] = 1.01 [Ca-IC] - 0.07 ppm, rZ = 0.998 (n = 21). The pH was measured on a 2-3 mL aliquot of samples using a Ross sure flow electrode designed for measuring dilute solutions in combination with a Metrohm pH meter. Measurements were performed with stirring at room temperature. Prior to measuring the pH value, the samples had been kept in the closed syringes. After measuring the fresh pH value, the samples were degassed by purging with N2 gas for 10 min and then the degassed pH was measured. A maximum deviation of 0.02 pH units was observed when measuring the pH 4 buffer as blind standards, indicating good performance of the electrode. The standard deviation of me precision of the pH measurements is estimated to 0.01 pH units. However, degassing of CO2 from the samples with the highest pc,,, could offset the pH by 0.1 pH unit or more as is discussed below. Dissolved organic carbon (DOC) in the porewater was measured with a Shimadzu TOC 500 Analyzer. The determination limit is around 2 ppm and the relative standard deviation of triplicate measurements was typically < 1%. 2.4. Charge Balance Calculations Accurate measurements of all important cations and anions allowed the organic anionic charge [X-l to be calculated as the difference in electrical charge (peq/L): [X-l = C( inorganic cationic charge) - Z( inorganic anionic charge + acetate)

( 1)

The organic anions formate, acetate, and oxalate were measured but only acetate was found in important amounts. Acetate was included as a separate species in the calculations; thus [X-l mainly represents large molecular weight organic (humic) anionic charge. PHREEQE (Parkhurst et al., 1990) was used to do these calculations, in order to correctly consider all protons associated with inorganic acid anions, and all possible HCO; and OH- ligands bound to metals. PHREEQE has the option to calculate the distribution of aqueous species while balancing the charge by adding a species chosen by the user. This mode was used to balance the electrical charge thereby yielding [X-l directly. All calculations were done for 25°C. The sources of error in the calculation of [X-l include the analytical errors involved in anion and cation analyses and the systematic errors in pH measurement of samples with high pco,. The sum of these errors is reasonably small (2-5%) as long as X- is the prevailing anion; this is true for the porewater samples in the upper sections of both profiles. In contrast, the error inherent in this calculation was unacceptably high in the deepest samples of both profile E9 and T12, where HCO; is the dominant anion; in these cases, [X-l is calculated as a small difference between two large numbers and its error increases considerably. In fact, in deeper porewaters with

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higher concentrations of cations and inorganic anions, the error associated with the calculation increases to 50% or more. In the deeper porewater samples, therefore, the organic anionic charge cannot reliably be determined using the charge balance approach. As a consequence, to calculate the distribution of aqueous species in these samples using PHREEQE, the organic anionic charge was estimated from the product of the measured DOC of the samples and the average charge density (see below). A second problem which becomes important for those samples in which HCO; is the dominant anion is the following. Measuring the fresh pH value of samples with very high pco, yields pH values that are slightly too high because of inevitable degassing of samples during or prior to the pH measurement. In contrast, C, in these samples was measured from aliquots which were injected directly from the sampling syringe into the chromatograph and in these cases no degassing occurred. Thus, the measured fresh pH may be not exactly in equilibrium with the measured CT if PC., is very high. Overestimation of the fresh pH leads to an overestimation of HCO; (as a free anion and as a ligand bound to metals) and, therefore, to an underestimation of [X-l. As noted above, for samples in the pH range 4-5 where X _ was the dominant anion, the electrical charge due to the organic anions (,ueq/l) could be calculated reliably as the anion deficit in fresh porewaters. This value could then divided by the measured DOC concentration (mgll) to yield the average density of electrical charge of the dissolved humic substances (,ueq/mg C). In order estimate the electrical charge density for DOC at higher pH values, those samples of profile T12 with X- as the dominant anion were degassed, thereby increasing the pH. Calculating the charge balance of these degassed samples (using degassed pH and CT ) provided an average charge density for the DOC over the pH range 4-7 (Fig. 3). A similar approach has been used to model organic-rich lake waters (Driscoll and Lethinen, 1994). To calculate metal speciation in the deepest porewaters of the profiles T12 and E9 using PHREEQE, therefore, the organic anionic charge was calculated as the product of the measured DOC concentration and the average charge density of DOC at the appropriate pH and included in the input dataset. 2.5. Calculation of Metal Speciation

4

5

6

0

9

PH

Fig. 4. Typical pH profile of bog porewaters at EGr. The porewaters were collected using peepers which meant that they were filtered to 0.2 pm in situ (Profile E6). The pH was measured in the lab before ( pHfresh ) and after sparging the samples with N, gas ( pHdegassed) The difference between these two profiles clearly shows the importance of dissolved CO, to the pH of these waters.

these porewaters. This hypothetical organic ligand resembles oxalate (which is not an important species in bog porewaters) with respect to metal binding properties but has a different acid-base chemistry. Because the concentration of the unprotonated ligand was obtained from the electroneutrality condition (charge balance), there was no need to make any assumptions regarding the acid-base properties of the L*- ligand. The PHREEQE database was extended to include the metal complexation data for oxalate given by Hummel ( 1992). The hypothetical bivalent ligand Lz- was assigned the following log K values for 1: 1 complexes, taken from the corresponding 1:1 metal oxalate complexes CaL (3.19), MgL (3.42), KL

To calculate the speciation of Al and Fe from the measured concentrations of Al,, FeT, Fe(lll), and Fe(ll), the relevant stability constants for humic complexes of A13+, Fe3+, and Fe’+ were needed. Many of the published binding constants for bidentate 1:l metal humate complexes agree reasonably well with the corresponding 1: 1 metal oxalate complexes. Thus a hypothetical, oxalate like bidentate ligand (L*-) was selected to model the complexation reactions taking place between the organic ligands and the dissolved metals in

7

(0.9), Fe(ll)L Fe(lll)L+

(4.1),

(14.1), AlL+ (7.6)

(2)

The log K values used for Ca, Mg, and Fe(l1) 1:l complexes agree well with the corresponding values for metal fulvic ligand complexes given in the GEOCHEM database (Fulvl, Parker et al., 1990). The value for Al is comparable to the log K given for 1: 1 organic ligand Al complexes reported by Driscoll and Lethinen ( 1994). The value for Fe3+ was estimated to be ten orders of magnitudes higher than the value for Fe*+ according to Fe(lll)/Fe(ll) measurements presented in this paper (see below). In an effort to explain the relatively high ratio of Fe( Ill) to Fe( II) in these anaerobic waters, the differential binding of Fe’+ and Fe*+ by the organic anions in the porewaters was evaluated by including the binding constants for these metals with citrate and EDTA (Hummet, 1992) and salicylate (Parker et al., 1990) in the PHREEQE data base. The PHREEQE database was also modified to include both Fe’+ and Fe3+ as master species; this allowed the measured Fe( lll)/Fe(ll) concentrations to be used to calculate pe. PHREEQE was also used to calculate the pco, and evaluate the saturation state of the waters with respect to the predominant primary and secondary mineral phases; for this latter calculation, the concentration of SiT as determined using ICP spectroscopy was included in the input tiles. 3. RESULTS

21

4

4.5

5

5.5

6

6.5

7

7.5

PH Fig. 3. Charge density as a function of pH for organic rich bog porewaters. Charge densities were calculated using degassed porewater samples from TGe (T12, samples l-20, see Appendix). Error bars: 2 1 standard deviation.

3.1. pH Profiles The pH of most samples increased markedly after degassing. In the deep profiles at EGr, for example, the pH of degassed samples were typically 2 or more pH units higher than the fresh samples (Fig. 4). Measurements of CT in degassed samples revealed CT concentrations (measured as

1149

Porewater profiles from peat bogs in the Jura Mountains

t -------

E 3o

zz. Q) 0

F

E

00

50

-

70

---

O

0

-o----I

6

I

6

2.

E ‘L 0” 40 ---

--I

601

--I

I

10

s,,/c,,

I

12

I

0

I

0.5

14

r.?

I

1.5

1

2

Concentration (PM)

(xl 000)

Fig. 5. Ratio of organic sulfur/organic carbon measured in the porewaters at TGe. Organic S decreases with depth (from around 17 PM to about 12 FM) while DOC increases. (Profiles T8, T9).

Fig. 6. Sulfide S and sulfate S in the porewaters at TGe (profile T2).

in the deeper sections of the EGr profiles down to 6 m, with an apparent minimum at intermediate depths of 2-3 m. HCO, using IC) below the detection limit of approximately 0.15 n-M. In contrast, fresh porewaters contained 2- 12 mM Cr. The pH profiles and CT before and after degassing shows clearly the importance of dissolved CO1 for the pH of these waters. Not only is the partial pressure of CO, far from equilibrium with atmospheric CO1 in the deep porewaters, but elevated concentrations were also characteristic of the shallow porewaters. For example at TGe, at depths of 1517 cm, log pco, was as high as -0.85 atm (profiles T12, Tll); at a depth of 29 cm at EGr, log pco, = -1.11 (E6) and at 15 cm log pco, = -1.23 (E7).

3.3. Iron(III)-Iron In the porewaters at TGe, Fe, was generally 5- 10 or more times higher than the waters at EGr (Fig. 7). A typical iron(III)/(II) profile at TGe (Tll ) shows Fe, increasing with depth from approximately 10 to 35 PM Fe, (Fig. 7a). In seven TGe profiles analyzed for Fe( III)/Fe( II), the trivalent redox state was always found in substantial concentrations with a typical Fe( III)/Fe( II) ratio being approximately 1:3. In the first 3-4 m at EGr, Fer is a factor of ten times less than in the porewaters at TGe. A profile of Fe(III)/Fe(II)

3.2. Sulfur Species The inorganic sulfur species sulfate, sulfite, thiosulfate, free sulfide, as well as acid volatile sulfide typically accounted for less than lo%, and sometimes less than 1% of total dissolved sulfur (ST ) . The dissolved S in these porewaters, therefore, is mainly in an organic form. If humic materials account for most of the organic matter in the waters (Thurman, 1985), then the S is most likely associated with humic matter. A plot of the ratio of organic S/organic C vs. depth shows that the dissolved humic substances near the surface at TGe are much richer in S compared with the DOC in deeper layers (Fig. 5). Inorganic sulfur in the porewaters was present only in small amounts as sulfide and sulfate. Other S oxyanions such as sulfite or thiosulfate were never detected (d.1. approx. 0.4 ,uM) . Figure 6 is a profile of sulfate S and sulfide S measured at TGe in Aug. 199 1 (profile T2 ) . Similar sulfide concentrations were measured in Dec. 91 (T3), while no measurable sulfide (d.1. 0.15 PM) was detected in profiles collected on three other dates. Sulfate concentrations measured in profiles T4 through T12 were typically below 0.5 /.LMwith higher values ( 1-2 PM) only found in some of the uppermost samples. The general picture, therefore, was the predominance of dissolved organic sulfur ( 15 -25 /.LM) and the presence of low concentrations of both sulfide and sulfate. At EGr sulfide was never detected. Sulfate concentrations and Sr, however, were similar to those found at TGe. Small concentrations of sulfate (0.1-0.3 PM) were still detectable

10

I

I

I

I

I

a g s 5 $ n

_

30

50

20

10

30

Concentration (uAUl

0

1

2

c

I

1

3

4

5

I

6

7

6

Concentration @M) Fig. 7. (a) Total Fe( III) and total Fe( II) in porewaters from TGe (profile T8). (b) Total Fe(II1) and total Fe(R) in porewaters from EGr (profile E6 )

1150

P. Steinmann and W. Shotyk

from 115 to 595 cm at EGr shows that the proportions of these two species is nearly the same (Fig. 7b). Two profiles collected from near-surface layers (E8 and E3) also show high concentrations of Fe(II1) compared to Fe(I1). In general, the concentration ratio Fe( III)/Fe( II) in these porewaters is approximately 1.

3.4. Charge Balance and Speciation Calculations for a Profile at TGe The comprehensive analyses of all important cationic and anionic species in a porewater profile at TGe (T12, given in the Appendix) allow the calculation of charge balance and a quantitative assessment of the acid-base chemistry of the waters. The sampling interval of this profile was 2.5 cm at depths from 12.5 to 60 cm and every 21 cm from 77 to 182 cm. Descending from the surface, the porewater profile is characterized by an increase in pH, bicarbonate, and alkalinity (Fig. 8). The increase in bicarbonate is due to the combined effect of increasing pH and increasing CT. A detailed charge balance calculated using PHREEQE is given in Fig. 9. In the uppermost 35 cm of the profile, with pH in the range 4.1-4.3, organic anions by far contribute most of the negative charge. In this interval H+ is an important cation, and H+ together with Ca2+ dominate the cationic charge. At depths of 77 cm and below, the pH values are much higher (5.4-6.5) and, the anionic charge is completely dominated by HCO; . Free protons become less important with increasing depth, while Ca*+ contributes most of the positive charge. Iron occurs as organic complexes in the upper part of the profile, whereas free Fe*+ is the most important Fe species in the bottom half of the profile. In the upper sections of the profile inorganic complexation of metals is quantitatively unimportant. Due to the increasing concentrations of bicarbonate and metal cations with depth, however, there is a corresponding decrease in the ratios of DOC/HCO, and DOG/metals, and Me HCO; complexes become important in deeper sections of the profile.

3.5. Charge Balance and Speciation Calculations for a Profile at EGr A summary of the acid base chemistry and the charge balance for a deep profile at EGr (E9, given in the Appendix) is shown in Figs. 10 and 11. The anionic charge is dominated by organic anions only in the first meter of the profile where the pH values are in the range 4.2-4.3. In this section of the bog, H+ is the dominant cation. From approximately 1 m to 4 m, bicarbonate accounts for nearly one-half of the anionic charge, and in the basal layers it becomes the most important anion and Ca2+ the dominant cation; in the deepest layers H+ largely disappears. In contrast to the TGe profile, at EGr ammonium plays an important role in the middle part of the profile, where it is by far the single most important cation. 4.

4.1. pH Regulation in Peat Bog Porewaters As seen in the charge balances given in Figs. 9 and 11, the main sources of H+ in these waters are dissolved humic substances and carbonic acid: these two acids control the pH of peat bog porewaters. This finding supports the work of Thompson et al. (1927) who showed that the low pH of peat bog surface waters were maintained either by dissolved organic matter (in the case of “dry bogs” with pH < 4.2) or by dissolved carbon dioxide (in the case of “wet bogs” with pH > 4.2). The contribution of protons by each of these two acids is given by the concentrations of [ HCO;] and [X -1, respectively. In the uppermost part of profile T12 ( 12-40 cm) the relative contribution of [X -1, defined as X% = lOO[X-]/([HCO;]

[Fe,1

I

5

670

5 (W

lo- 1

+ [X-l)

(3)

ranges from 81% to 91% (Fig. 9). From 40 cm down to 98 cm the relative contribution of [X -1 declines continually to a value of about 50%, despite an increase in its absolute concentration; this is due to a disproportionate rise in [HCO;]. Below 98 cm [HCO;] continues to increase,

[HCO,’ 1

4

DISCUSSION

!Y

I

-0.75 -0.5

0.2 0.4

btm)

bedI)

[Cal

0

250 CPM)

500

1.5

3

MM)

[Fe,,,,,] and [Cal in a porewater profile from TGe (T12). Fig. 8. PH, lHCO?I, log PC~,. [X-l, [HCO;] and [X-] were calculated using PHREEQE as described in the text. A complete summary of the chemical composition of this porewater profile is given in the Appendix.

Porewater

profiles

from peat bogs in the Jura Mountains

anions

1151

cations ?? HCO,(-Me)

15

0

? ? L2-

35

EI

55

“’

HCOj (free ligand)

L(-Me)

?? so;* ? ? Cl ? ? H,PO,

77 98

?? H’

119

N

NH,

0

Fe”

? ? K+

140

6J

161

Na’

? ? Mg2’ ? ?Ca2’

182 -100

-50

0

50

H

MeHCO,

B

Me (-L)

100

charge (%) Fig. 9. Charge balance for the porewater profile T12 (TGe) calculated using PHREEQE. Me refers to Ca”, Mg*+, and Fe’+. A complete summary of the chemical composition of this porewater profile is given in the Appendix.

while the share of [X -1 decreases to approximately 2% (Fig. 9). At EGr the relative importance of [X-l ranges from 80% to 90% in the first meter of the profile (Fig. 11). In another near surface profile at EGr (E7) from 22 cm to 60 cm [X-l was responsible for 85% to 95% of protons released to the solution. Although [ HCO ;] in these cases contributes only 5- 15% of [H+] , a significant rise in pH following degassing was observed in all samples (e.g., from 4.07 to 4.13). Between 1 and 4 m the relative contribution of [X -1 is only 50%, and below 4 m only 15-20% (Fig. 11). The concentration of free H+ in these porewaters is the sum of acid-producing and acid-consuming reactions in the fluids. The most important acid-producing reactions are the dissociation of organic acids and H&O?. Counteracting these processes are the acid-consuming reactions which in-

elude biological sulfate and nitrate reduction, mineralisation of organic N compounds (ammonification), and mineral dissolution. The vertical pH profiles in these two bogs (in general, pH 4 at the surface increasing to pH 7 at depth) indicate that acid-producing reactions dominate in the uppermost, ombrogenic layers while acid-consuming reactions dominate in the deeper, minerogenic layers. By isolating specific acidproducing and acid-consuming reactions, the importance of individual processes can be quantified. To do this using the profile at EGr, it is helpful to separate the bog vertically into three compartments: the porewaters in the uppermost part of the bog which are derived directly from rainwater (the ombrogenic zone), the porewaters from basal layers which are influenced by reaction with the underlying mineral sediments (the minerogenic zone) and a transition zone between these two extremes.

I

r

I

[Cal

I\ [Fe,1

1 (mW

2

3-1.5

-1 (atm)

-0.5

0.2

0.4 0.6

(mN)

10

(IIN)

20

30

0.5

1

(mM)

Fig. 10. pH, [HCO;I, log PCO?, WI, [Fe,,,,J, and [Cal in a porewater profile from EGr (E9). [HCO; ] and [X-l were calculated using PHREEQE as described in the text. A complete summary of the chemical composition of this porewater profile is given in the Appendix.

P. Steinmann and W. Shotyk

1152

anions

cations ?? HC03(-Me) E! ? ?HCO; ? ?L*- (free ligand) El

L(-Me) AC‘ (acetate)

EJ

z 235 S c 295 ‘L 8 355

Cl

?? Ii+ H

NH,

? ?Fe*’ ?? K’ i8

Nat

•I

Mg2

? ?Ca2’ ? ?MeHCO, IxI

-100

-50

0

50

Me(-L)

100’

charge (%) Fig. 11. Charge balance for the porewater profile E9 (EGr) calculated using PHREEQE. Me refers to Ca’+, Mg*+, and Fe*+. A complete summary of the chemical composition of this porewater profile is given in the Appendix.

4.1.1. Acid-base reactions in the ombrogenic zone

A summary of the average chemical composition of rainwater from this part of Switzerland and the porewaters in the ombrogenic zone of EGr are given in Table 1. Also given are the relevant chemical reactions which are needed to explain the evolution from rainwater to bog porewater in the ombrogenic zone. The higher chlorine content of ombrogenie porewater relative to rainwater can be explained by evaporative concentration of the solution. The composition of such an evaporated rainwater is given in Table 1; it was calculated by increasing the concentrations of all ions except (H’) and (HCO;) by the same factor as Cl (assuming constant pco? during evaporation). Table 1 shows that SOi- concentrations in the porewaters are much lower than the corresponding average concentration in rainwater. The most likely sinks for sulfate are sulfide and organic sulfur. While sulfite and thiosulfate are conceivable products also, recent studies of S transformations in anoxic sediments have shown that these intermediates are mainly the products of sulfide oxidation, and not sulfate reduction (Luther et al., 1986; Jorgensen, 1990; Fossing and Jorgensen, 1990; Elsgaard and Jorgensen, 1992; Thamdrup et al., 1994). In the porewaters at EGr and TGe both SO:- and S@- were always less than the detection limit of approximately 0.4 PM. Polysulfide species can be ignored as they are unimportant at pH values below 6 (Morse et al., 1987). As noted earlier, almost all of the S in the porewaters is in the organic matter fraction. Therefore, the formation of organic S compounds is by far the most important sink for sulfate, while the reduction of sulfate to sulfide is negligible. The reaction which describes the transformation of sulfate to reduced organic sulfur can be written as SO:- + 2CH,O + 2H + = ( HPS)org + 2C02 + 2HzO (4) which indicates that 2 moles of H’ are consumed per mole of sulfur reduced. An additional fraction of sulfate may have

been lost as volatile, reduced gases, but this would have the same effect on the number of moles of H+ consumed SO:- + KHZ0

+ 2H+ = S(CH,),(g)

+ 3C02 + 3H20

(5)

For the purpose of the calculation, therefore, the specific product of sulfate reduction is not important, since all important sinks of sulfate in these porewaters consume two moles of protons for each mole of sulfate ions. Nitrate is a major anion in rainwater but is absent in the ombrogenic porewater (Table 1). There are various reactions that can account for the quantitative removal of nitrate. The reduction of nitrate to nitrite is not important here because nitrite concentrations in the porewaters are very low ( <0.4 PM), and more importantly because this reaction does not affect H+ Nitrogen is a growth limiting element in bogs (Small, 1972; Urban and Eisenreich, 1988), thus assimilation of N by plants and microorganisms is the most likely sink for nitrate; this reaction consumes one mole of H+ per mole of nitrate removed and can be written as NOT + 2CHz0 + H+ = (NH,),,,

+ 2C02 + Hz0

(6)

Gaseous species such as N2 and N20 were not measured as part of the present study, even though these are also possible products of NO; reduction. However, the production of these species also consumes one mole of H+ per mole of NO; reduced and, therefore, the net effect of these reductions on H+ would be the same. While it is not possible from the data presented here to determine which specific reactions are involved in NO; reduction, from the analytical results which are available and the approach which has been taken to evaluate this data, it is in fact possible to quantify the importance of nitrate reduction as an H+-consuming reaction. Acid producing reactions include the dissociation of organic acids and carbonic acid, and the conversion of NH: in rainwater inputs to organic N in the peat; this latter reaction produces one mole of protons for each mole of NH,+ converted.

Porewater profiles from peat bogs in the Jura Mountains

Il.53

Table 1: Average composition of rainwater (column l), ombrogenic pore water (co1 4) a& extent of acid-consuming and acid-producing reactions needed during the genesis of ombq ;enic pore water (co1 3) at EGr. 1 2 spenes wnwater evaporated smk/source of porewater rainwater EGr 50 cm H+

PM

PM

8.7

12.5

0.0 a

12.5

22.4

32.1

-63.8 b

0.2

28.5

41.0

-41.0 c

0.0

0.0

0.0

3.6 ’

3.6

AC-

0.0

0.0

16.9 e

16.9

HCOj

0.4

0.3

19.7 f

20.0

X-

0.0

0.0

95.0 g

95.0

H+

11.0

15.5

-43.4 h

58.9

Na+

6.1

8.7

-10.0 i

18.7

NH;

41.8

60.1

23.2 j

36.8

K+

2.3

3.3

-0.5 i

3.8

Mg2+

1.6

2.4

-1.0 i

2.9

Ca2+

9.0

12.9

5.8 ’

10.0

Fe2+'3+

0.0

0.0

-4.5 k

1.8

lt

bohnce

0.0

wdl

0.0

0.0

PM

0.0

4 Cl- is assumed to be conservative and used to correct for evaporation

“)SO~‘+2c~0+2Hf=

(H.$)Wg+2C02+2~0

c)Noj+2C%O+H+=

(W&W~ + 2 CO, + Hz0

d, H2m; + H+ = (H$Q,)ors e)HAc= AC‘ + H’ (AC = acetate) OH,CO,

=

HCOj+H+

g)HX = x- +H+

(X- = humic anion)

h) negative values indicate lower pH in the final solution 0 exchange: 1 mole Hf for each mole (Na’, K’); 2 moles H’ for each mole (Mg’+, Ca”) j) NH; = (NH&m

+ H+

The total amount of H’ consumed in the ombrogenic zone is approximately 164 peq L -’ (Table 1) Approximately 40% of this is due to sulfate reduction and 25% to nitrate reduction. By far the most important acid producing reaction is organic acid dissociation: this explains almost 70% of the H’ produced, with acetate responsible for 15% of this value. The remaining H+ produced can be attributed to the dissociation of H2COx and the conversion of NH: to organic N. 4.1.2. Acid-base chemistry of the transition zone The changes in acid base chemistry which take place in this zone can be calculated by comparing the average chemical composition of the porewaters in the ombrogenic zone (e.g., 50 cm, given in Table 2) with the transition zone; this is the zone between the ombrogenic and the minerogenic zones and is represented in chemical composition by the

porewater at 355 cm in Table 2. This transition zone corresponds to the more decomposed layer of the peat profile. Again, the relative importance of each of the reactions discussed above was calculated and the results shown in column 3 of Table 2. In these deeper layers of the bog NH: is the dominant cation and ammonification as the dominant sink for H+, accounting for 70% of the total (615 PeqlL) of protons consumed. As is seen from the ammonium profile in the Appendix, some part of the NH: may have been introduced to section of the bog by upward diffusion in response to the NH: concentration gradient, and in exchange for H+ diffusing downward. From these results, therefore, it is not possible to precisely identify the depths at which active ammonification is taking place. A considerable sink for H+ in the middle part of the profile arises from the dissolution of Ca*+ from minerals and/or cation exchange of H+ for Ca *+ Calcium ions from these processes

1154

P. Steinmann and W. Shotyk Table 2: Extent of acid-u~nsumin~and acid-pmducingteaclions (colomn 2) during uaters (cd 3) fmm near surface porn waters (co1 1). develoPme”t Column 4 shows off”““d”pg” the extent o actd-base textion neededfor the generation minemaophic porn waters (cd 5) at EGr species

1

1. 2 stwsoluceof H+

s%E PM

l.wl

cl-

12.5

-2.3

SOi-

0.2

-0.4 b

NO;

0.0

I

3 Pomwat= 355cmEGr

I

PM

’

0.0

I

4 ~somceof H+ w/l

5 porcwat= 535cmEGr PM

’

10.2

-0.0

0.0

0.0

10.2 0.0

0.0

0.0

0.0

Hz%

3.6

-3.4 d

0.1

3.2 d

3.3

Ac-

16.9

-6.3 =

10.6

9.3 =

19.9

Hco;

20.0

281.8 f

301.8

2154.4 f

2456.2

X-

95.0

287.0 g

382.0

-74.0 m

308.0

H’

58.9

45.7 h

13.2

11.6h

Na+

18.7

-0.5

i

19.2

-20.0

NH;

36.8

-432.3 n

469.1

K+

3.8

-7.1

-27.3

2.9

-24.2

15.0

-173.3

ca2+

10.0

i i -132.0 i

11.6

Mg*+

76.0

-1447.7

Fc2+“+

1.8

-5.3 k

3.9

0.0

0.0

bolom

0.0

1.6

i

39.2

-385.5 n

854.6

i i i

101.6

-50.8 ’ 0.0

38.9

799.9 24.2 0.0

1)Cl‘ may be lost due to incoqxxation into organic matter m)mineralizationof humic material “) (NH&,re + H+ = NH; For the other foomotes see Table 1.

consume 22% of the H+ (Table 2). Again, the depths at which calcite dissolution is taking place may well be lower in the profile, the Ca*+ ions being introduced by upward diffusion (see concentration gradient in the Appendix). The dominant sources of H+ in the transition zone are the dissociation of carbonic and organic acids and these are about equal in importance. 4.1.3. Acid-base

chemistry

of the minerogenic

zone

In a third step the chemical composition of the porewaters in the transition zone (e.g., 355 cm at EGr) can be compared with the porewaters in the minerogenic zone (e.g., 535 cm at EGr) . It is assumed that the porewaters in the minerogenic zone are the product of reactions which have taken place between fluids having a starting composition comparable to those of the transition zone with the underlying mineral sediments. Support for this assumption is the fact that penetration of the peatlands by external waters is prevented by the surrounding karstic, highly permeable bedrocks (see Materials and Methods). The results of the calculations are presented in column 4 of Table 2. Bicarbonate is by far the dominant anion in the deep porewaters. While the dissociation of carbonic acid is, therefore, the main source of protons, this is nearly balanced by the H+ consumed via mineral dissolution and ammonification. In particular, the release of base cations from the weathering of minerals in the underlying sediments (mainly Ca2+ and Mg*‘), is responsible for 77% of all H+ consumed. Like Ca, Fe also shows increased concentrations with increasing depth in the profile (Fig. 7a,b, Appendix). The

consumption of protons per mole of Fe dissolved depends on the final oxidation state of the dissolved Fe ions. In the footnotes to Table 1 a possible dissolution reaction for iron hydroxides is given which produces equal amounts of divalent and trivalent Fe and consumes 2.5 moles of protons per mole of Fe dissolved. Again, much of the Fe mineral dissolution may take place in the underlying sediments with Fe species diffusing upward in exchange for protons diffusing downward. Table 2 shows that the mineralization of dissolved organic matter (reaction m in footnotes to the table) is an additional sink of protons in the minerogenic layers, resulting in a decrease of the organic anionic charge despite increasing pH. Mineralization of dissolved organic matter is also reflected by the decrease in DOC in the lowermost peat layers (see Appendix). 4.2. Speciation of S and Fe in Bog Porewaters 4.2.1. Sulfate, suljide, and the apparent redox potential Sulfate-reducing bacteria are obligate anaerobes (Bemer, 1984), and the presence of measurable concentrations of sulfide in the waters at TGe show that the waters are strictly anoxic (Bemer, 198 1) . Given the pH values and the measured concentrations of sulfate, and sulfide, the redox potential of the porewaters can be calculated using (Stumm and Morgan, 198 1) pe = 41/8 + l/8 log ((SO:]/ where {

{ H2So}) - IO/8 pH

(7)

} indicates activities. This approach has recently

Porewater profiles from peat bogs in the Jura Mountains

been used to quantify the redox potential in other organicrich freshwater sedimentary environments (Marnette et al., 1993 ) . At depths below 60 cm at TGe, with a pH of approximately 4 and the concentrations of sulfate and sulfide almost equal, pe = 0. In other porewater profiles from this bog, sulfide was less than the limit of detection (0.15 PM) at the same pH. However, the equation given above shows that a tenfold change in the sulfate/sulfide ratio only slightly affects the calculated redox potential. Slightly more positive redox potentials further up in the profile are indicated by lower sulfide and higher sulfate concentrations. Unfortunately, the uppermost part of the profile is the most difficult zone to collect porewaters using peepers because of the seasonal changes in depth to water table. At higher pH values deeper in the profile, the pe is lower (e.g., - 1.1 at pH 5). We, therefore, estimate that the pe of these waters over the pH range 4-6 varies from 0 to approximately -3. Very similar redox potentials have been calculated using sulfate/ sulfide concentrations in other bog porewaters (Shotyk et al., 1992). At EGr, the concentrations of sulfate were similar to those at TGe, but sulfide at EGr was less than the limit of detection. Again, however, the calculated pe of the waters is not very sensitive to the sulfate/sulfide ratio. We, therefore, assume that the sulfate/sulfide ratio at EGr in not more than a factor of ten times that of TGe, and that the pe values at EGr, therefore, are similar to those at TGe. 4.2.2. Fe( III), Fe(II), and the apparent redox potential Iron-reducing bacteria also are obligate anaerobes (Lovely, 1987) and the presence of relatively high concentrations of Fe’* in these waters is consistent with the interpretation that the bog porewaters are anoxic. The presence of measurable concentrations of Fe(II1) in these same waters at first seems inconsistent with the Fe( II) results in the sense that the Fe(II1) is suggestive of a more oxygenated condition. Perhaps even more surprising, not only is ferric Fe present, the measured concentrations of Fe( III)T and Fe( II)T at TGe are in the ratio of 1:3, and at EGr in the ratio 1: 1. In other words, in these anoxic waters, both oxidation states of Fe are nearly equally important. One possible explanation for the relatively high concentrations of Fe(II1) in anoxic waters is complexation of Fe3+ by organic ligands. The high stability of complexes of Fe( III) by organic ligands such as humic substances is known from a number of studies (Senesi et al., 1977; Rausaet al., 1994; De Haan et al., 1990; Ephraim and Marinsky, 1990). The ability of many natural organic ligands to stabilize the trivalent oxidation state of iron is discussed in detail by Luther et al. ( 1992). The relative importance of iron species at TGe (profile T12) is shown in Fig. 12. Divalent iron occurs as free Fe’+ and as organic complex, while Fe3+ is present only as organic complexes. The increasing importance of bicarbonate as a ligand for Fe( II) in the deeper peat layers are due to increasing CT and pH. As a guide to the concentrations of Fe3+ which might be expected in an anoxic porewater at pH 4 with organic ligands absent, the ratio of the thermodynamic activities of Fe3+ and Fe’+ can be calculated as follows (Stumm and Morgan, 1981):

z

77

s AZ E

96

$

119

1155

140 161 0

20

40

60

60

100

molalities of iron species (“72)

Fig. 12. Important Fe-species in the porewater profile of TGe (T12). Calculated using PHREEQE.

pe = 13.01 + log ((Fe’+ ]/ [Fe*+ ))

(8)

Liang et al. (1993) used a similar approach to calculate the pe of groundwaters. Recall that the activity ratio of ( SO: }/ ( H2So } indicated a pe value of approximately 0 at pH 4. Using this pe value, the predicted activity ratio is log ( Fe3+ }/ ( Fe’+ } = - 13. In other words, the activity ratio of ( Fe3+ }/ ( Fe*+ } calculated using the pe value obtained from (SO:}/ ( H2So} is more than twelve orders of magnitude lower than the measured Fe(III)r/Fe(II).r. It is of course well known that in many natural waters there exists a lack of chemical equilibrium between redox couples (Morris and Stumm, 1967) because of slow kinetics or independent biogeochemical cycling. In groundwaters, for example, the lack of agreement between the redox potentials indicated by ( Fe3+ }/ ( Fe’+ } and ( SO: ) / ( H2So ) is well documented and ranges from 6 to 8 pe units (Lindberg and Runnels, 1984). The predicted activity ratio of ( Fe3+ ) / ( Fe’+ ) given above ( 10-13) is not intended to suggest that the two redox couples (S and Fe) in bog waters were expected to be in equilibrium. Instead, this calculation is simply used as a guide to indicate what concentrations of Fe3+ might have been expected in an anoxic porewater at pH 4 at pe 0 in the absence of complex-forming organic ligands. Given the high concentrations of dissolved organic carbon in these porewaters (2-7 mM), it is reasonable to enquire whether organic complexation of Fe(II1) may have reduced the thermodynamic activity of Fe’+ sufficiently to achieve an ( Fe3’ ]/ (Fe*+ } activity ratio comparable to the one predicted for pH 4, pe 0. In fact, for organic complexation to result in such a low ratio of ( Fe3+ } / ( Fe2+ } , the binding constants for the dissolved organic materials with Fe”+ would have to be comparable to those of oxinate and hem derivates which are among the most stable Fe( III) complexes known (Burgess, 1989), and the affinity of the organic ligand for Fe’+ would have to be approximately 10 I3 times greater than that for Fe*+. PHREEQE was used to evaluate the effect of a number of well characterized organic ligands on the activity ratio of free ( Fe3+ ]/ (Fe*+ ) using the measured ratios of Fe(III)/ Fe( II). The calculations were done as described earlier, except that the organic ligand to balance the charge was varied. Some typical ( Fe3+ )/ (Fe*+ ) ratios obtained from these

P. Steinmann and W. Shotyk

1156

Table 3: Dcpmdenccofth~~alcUlatcd activity ratio (~)/(I+)

fd~~tkfypc~~~@c

C!alc~lati~ns fa a pat waterramplehornTGe@asurcd tetal Fe@)Ipco = 3.0;depth20cm;pH= 4.05,T12)weredoneusing PHREEQE. ligd8

chosen.

organic ligmd

log((Fd+)l{Fe+)

PC

ozsfhte

-0.51 -2.98 -4.53 -6.09 -1.76

L= (hypot&ic)

-10.21

2.8

FJJTA

-11.98

1.03

no spwialion considued ilmgmic ligatKls only

salicylfue

ciuatc

calculations are shown in Table 3.Calculating the speciation using oxalate as an organic ligand (instead of the hypothetical organic ligand L*-) the resulting activity ratio log ({Fe’+ )/ {Fe’+ )) = -7.8. Using EDTA whose binding constants are 27.6 and 16 for Fe(II1) and Fe(II), respectively, the calculated ratio (log ( Fe3+ }/ { Fe*+ } = - 11.98) is comparable to the expected ratio (log ( Fe3+ }/ {Fe*+ } = - 13 ) at pH 4, pe 0. The results summarized in Table 3 indicate that preferential complexation of Fe3+ over Fe*’ by the dissolved organic matter in these waters could produce activity ratios of { Fe3+ ) / ( Fe*+ ) which are many orders of magnitude lower than the measured ratios of Fe(III)r and Fe( 1I)r. 4.3. Fe( III), Fe( II), and Organic Complexation in Organic-Rich Natural Waters

of Fe

The speciation of Fe for profile T12 discussed above is shown in Fig. 12. In both profiles (T12 and E9) organic complexes are the dominant species. In deeper layers, especially at TGe, FeHC03 complexes become more important as CT increases and DOC decreases (see Appendix). The association between Fe and organic matter in natural waters was established well over a century ago (e.g., Hunt, 1865 ) . More recently fractionation studies have shown that most of the Fe in the waters from lake sediments (Krom and Sholkovitz, 1978) and bog lakes (Koenings, 1976) is complexed by organic matter. Even in seawater where both Fe and organic ligands are present only in trace concentrations, most of the Fe is organically bound (Gledhill and van den Berg, 1994; Wu and Luther, 1995). Using competitive ligand complexation techniques to study these processes, the log K values for organic complexes of Fe3+ in seawater were found to range from approximately 21 (Gledhill and van den Berg, 1994) to 23 (Wu and Luther, 1995), values which are comparable to those of EDTA which was used in the calculations described above. Lewis et al. (1995) examined the thermodynamic stability of Fe(II1) complexes of catecholate-type siderophores extracted from marine bacteria and reported log K values in the range 37.6-43.6, values which are significantly higher than those reported for hydroxamate siderophores (log K = 29-32.5). Not only do siderophores form extremely strong complexes with Fe( III), with stability constants ranging up to log K-53 (Wu and Luther, 1995), they form much less stable complexes with Fe(II) (Lewis et al., 1995).

12.5 10.03 8.48 6.92 5.25

The thermodynamics and kinetics governing the solubility of Fe and the speciation of ferrous and ferric iron in solutions devoid of organic ligands can only be applied with great caution to systems containing appreciable concentrations of dissolved organic ligands (Stumm and Lee, 1960; Hem, 1972). In addition to the direct effects of organic complexation of Fe in solution, organic ligands may promote the rates of dissolution of iron oxides (LaKind and Stone, 1989; Deiana et al., 1995). Such ligands accomplish this by adsorbing to the hydr ( oxide) surface and exchanging electrons with Fe(III), with those forming inner sphere complexes especially efficient (Sulzberger et al., 1989). This process has been discussed in detail by Sulzberger et al. (1989) who note that organic ligands may promote the rates of both reductive and nonreductive dissolution of iron oxides. Luther et al. (1992) have proposed a comprehensive biogeochemical cycle for Fe in salt-marsh sediments and porewaters; within this cycle organic ligands promote the reductive dissolution of Fe( III) minerals **Fe’+(LL)

+ FeOOH + H+ s **Fe3+(LL)

+ Fe*+(LL) + 20H-

(9)

where LL refers to multidentate organic ligands and the asterisks indicate the original Fe( II) in solution. Due to complexation by the organic ligand which favors the trivalant oxidation state of Fe the dissolved Fe( II) acts as an electron donor in the reductive dissolution of the Fe( III) mineral. The resultant organic Fe( III) species may in salt marsh waters be reduced subsequently by sulfide. In the bog waters, however, sulfide concentrations are very low, and this may help explain why Fe(II1) organic complexes are relatively stable. Reduction of Fe(II1) by the organic complex itself is also suggested in the Fe cycle of salt marsh porewaters. It seems, therefore, that some type of organic ligands can stabilize Fe(II1) while others tend to reduce it. Luther et al. (1996) found that Fe(II1) generally accounted for < 10% of total dissolved Fe in anoxic salt marsh sediments. They found that most of the organically-bound Fe(III) was in the relatively large (100-5000) molecular weight class compared to Fe(I1) which was mainly in the
Porewater profiles from peat bogs in the Jura Mountains mechanisms. In the surface layers, dissolved Fe(II1) production can be sustained via Fe(I1) oxidation by dissolved O2

possible

4Fe2+ + O2 + 4H + = 4Fe3+ + 2H20

(10)

In the deeper, anoxic layers, ferric iron concentrations can be maintained by Fe(II1) solid-phase reduction with Fe(B) organic complexes as mentioned above. 4.4. Diagenesis of S and Fe in Ombrogenic Peat Bogs, and Comparison with Other Anoxic Wetland Sediments 4.4. I. Sulfur diagenesis The sole source of sulfur to the ombrogenic peats at EGr is atmospheric deposition (Shotyk and Steinmann, 1994; Shotyk, 1996a,b). During a recent eight year period, annual precipitation in this part of Switzerland averaged 22 PM sulfate. In the uppermost 2-3 m at EGr, and in the first 40 cm at TGe, the concentrations of ST in the porewaters ( 1525 ,uM) are comparable to rainwater concentrations which are approximately one thousand times lower than the seawater average (Berner and Berner, 1987). Not only is the total depositional flux to the bog low, but this is also the sole source of S to the plants living at the bog surface. Most plants contain on the order of 0.2% S, with 1 mM sulfate in solution generally required for optimal growth (Anderson, 1978). The plants growing on the surface of the bog, therefore, receive fifty times less S than this and S, along with N and P (Small, 1972; Malmer, 1993), are growth-limiting plant nutrients. The sulfate which is supplied to the bog surface is rapidly taken up by plants (Urban and Bayley, 1986) and incorporated in organic S compounds (Urban et al., 1989). It has been suggested that during the summer when the water table in bogs is drawn down, some of this organic S may be oxidized to sulfate (Bayley et al., 1986) which would then be available for dissimilatory reduction. However, this process does not appear to be of great importance in the Jura bogs: in the porewaters collected from June to December, most (90-99%) of the S in the porewaters was in an organic form. Therefore, the combination of efficient biological scavenging of sulfate by the living plant layer and conversion to organic sulfur compounds greatly reduces the availability of sulfate for microbial reduction. In contrast to bogs, sulfate is readily available for dissimilatory sulfate reduction in the anoxic sediments of salt marshes which are regularly inundated by seawater (Howarth and Jorgensen, 1984; Swider and Ma&in, 1989). In this environment sedimentary pyrite is the ultimate authigenic reservoir for most of this S (Lord and Church, 1983; Howarth and Giblin, 1983). While the availability of decomposable organic matter may limit sulfate reduction in salt marshes (Howarth and Teal, 1979; Griffin and Rabenhorst, 1989)) this seems an unlikely limitation in bogs where the substrate (peat) is 97-99% by weight organic matter and the porewaters are rich in DOC. The availability of sulfate, therefore, seems the most likely factor limiting sulfate reduction in bogs. The complexity of the sulfur cycle in wetlands, and its relationship to the Fe cycle, has been reviewed by Luther

1157

and Church ( 1992). They note that in general, inorganic S compounds predominate in marine systems whereas in freshwater systems, organic S compounds are most important. This statement is true in general because the low concentrations of sulfate which characterize most freshwater systems limit sulfate reducing bacteria (Cappenberg, 1974), allowing methanogens to compete successfully for available substrates such as acetate (Cappenberg, 1974; Martens and Bemer, 1974). Organic S has been found to be the most important S pool in many freshwater lakes, accounting for up to 90% of the total S in the sediments (Hesse, 1958; Mitchell et al., 1984; Marnette et al., 1993). Independent studies have shown that this organic S forms from sulfate reduction in the sediments as opposed to settling of particulate organic S derived from the watershed (Nriagu and Soon, 1985; Rudd et al., 1986). In some freshwater lakes, however, inorganic S forms predominate (White et al., 1989; Giblin et al., 1990) and in lakes subjected to significant anthropogenic atmospheric S inputs, organic S may be insignificant (Carignan and Tessier, 1988). The sulfate concentrations supplied to lakes is highly variable, depending upon both geological and anthropogenic factors, and the dominant forms of S in the sediments may be either organic or inorganic. In contrast, ombrogenic bogs represent a unique natural water system in that all of the inputs are provided exclusively by atmospheric deposition. Thus, the S inputs are generally low enough to prevent dissimilatory sulfate reduction, and organic forms of S account for most of the S in freshwater peats (Casagrande et al., 1977, 1980; Brown, 1985; Bustin and Lowe, 1987), with the majority being carbon-bonded S (Lowe and Bustin, 1985; Lowe, 1986; Wieder et al., 1987; Wieder and Lang, 1988). 4.4.2. Iron diagenesis In marine sediments and salt marshes, the formation of Fe sulphides is an important sink for both Fe and S (Bemer, 1984), and the chemistry of sedimentary iron sulfide formation has been reviewed in detail by Morse et al. ( 1987) and Rickard et al. ( 1995). In marine sediments, porewaters are typically saturated with respect to Fe monosulfides, and pyrite forms slowly from the gradual conversion of iron monosulfides (Bemer, 1970, 1984). Here, sulfate is present in abundance and the main factor limiting pyrite formation is the availability of metabolizable organic matter (Bemer, 1970, 1984). In contrast, pyrite forms rapidly in the sediments of salt marshes (Howarth, 1978; Howarth and Teal, 1979; Giblin and Howarth, 1984; Kostka and Luther, 1995 ) Howarth ( 1978), for example, showed that the porewaters are undersaturated with respect to Fe monosulfides (e.g., Ksp mackinawite = 2.75 X lo-“) but saturated with respect to pyrite ( Ks, = 2.4 X 10 -**). Salt marsh sediments are made up mainly of detrital inorganic material, with total Fe concentrations typically on the order of 2-3 wt% (Kostka and Luther, 1994). In these sediments the formation of Fe sulfides may be limited only by the availability of easily reducible Fe (King et al., 1982). The recent reports of Fe transformations in salt marshes reveal a very dynamic Fe cycle in which the transformations between Fe( III) oxides and pyrite take place rapidly ( Kostka and Luther, 1995 ) In these pore-

P. Steinmann and W. Shotyk

1158

waters, concentrations of dissolved sulfide may reach 10,000 FM and dissolved Fe(I1) 1200 /.LM (Kostka and Luther, 1995). In peat bogs, the formation of Fe sulfides is prevented both by the low availability of sulfate (see above) and reducible Fe. In contrast to the salt marshes, ombrogenic bog peats receive their mineral matter exclusively from the atmosphere and as a result, usually contain no more than l-3% by weight total inorganic material (Shotyk, 1988). The total concentration of Fe in the ombrogenic peats at EGr, for example, is usually on the order of 500 ppm or less (Steinmann and Shotyk, 1997). Even if most of this Fe is supplied as small particles ( pi:10 pm) of goethite (Hoffmann et al., 1994) which should be highly reactive, these Fe concentrations are still a factor of fifty times lower than in salt marsh sediments. In addition, as we have shown here, organic complexation accounts for most of the dissolved Fe, yielding very high ferric to ferrous Fe ratios and further decreasing the availability of Fe for sulphide mineral formation. As a result of the low concentrations of total Fe in the peat combined with the extent of organic complexation and the low pH of the waters, the porewaters are highly undersaturated with respect to Fe sulphides. The diagenesis of Fe in the porewaters described here explains why Fe sulphide minerals have never been found in ombrogenic peats (Shotyk, 1988; Shotyk et al., 1992). While Fe sulphides may form in minerogenic freshwater peats if there is a local source of sulfate (Rost, 1922; Papunen, 1966; Urban et al., 1989), they are most commonly found in peats which have formed in brackish waters (Casagrande et al., 1977; Postma, 1982, 1983). Two other possible inorganic controls on the Fe concentrations in freshwater peats are the formation of siderite, FeC03, and vivianite, Fe3 ( PO4)2 (Table 4). However, the formation of siderite is possible only in anoxic, CO*-charged peats when the pH is above 7 (Puustjarvi, 1952; Postma, 1977; Shotyk et al., 1992). Thus, the formation of siderite is restricted to neutral to alkaline minerogenic fens and swamps (van Bemmelen,

1899; Kivinen, 1936; Bushinskii, 1946; Puustjarvi, 1952; Lukashev and Kovalev, 1969; Postma, 1977, 1981, 1982, 1983; Selmer-Olsen and Lie, 1983; Virtanen, 1994). While vivianite is sometimes found in peats (Douglas, 1767; van Bemmelen, 1899), the formation of this mineral requires a significant external supply of phosphorus (Puustjkvi, 1952; Postma, 1977, 1981, 1982, 1983). Recent studies of vivianite formation in Finnish peats revealed P concentrations in the vivianite-bearing peats up to one hundred times higher than background levels (Vi&men, 1994). Thus, the formation of Fe sulphide, carbonate, and phosphate minerals in ombrogenie bogs is not possible. Table 4 illustrates that porewaters in the ombrogenic sections of the bogs EGr and TGe are highly undersaturated with respect to siderite and vivianite. With increasing depth the extent of undersaturation becomes less important; saturation with respect to siderite is only reached in the CO,-charged bottom part at TGe, where pH approaches neutrality. If mineral precipitation reactions do not regulate the concentration of Fe in ombrogenic bog waters, what is the ultimate fate of this dissolved Fe? The relatively high concentrations of dissolved Fe relative to typical natural waters are characteristic of peat bogs (Smoroditsev and Adova, 1929; Bushinskii, 1946; Kovalev and Generalova, 1969; Clausen and Brooks, 1983; Urban et al., 1987). Relatively high concentrations of dissolved Fe in bog waters, however, can only be maintained when O2 is absent (Sparling, 1966). As the waters in ombrogenic bogs slowly seep out into the surrounding lagg zone (Keough and Pippen, 1984)) they become oxygenated, sometimes the result of this is seen as oily films which begin to form at the air/water interface; these films are a colloidal mixture of organic matter and ferric hydroxide (Puustjlrvi, 1952; Henrot and Wieder, 1990). While organic complexation can certainly reduce the rate of chemical oxidation of ferrous Fe (Theis and Singer, 1974; Sung and Morgan, 1980; Millero, 1989), the production of these iron oxides is catalyzed by iron-oxidizing bacteria which accelerate the process (Thunmark, 1942; Crerar

Table 4: SaNration indims for selected imn phasesat various depths in the TGe. (Tl2) and EGr (E9) pmtXcs, calculated fmm data given in the Appendix usinS PHREEQE. (ml: not de.te.rmi& PO4 or HS- below &&on

site.,depth

F&i pmipitate a

0

siderite.(Few) b3 W/K)

limit)

vivianite (Fe3(PO4)rsHfi) log wiw

TGe, 12 cm

-7.42

-5.39

-17.26

TGe. 30 cm

-8.17

-4.83

-15.55

-10.56

-4.06

-10.66

ad.

-1.53

n.d.

TGe, 98 cm

n.d.

-0.61

-6.73

TGe, 140 cm

n.d.

0.74

-3.04

EGr. 55 cm

n.d.

-5.98

-17.25 -16.35

TGe, 50 cm TGc, 17 cm

n.d.

-4.87

EGr, 295 cm

cd.

-4.55

ad.

EGr, 415 cm

cd.

-3.69

-13.54

EGr, 535 cm

ad.

-1.47

-6.58

EGr, 175 cm

Porewater

profiles

from peat bogs in the Jura Mountains

et al., 1979; Ghiorse and Chapnick, 1983; Lahdesmaki, 1984; Wheatley, 1988). The reductive dissolution of iron oxides which are deposited on the bog surface and the seepage of anoxic, Fe-bearing fluids to the oxygenated surface layers may explain why the concentrations of Fe in peats are generally higher around the margins of the bog (Fuustj&vi, 1952; Holmen, 1964). The ultimate oxidation product is bog iron ore, a complex mixture of amorphous and crystalline iron oxides and organic matter (Griffin, 1893; Ljunggren, 1953, 1955;

De Geyter

5.

et al.,

1985).

SUMMARY AND CONCLUSIONS

Chemical analysis of all important inorganic solutes in the porewaters from two peat bogs allowed the relative importance of CO2 and organic acids to the pH of these waters to be quantified. In the porewaters of the upper sections of the bogs where inputs of inorganic solids arise exclusively from atmospheric deposition (i.e., in the ombrogenic layers), the waters are acidic (pH 4), H+ is the dominant cation, and humic materials the dominant anions. The dissociation of organic acids is the most important source of protons, while sulfate and nitrate reduction are important proton consuming reactions. In deeper, minerotrophic layers, the pore fluids are neutralized to pH 7 by dissolution of mineral matter in the sediments; in this zone, the pH is controlled by dissolved CO* species: HCO; is the most important anion and Ca*+ the dominant cation. At intermediate depths, in between these two extremes, there is a transition zone whose pH is 4.5 or higher; here, organic acids and dissolved CO1 are equally important acids, and NH,f may be the single most important cation. The mineralisation of organic N compounds (ammonification) which is the source of NH: to these waters is an important sink for protons. Most of the dissolved S in the waters (90-99%) is in an organic form. While the presence of measurable sulfide in the porewaters at TGe indicate that the waters are anoxic, sulfate is quantitatively more important. Rainwater is the only source of sulfate to the surface layers of the bogs, and the low rate of supply combined with the rapid removal and conversion to organic S by the plants appears to limit dissimilatory sulfate reduction in the porewaters. The presence of Fe(I1) in the porewaters from both bogs supports the view that the waters are anoxic, but the Fe( III)/ Fe( II) ratio is only 1:3 at TGe, and essentially 1: 1 at EGr. The formation of stable organic complexes of Fe3+ may account for the unexpectedly high ratio of Fe( III)/Fe( II) in these anoxic waters. Acknowledgments-We are grateful to Professor Albert Matter of this Institute for providing laboratory facilities and equipment. Many thanks to Dr. Peter Blaser and Dr. JGrg Luster (at WSL, Birmensdorf) for ICP and DOC analyses and fruitful discussions. An introduction to PHREEQE by Dr. Nick Waber and field assistance by Mrs. Annatina Janett is highly appreciated. This study was financially supported by the Canton of Bern (SEVA Lottofonds) and the Swiss National Science Foundation (Grants 21 30207.90 and 2036371.92). Sincere thanks are also extended to Dr. Eric Reardon and to three anonymous reviewers for their helpful comments on an earlier version of this manuscript.

Editorial

handling:

E. J. Reardon

1159

REFERENCES Anderson J. W. (1978) Sulphur in Biology. Edward Arnold. Armstrong A. C. ( 1995) Hydrological model of peat-mound form Earth Surf: Proc. with vertically varying hydraulic conductivity. Landforms Z&473-477. Baden W. and Eggelsmann R. ( 1963) Zur Durchllssigkeit der MoorbBden. Z. Kulturtechn. Flurberein. 4, 226-254. Bayley S. E., Behr R. S., and Kelley C. A. (1986) Retention and release of S from a freshwater wetland. Water Air Soil Polluz. 31, 101-114. Bemer E. K. and Bemer R. A. (1987) The Global Water Cycle. Prentice-Hall. Bemer R. A. ( 1970) Sedimentary pyrite formation. Amer. J. Sci. 268, l-23. Bemer R. A. ( 1981) A new geochemical classification of sedimentary environments. J. Sediment. Petrol. 51, 359-365. Bemer R. A. ( 1984) Sedimentary pyrite formation: An update. Geochim. Cosmochim. Acta 48, 605-615. Boelter D. H. ( 1969) Physical properties of peats as related to degree of decomposition. Soil. Sci. Sot. Amer. Proc. 33, 606-609. Brand1 H. and Hanselmann K. W. ( 1991) Evaluation and application of dialysis porewater samplers for microbiological studies at sediment-water interfaces. Aquat. Sci. 53, 55-73. Brown K. A. (1985) Sulphur distribution and metabolism in waterlogged peat. Soil Biol. Biochem. 17, 246-253. Burgess J. ( 1989) Ions in solution: Basic Principles of Chemical Interactions. Wiley and Sons. Bushinskii G. I. (1946) The conditions in the formation of siderites, vivianites, and brown iron ores in the peat bogs of White Russia. Byull. Moskov. Obschchestva Ispytat. Prirody, Otdel Geol. 21, 65-80. Bustin R. M. and Lowe L. E. ( 1987) Sulphur, low temperature ash, and minor elements in humic-temperate peat of the Fraser River Delta, British Columbia. J. Geol. Sot. London 144, 435-450. Cappenberg T. E. ( 1974) Interrelations between sulfate-reducing and methane-producing bacteria in bottom deposits of a freshwater lake. I. Field observations. Ant. Leeuw. J. 40, 285-295. Carignan R. and Tessier A. ( 1988) The codiagenesis of sulfur and iron in acid lake sediments of southwestern QuBbec. Geochim. Cosmochim. Acta 52, 1179- 1188. Carter G. S. (1955) Thepapyrus Swamps of Uganda. W. Heffer and Sons. Casagrande D. J., Gronli K., and Sutton N. ( 1980) The distribution of sulfur and organic matter in various fractions of peat: Origins of sulfur in coal. Geochim. Cosmochim. Acta 44, 25-32. Casagrande D. J., Siefert K., Berschinski C., and Sutton N. (1977) Sulfur in peat-forming systems of the Okefenokee Swamp and Florida Everglades: Origins of sulfur in coal. Geochim. Cosmochim. Acta 41, 161-167. Casey W. H. and Lasaga A. L. ( 1987) Modeling solute transport and sulfate reduction in marsh sediments. Geochim. Cosmochim. Acta 51, 1109-1120. Clausen J. C. and Brooks K. N. ( 1983) Quality of runoff from Minnesota peatlands: I. A characterization. Water Res. Bull. 19, 763767. Clymo R. S. (1983) Peat. In Mires: Swamp, Bog. Fen, and Moor. (ed. A. J. P. Gore); pp. 159-224. Ecosyi. Woyld. 4A, Elsevier. Crerar D. A., Knox G. W.. and Means J. L. c 1979) Biogeochemistrv of bog iron in the New Jersey Pine Barrens.’ Che%. Geol. 24, 111-135. Crum H. (1987) A Focus on Peatlands and Peat Mosses. Univ. Michigan Press. Damman A. W. H. ( 1986) Hydrology, development, and biogeochemistry of ombrogenous peat bogs with special reference to nutrient relocation in a western Newfoundland bog. Canadian J. Bot. 64, 384-394. Damman A. W. H. ( 1987) Variation in ombrotrophy: Chemical differences among and within ombrotrophic bogs. Proc. Symp. ‘87 Wetlands/Peatlands, 85-93. Dau J. H. C. ( 1823) Neues Handbuch iiber den Tor$ J. C. Hinrich’sche Buchhandlung. De Geyter G., Vandenberghe R. E., Verdonck L., and Stoops G. ( 1985) Mineralogy of Holocene bog-iron ore in northern Belgium. N. .I. M. Abh. 153, I- 17.

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De Haan H., Jones R. I., and Salonen K. (1990) Abiotic transformations of iron and phosphate in humic lake water revealed by double-isotope labeling and gel filtration. Limnol. Oceanogr. 3549 l497. Deiana S., Gessa C., Marchetti M., and Usai M. (1995) Phenolic acid redox properties: pH influence on iron (III) reduction by caffeic acid. Soil Sci. Sot. Amer. J. 59, 1301-1307. Douglas S. ( 1767) Experiments and observations upon a blue substance, found in peat moss in Scotland. Phil. Trans. Roy. Sot. London 181-188. Driscoll C. T. and Lehtinen D. ( 1994) Modeling the acid-base chemistry of organic solutes in Adirondack, New York, lakes. Wuter Resow. Res. 30, 297-306. Elsgaard L. and Jorgensen B. B. (1992) Anoxic transformations of radiolabeled hydrogen sulfide in marine and freshwater systems. Geochim. Cosmochim. Acta 56, 2425-2435. Ephraim J. H. and Marinsky J. A. ( 1990) Ultrafiltration as a technique for studying metal-humate interactions: Studies with Fe and Cu. Anal. Chim. Acta 232, 171- 180. Fossing H. and Jorgensen B. B. (1990) Oxidation and reduction of radiolabeled inorganic sulfur compounds in an estuarine sediment, Kysing Fjord, Denmark. Geochim. Cosmochim. Acta 54, 27312742. Fritz S. F. and Popp R. K. (1985) A single dissolution technique for determining Fe0 and Fe203 in rock and mineral samples, Amer. Miner. 70,961-968. Gerber E. and Monbaron M. ( 1990) Die Torfmoore des nordlichen Faltenjuras in ihrem morpho-strukturalen Zusammenhang. UKPIK 7, 31-44. Ghiorse W. C. and Chapnick S. D. ( 1983) Metal-depositing bacteria and the distribution of manganese and iron in swamp waters. Ecol. Bull. (Stockholm) 35, 367-376. Giblin A. E. and Howarth R. W. ( 1984) Porewater evidence for a dynamic sedimentary iron cycle in salt marshes. Limnol. Oceanogr. 29, 47-63. Giblin A. E., Likens G. E., White D., and Howarth R. W. (1990) Sulfur storage and alkalinity generation in New England lake sediments. Limnol. Oceanogr. 35, 852-869. Gledhill M. and van den Berg C. M. G. (1994) Determination of complexation of iron (III) with natural organic complexing ligands in seawater using cathodic stripping voltammetry. Mar. Chem. 47, 41-54. Gore A. J. P. (ed.) (1983) Mires: Swamp, Bog, Fen, and Moor. Ecosys. World 4 (2 volumes). Gorham E., Eisenreich S. J., Ford J., and Santelmann M. V. (1985) The chemistry of bog waters. In Chemical processes in Lakes (ed. W. Stumm), pp. 339-363. Wiley. Gottlich K. (ed.) ( 1990) Moor- und Torf kunde. 3rd German ed. E. Schweizerbart’sche Verlagsbuchhandlung. Griffin P. H. (1893) The Manufacture of Charcoal-Iron from the Bog- and Lake-Ores of Three Rivers District, Province of Quebec, Canada. Amer. Inst. Min. Met. Eng. Trans. 21, 974-992. Griffin T. M. and Rabenhurst M. C. ( 1989) Processes and rates of pedogenesis in some Maryland tidal marsh soils. Soil Sci. Sot. Amer. J. 53, 862-870. Hem J. D. ( 1972) Chemical factors that influence the availability of iron and manganese in aqueous systems. Geol. Sot. Amer. Spec. Pap. 140, 17-24. Hemond H. F. (1980) Biogeochemistry of Thoreou’s bog, Concord, Massachusetts. Ecol. Monogr. 50, 507-526. Henrot J. and Wieder K. ( 1990) Processes of iron and manganese retention in laboratory peat microcosms subjected to acid mine drainage. J. Environ. Qual. 19, 312-320. Hesse P. R. (1958) The distribution of suluhur in the muds. water and vegetation of Lake Victoria. Hydrob;:ologia 11, 29-39. Hesselman H. (1910) Om vattnets syrehalt och dess inverkan pa skogsmarkens forsumpning och skogens vaxtlighet. Meddelanden fran Statens Skogsforsoksanstalt 7,91- 125 (in Swedish with German summary) Hill B. M. and Siegel D. I. (1991) Groundwater flow and the metal content of peat. J. Hydrol. 123, 21 l-224. Hoag R. S. and Price J. S. (1995) A field-scale, natural gradient solute transport experiment in peat at a Newfoundland blanket bog. J. Hydrol. 172, 171-184.

and W. Shotyk Hobbs N. B. ( 1986) Mire morphology and the properties and behaviour of some British and foreign peats. Quart. J. Engin. Geol. London 19, 7-80. Hoffman P., Sinner T., Dedik A. N., Karandashev V. K., Malyshev A. A., Weber S., and Ortner H. M. (1994) Iron in atmospheric aqueous and particulate samples. Fres. J. Anal. Chem. 350, 3437. Holmen H. (1964) Forest ecological studies on drained peatland in the Province of Uppland, Sweden. Parts I-III. Stud. Forest. Suecica 16. l-23 I Hordijk C. A., van Engelen J. J. M., Jonker F. A., and Cappenberg T. E. ( 1989) Determination of total sulfur in freshwater sediments by ion chromatography. Water Res. 23, 853-859. Howarth R. W. ( 1978) Pyrite: Its rapid formation in a salt marsh and its importance in ecosystem metabolism. Science 203, 4951. Howarth R. W. and Giblin A. ( 1983) Sulfate reduction in the salt marshes at Sapelo Island, Georgia. Limnol. Oceanogr. 28, 7082. Howarth R. W. and Jorgensen B. B. (1984) Formation of “S-labelled elemental sulfur and pyrite in coastal marine sediments Geochim. Cosduring short-term %O:- reduction measurements. mochim. Actu 48, 1807- 18 18. Howarth R. W. and Teal J. M. ( 1979) Sulfate reduction in a New England salt marsh. Limnol. Oceanogr. 24, 999-1013. Hummel W. ( 1992) Thermodynamic Database for Organic Ligands. PSI Report TM-41 -91-43 PES-91-577. Hunt T. S. (1865) Contributions to the chemistry of natural waters. Amer. J. Sci. 89, 1766193. Ingram H. A. P. ( 1982) Size and shape in raised mire ecosystems: a geophysical model. Nature 297, 300-303. Johnson C. W. (1985) Bags af the Northeast. Univ. Press New England. Joray M. (1942) L’etang de la Gruyere. Etude pollenanalytique et stratigraphique de la tourbiere. Mm. leve’geoobot. de la Suisse 25, l-117 (in French). Jorgensen B. B. (1990) The sulfur cycle of freshwater sediments: Role of thiosulfate. Limnol. Oceanogr. 35, 1329- 1342. Keough J. R. and Pippen R. W. (1984) The movement of water Canadian J. Bot. 62, from peatland into surrounding groundwater. 835-839. King G. M., Klug M. J., Wiegert R. G., and Chalmers A. G. ( 1982) Relation of soil water movement and sulfide concentration to Spartina alternifiora production in a Georgia salt marsh. Science 218, 61-63. Kivinen E. (1935) Uber Elektrolytgehalt und Reaktion der Moorwasser. In Agrogeologisia Julkaisuju, Vol. 38, p. 71 (in German). Kivinen E. (1936) Zur Kenntnis der Eisenkarbonatvorkommnisse in den Mooren Finnlands. 2. Pflanzenernahr. Dtingung Bodenk. 45, 96-104. Kivinen E. and Pakarinen P. ( 1981) Geographical distribution of peat resources and major peatland complex types in the world. Ann. Acad. Sci. Fennicae, A. III 132. Koenings J. P. (1976) In situ experiments on the dissolved and colloidal state of Fe in an acid bog lake. Limnol. Oceanogr. 21, 674-683. Kostka J. E. and Luther G. W. ( 1994) Partitioning and speciation of solid phase Fe in saltmarsh sediments. Geochim. Cosmochim. Acta 58, 1701-1710. Kostka J. E. and Luther G. W. (1995) Seasonal cycling of Fe in saltmarsh sediments. Biogeochem. 29, I59- 18 I. Kotilainen M. J. ( 1927) Untersuchungen tiber die Beziehungen zwischen der Pflanzendecke der Moore und der Beschaffenheit, besonders der Reaktion des Tortbodens. In Suomen Suoviljelysyhdistys. Tieteelisi Julkaisuja, Vol. 7, p. 219 (in German). Kovalev V. A. and Generalova V. A. (1969) Geochemical aspects of the movement of Fe in recent peat bogs of Byelorussia. Geochem. Intl. 6, 144-153. Krom M. D. and Sholkovitz E. R. ( 1978) On the association of Fe and Mn with organic matter in anoxic marine pore waters. Geochim. Casmochim. Acta 42, 607-611. Lahdesmaki P. (1984) Humic acids and microorganisms stimulate the precipitation of iron in bogs. Sue. 35, 35-39. LaKind J. S. and Stone A. T. ( 1989) Reductive dissolution of goe-

Porewater

profiles

from peat bogs in the Jura Mountains

thite by phenolic reductants. Geochim. Cosmochim. Acta 53,961971. Landva A. O., Korpijaakko E. O., and Pheeney P. E. (1986) Notes on the original von Post peat and peatland classification system. In Advances in Peatlands Engineering pp. 17-29. Nat]. Res. Count. Canada. Lewis B. L. et al. ( 1995) Voltammetric estimation of iron(III) thermodynamic stability constants for catecholate siderophores isolated from marine bacteria and cyanobacteria. Mar. Chem. 50, 179-188. Liang L., McCarthy J. F., Jolley L. W., McNabb J. A., and Mehlhom T. L. ( 1993) Iron dynamics: Transformation of Fe(II)/Fe( III) during injection of natural organic matter in a sandy aquifer. Geochim. Cosmochim. Acta 57, 1987-1999. Lindberg R. D. and Runnells D. D. (1984) Ground water redox reactions: An analysis of equilibrium state applied to Eh measurements and geochemical modeling. Science 225, 925-927. Ljunggren P. ( 1953) Some data concerning the formation of manganiferous and ferriferous bog ores. Geol. Foren. Forhandl. Bd. 75, 277. Ljunggren P. ( 1955) Geochemistry and radioactivity of some Mn and Fe bog ores. Geol. Foren. Forhandl. Bd. 77, 33. Loiselle A. J. ( 1986) Development of a technique for sampling interstitial waters in peatlands. In Advances in Peatlands Engineering pp. 174- 183. Nat]. Res. Count. Canada. Lord C. J. and Church T. M. (1983) The geochemistry of salt marshes: Sedimentary iron diffusion, sulfate reduction, and pyritization. Geochim. Cosmochim. Acta 47, 138 1- 1391. Lovely D. R. ( 1987) Organic matter mineralization with the reducJ. 5, 375-399. tion of ferric iron: A review. Geomicrobiology Lowe L. E. ( 1986) Application of a sequential extraction procedure to the determination of the distribution of sulfur forms in selected peat materials. Canadian J. Soil Sci. 66, 337-345. Lowe L. E. and Bustin R. M. ( 1985) Distribution of sulphur forms in six facies of peats of the Fraser River Delta. Canadian J. Soil Sci. 65, 531-541. Lukashev K. I. and Kovalev V. A. (1969) Mineralogical-geochemical systems of iron in recent peat bogs. Dokl. Akad. Nauk SSSR 187, 1390-1393. Luther G. W. and Church T. M. ( 1992) An overview of the environmental chemistry of sulphur in wetland systems. In Sulphur Cycling on the Continents (ed. R. W. Howarth et al.), pp. 125- 142. Wiley. Luther G. W., Church T. M.. Giblin A. E., and Howarth R. W. ( 1986) Speciation of dissolved sulfur in salt marshes by polarographic methods. In Organic Marine Chemistry (ed. M. L. Sohn) ; ACS Symp. Ser. 305, 340-355. Luther G. W., Kostka J. E., Church T. M., Sulzberger B., and Stumm W. ( 1992) Seasonal iron cycling in the salt marsh sedimentary environment: The importance of ligand complexes with Fe( II) and Fe( III) in the dissolution of Fe(II1) minerals and pyrite, respectively. Mar. Chem. 40, 81- 103. Luther G. W., Shellenbarger P. A., and Brendel P. J. (1996) Dissolved organic Fe( III) and Fe( II) complexes in salt marsh porewaters. Geochim. Cosmochim. Acta 60, 95 l-960. Malmer N. ( 1986) Vegetational gradients in relation to environmental conditions in northwestern European mires. Canadian J. Bat. 64, 375-383. Malmer N. ( 1993) Mineral nutrients in vegetation and surface layers of Sphagnum-dominated peat-forming systems. Adv. Blyology 5, 223-248. Malmstrnm C. (1923) Degero stormyr. En botanisk, hydrologisk och utvecklingshistorisk undersokning over ett nordsvenskt myrkomplex. Meddelanden fran Statens Skogsforsoksanstalt 20, l206 (in Swedish with German summary). Marnette E. C. L., Van Breemen N., Hordijk K. A., and Cappenberg T. E. ( 1993) Pyrite formation in two freshwater systems in the Netherlands. Geochim. Cosmochim. Acta 57, 4165-4177. Martens C. S. and Bemer R. A. (1974) Methane production in the interstitial waters of sulfate-depleted marine sediments. Science 185, 1167-I 169. McKnight D., Thurman E. M., Wershaw R. L., and Hemond H. ( 1985) Biogeochemistry of aquatic humic substances in Thoreau’s Bog, Concord, Massachusetts. Ecol. 66, 1339- 1352.

1161

Miller0 F. J. (1989) Effect of ionic interactions on the oxidation of Fe(I1) and Cu(I) in natural waters. Mar. Chem. 28, 1- 18. Mitchell M. J., Landers D. H., Brodowski D. G., Lawrence G. B., and David M. B. ( 1984) Organic and inorganic sulfur constituents of the sediments in three New York lakes: Effect of site, sediment depth and season. Water Air Soil Pollur. 21, 231-245. Morris J. C. and Stumm W. (1967) Redox equilibria and measurements of potentials in the aquatic environment. In Equilibrium Concepts in Natural Water Systems (ed. W. Stumm); Adv. Chem. 67, 270-285. Morse J. W., Miller0 F. J., Cornwell J. C., and Rickard D. ( 1987) The chemistry of the hydrogen sulfide and iron sulfide systems in natural waters. Earth Sci. Rev. 24, l-42. Nriagu J. 0. and Soon Y. K. ( 1985) Distribution and isotopic composition of sulfur in lake sediments of northern Ontario. Geochim. Cosmochim. Acta 49, 823-834. Papunen H. ( 1966) Framboidal texture of the pyritic layer found in a peat bog in S.E. Finland. Compres Rendus Sot. Geol. Finland 38, 117-125. Parker D. R., Norvell W. A., and Chaney R. L. ( 1990) GEOCHEMPC Version 2.0, Department of Soil and Environmental Sciences, University of California, Riverside. Parkhurst D. L., Thorstenson D. C., and Plummer L. N. (1990) PHREEQE (Water-Resources Investigations Report No. 80-96, revised edition). USGS. Perdue E. M., Beck K. C., and Reuter J. H. ( 1976) Organic complexes of Fe and Al in natural waters. Narure 260, 418-420. Postma D. (1977) The occurrence and chemical composition of recent Fe-rich mixed carbonates in a river bog. J. Sediment. Petrol. 47, 1089-1098. Postma D. ( 1981) Formation of siderite and vivianite and the porewater composition of a recent bog sediment in Denmark. Chem. Geol. 31, 225-244. Postma D. (1982) Pyrite and siderite formation in brackish and freshwater sediments. Amer. J. Sci. 282, 115 I- 1 183. Postma D. ( 1983) Pyrite and siderite oxidation in swamp sediments. J. Soil Sci. 34, 163-182. of iron in peat soils. Acta Puustjtivi V. ( 1952) The precipitation Agral. Fenn. 78, 72. Ramann E. ( 1895) Organogene Ablagerungen der Jetztzeit. Neves Jahrb. Mineral Geol. Palacontol. 10, 119- 166 (in German). Rausa R., Deiana S., and Gessa C. (1994) Reduction of iron(II1) by synthetic humic acids obtained via coal oxidation. In Humic Substances in the Global Environment and Implications on Human Health (ed. N. Senesi and T. M. Miano), pp. 1177- 1182. Elsevier. Reuter J. H. and Perdue E. M. ( 1977) Importance of heavy metalorganic solute interactions in natural waters. Geochim. Cosmochim. Acta 41, 325-334. Rickard D., Schoonen M. A., and Luther G. W. (1995) Chemistry of iron sulfides in sedimentary environments. In Geochrmical Transformations of Sedimentary Sulfur (ed. M. A. A. Schoonen), pp. 168- 193. American Chemical Society. Rost C. 0. (1922) Occurrence of sulfides in Minnesota peat soils. Soil Sci. 14, 167- 174. Rudd J. W. M., Kelly C. A., and Furutani A. (1986) The role of sulfate reduction in long term accumulation of organic and inorganic sulfur in lake sediments. Limnol. Oceanogr. 31, 128 I- 1291, Selmer-Olsen A. R. and Lie 0. (1983) Jerndarbonatutfelling (Sideritt) i myr. Saertrykk av Jord og Myr nr. 4, 13 I- 140 ( in Norwegian with English summary). Senesi N., Griffith S. M., Schnitzer M., and Townsend M. G. ( 1977) Binding of Fei’ by humic materials. Geochim. Cosmochim. Actu 41,969-976. Shannon R. D. and White J. R. (1996) The effects of spatial and temporal variations in acetate and sulfate on methane cycling in two Michigan peatlands. Limnol. Oceanogr. 41, 435-443. Shotyk W. ( 1987) European contributions to the geochemistry of peatland waters, 1890- 1940. Proc. Symp. ‘87 Wetlands/Peatlands, 115-125. Shotyk W. ( 1988) Review of the inorganic geochemistry of peats and peatland waters. Earth Sci. Rev. 25, 95- 176. Shotyk W. (1989a) An overview of the geochemistry of methane dynamics in mires. Intl. Pear J. 3, 25-44.

1162

P. Steinmann

Shotyk W. (1989b) The chemistry of peatland waters. Water Quality 103. Bull. 14, 47-58, Shotyk W. (1996a) Peat bog archives of atmospheric metal deposition: Geochemical evaluation of peat profiles, natural variations in metal concentrations, and metal enrichment factors. Env. Revs. 4, 149- 183. Shotyk W. ( 1996b) Natural and anthropogenic enrichments of As, Cu, Pb, Sb, and Zn in ombrotrophic vs. minerotrophic peat bog profiles, Jura Mountains, Switzerland. Water Air Soil Pollut. 90, 375-405. Shotyk W. and Steinmann P. ( 1994) Porewater indicator of rainwater-dominated vs. groundwater-dominated peat bog profiles, Jura Mountains, Switzerland. Chem. Geol. 116, 137-146. Shotyk W., Nesbitt H. W., and Fyfe W. S. (1992) Natural and anthropogenic enrichments of trace metals in peat profiles. Intl. J. Coal Geol. 20, 49-84. Siegel D. I. and Glaser P. H. (1987) Groundwater flow in a bog/ fen complex, Lost River peatland, northern Minnesota. J. Ecol. 75,743-754. Small E. (1972) Ecological significance of four critical elements in plants of raised Sphagnum peat bogs. J. Ecol. 53, 498-503. Smorodintsev I. A. and Adova A. N. (1929) Jelesystiyeh regim torfyanix vod. Ruskii Zhumal Tropicheskoi Meditsinii 7, 143-153 (in Russian with French summary). Sparling J. H. (1966) Studies on the relationship between water movement and water chemistry in mires. Canadian J. Bot. 44, 747-758. Steinmann P. ( 1995) Geochemistry of atmospheric inputs and pore water genesis in two contrasting Sphagnum bog profiles, Jura Mountains, Switzerland. Ph.D. thesis, Univ. Berne. Steinmann P. and Shotyk W. (1995a) Ion chromatography of organic-rich natural waters from peatlands: 3. Improvements for measuring anions and cations. J. Chromatogr. A 706, 28 l-286. Steinmann P. and Shotyk W. (1995b) Ion chromatography of organic-rich natural waters from peatlands: 4. Fe+2 and Fe+3. J. Chromatogr. A 706, 293-299. Steinmann P. and Shotyk W. (199%) Ion chromatography of organic-rich natural waters from peatlands: 5. Sulfide and acid volatile sulfur (AVS). J. Chromatogr. A 706, 287-292. Steinmann P. and Shotyk W. (1996) Sampling anoxic porewaters in peatlands using peepers for in situ filtration. Fres. J. Anal. Chem. 354, 709-713. Steinmann P. and Shotyk W. ( 1997) Geochemistry, mineralogy, and geochemical mass balance on major elements in two peat bog profiles (Jura Mountains, Switzerland). Chem. Geol., (in press). Stucki J. W. and Komadel P. (1988) Q uantitative assay of minerals for Fe2+ and Fe3+ using l,lO-Phenanthroline: III. A rapid photochemical method. Clays Clay Miner. 36, 379-381. Stumm W. and Lee G. F. (1960) The chemistry of aqueous iron. Schw. Z. Hydrol. 22, 295-319. Stumm W. and Morgan J. J. ( 1981) Aquatic Chemistry (2nd ed.). Wiley-Intersci. Succow M. ( 1988) Landschaftiikologische Moorkunde. Gebrtider Borntrlger (in German). Sulzberger B., Suter D., Siffert C., Banwart S., and Stumm W. (1989) Dissolution of Fe(II1) (hydr)oxides in natural waters; laboratory assessment on the kinetics controlled by surface coordination. Mar. Chem. 28, 127- 144. Sung W. and Morgan J. J. ( 1980) Kinetics and product of ferrous iron oxygenation in aqueous systems. Environ. Sci. Techn. 14, 561-568. Swider K. T. and Mackin J. E. (1989) Transformations of sulfur

and W. Shotyk compounds in marsh-flat sediments. Geochim. Cosmochim. Acta 53, 2311-2323. Tarutis W. J., Unz R. F., and Brooks R. P. (1992) Behaviour of sedimentary Fe and Mn in a natural wetland receiving acidic mine drainage, Pennsylvania, U.S.A. Appl. Geochem. 7, 77-85. Thamdrup B., Finster K., Fossing H., Wtirgler H. J., and Jorgensen B. B. (1994) Thiosulfate and sulfite distribution in porewater of marine sediments related to manganese, iron, and sulfur geochemistry. Geochim. Cosmochim. Acta 58, 67-73. Theis T. L. and Singer P. C. (1974) Comulexation of iron (II) bv organic matter and its effect on iron (II) oxygenation. Environ. Sci. Tech. 8, 5699573. Thompson T. G., Lorah J. R., and Rigg G. B. (1927) The acidity of the waters of some pudget sound bogs. J. Amer. Chem. Sot. 49, 2981-2988. Thunmark S. (1942) Uber rezente Eisenocker und ihre Mikroorganismengemeinschaften. Uppsala Univ. Geol. Inst. 29, l-286 (in German). Thurman E. M. (1985) Organic Geochemistq of natural waters. Nijhoff/Junk. Urban N. R. and Bayley S. E. (1986) The acid-base balance of peatlands: A short-term perspective. Water Air Soil Poll. 30,791800. Urban N. R. and Eisenreich S. J. (1988) Nitrogen cycling in a forested minnesota bog. Canadian J. Bot. 66, 43.5-449. _ Urban N. R.. Eisenreich S. J.. and Gorham E. (. 1987) , Proton cvcline < ., in bogs: Geographical variation in northeastern North America (ed. T. C. Hutchinson and K. M. Meema); Effects ofAcidic Deposition on Forests, Wetland, and Agricultural Ecosystems; NATO ASI Series G16, 577-598. Urban N. R., Bayley S. E., and Eisenreich S. J. ( 1989) Export of dissolved organic carbon and acidity from peatlands. Water Resour. Res. 25, 1619-1628. van Bemmelen J. M. (1899) Uber das Vorkommen, die Zusammensetzung und die Bildung von Eisenanhlufungen in und unter (in German). Mooren. Z. Anal. Chem. 12, 313-379 Virtanen K. (1994) Geological control of iron and phosphorous precipitates in mires of the Ruukki-Vihanti area, central Finland. &ll.-Geol. Surv. Finland 375, l-69. von Post L. and Granlund E. ( 1925) Sodra Sveriges tortvtilluangar. Sveriges Geol. Undersok., ser. C,’No. 335, Arsbok 19. (in-Swedish). Waksman S. A. ( 1930) Chemical composition of peat and the role of microorganisms in its formation. Amer. J. Sci. 19, 32-54. Websky J. ( 1864) Beitrlge zur Erkenntnis der Zusammensetzung und Bildung des Torfes. J. Prakt. Chem. 92,65-96 (in German). Welten M. (1964) La Tourbiire de Genevez. M. H. Natutforsch. Ges. Bern. 21, 67-73 (in German). Wheatley R. E. (1988) Ochre deposits and associated bacteria in some field drains in Scotland. J. Soil Sci. 39, 253-264. Wheeler B. D. and Giller K. E. (1984) The use of dialysis cells for investigating pore water composition in wetland substrata, with particular reference to dissolved iron and sulphide. Comm. Soil Sci. Plant Anal. 15, 707-716. White J. R., Gubala C. P., Fry B., Owen J., and Mitchell M. J. ( 1989) Sediment biogeochemistry of iron and sulfur in an acidic lake. Geochim. Cosmochim. Acta 53, 2547-2559. Wieder R. K. and Lang G. E. ( 1988) Cycling of inorganic and organic sulfur in peat from Big Run Bog, West Virginia. Biogeochem. 5, 221-242. Wu J. and Luther G. W. ( 1995) Complexation of Fe( III) by natural organic ligands in the Northwest Atlantic Ocean by a competitive ligand equilibration method and a kinetic approach. Mar. Chem. 50,159-177. .

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3.05

145.1' 3.19

0.38‘ 12.96' 130.7: 3.08

31.26. 0.33

19.96

15.80

6.82

10.31

51.10.

9.96, 93.56: 2.48

9.22

8.27

8.64

7.53

6.58

6.95

5.76

116.3 10.03: <0.15' 571.3 lpi.!+

89.78

3.21

1.36. 11.73

26.88: 0.59

27.62, 1.16;

27.40

229 ._ : i~.236.9 ~ 33.88 .~ .* 7:71 332.9: 340.1 50.02,

126.7

23.02

21.55

20.08: 26.77

f8.&

17.57

15.24 20.33

12.93: lj.24

72.72; 16.9s s6.23, 14.25i.l.94

26.49

15.63 26.62

12.50; 16.67

11.72

10.94. 14.58

11.33; 75.11 26.05

0.84

24.67.; 1.79

1.50,

;1.;3: 15.64' 2i.7i3.3.4i.

10.61

0.36

3.33 3170' 40.77:

16.30: 3.441_ 0.87 1.02 .‘ 4.07 5.02

7.821

1.28

0.64

T

72.16.

79.96

73.98

58.63,

55.50,

49.31

34.25

37.84

41.44'

44.82

48.20

49.36

50.52

53.70

56.89

60.1;

63.40

62.74

62.09

56.34

50.59

43.73

36.88

33.61

0.02

0.02

0.02

0.03

0.04

0.05

0.06

0.09

0.11

0.11

0.12

0.12

0.11

0.12

0.13

0.15

0.17

0.18

0.20

0.22

0.24

0.27

0.29

0.29

30.33 .._ 0.30 30.33 0.30

<0.615

9.59:

6.96 8.51

9.65 10.19 9.37

4.25 27.08

4.55

4.74

5.oi

4.90

4.70

4.89

5.48

5.68

5.82

5.87

5.92

5.70

5.48

7.42 11.14

10.35 11.68 11.99 ____.__

12.99

14.10

20.64

29.76, 6.29

33.00

34.73

36.25

36.77

38.99

37.22

38.97

36.72, 5.44

35.93, 5.41

33.20

3l.zj!, 4.39

31.29

31.33

30.95

28.76

4.40

21.27 25.39

4.16 4.18

24.13

10.8:

CO.1
0.41 0.17 0.14 0.00 0.12 0.38 2.37 3.27

0.02

0.21

0.20

0.08

0.06

0.00

4.48

5.80

6.40

8.36

6.32

5.44

4.57

4.14

3.71

3.03

2.35

1.25

0.15

0.18

0.21

o.ri

0.08

0.06

0.09

7.32

7.04

5.05

3.16

1.97

2.71

2.92

mM

DOC

0.31

0.16

0.36

0.39

0.32

0.35

0.29

0.07

0.00

0.14

0.26

0.38

0.49

0.37

0.25

0.68

1.10

0.73

0.37

0.38,

0.38

0.43

0.48

1.66

1.61

co.4

~0.4.

<0.4

<0.4

co.4

co.4

co.4

co.4

co.4

<0.4

~0.4

co.4

co.4

co.4

co.4

co.4

co.4

co.4

<0.4

co.4

co.4

co.4

co.4

cO.4'

co.4

1.28

1.80

2.42

3.30

3.79

3.54

5.30

4.73

4.64

4.91

4.90

4.99

4.92

5.1.5

5.44

4.74

5.95

5.5:

5.lC

4.4i

5.0:

4.s

3.9'

3.3!

3.8:

3.8.

5.44 co.4

19.91

54.45. 5.35

86.48

15.98

16.11

7.51

1.92

0.00,

0.00

pM

AC

1.81


0.27

0.09

FM

SO4

pM

PO4

24.13

10.26

10.68' IO.+!

57.21, co.015 10.98

852.7. 3.6.77 1Ol.t3 798.3; 0.49' 81.63

14.70; 9.37 14.44

9.44 ” 21.79 19.83 11.05 _

7.88

6.33

39.19

9.49, 12.65' 22.79

7.06 8.28 .

5.91.

4.74

4.74.

18.21; 24.28

37.28: +4,33.89

i.28

9.83.

9.10

1.99 2.91

8:.!4

mM

4.54

CT

cl pM_

~0.015 10.25

0.03

0.02

0.03

0.05

0.08

co.016

66.6:3. 422.8: 0.65

49.59

Zn

PM

73.62

22.84' 465.01 16.21 2i8f 3 1iQ.i' 1.40

8.07

4.55' 10.25! - . 13 ..67 ..

1.45

3.92

2.94:

36.88

16.52

2.34

18.61

28.491 2.15

Si

FM

Al

FM

PM

Ca

19.18 469.1; 11.59 lk.Qt1. 76.02

14.44 272.6

131.0; 2.69.

0.40

11.18

!,cs

1.38

0:41

0.46

9.65;

K PM

3.39. 17.09 25.89: 2.43

.

20.62

1.73, 2.30

PM

NH4’

2.55,

0.85

0.58

Na NM

Fer

:

pM

w

Fe(lll Fe(ll)

4.28

4.26

4.31

8.59

.5.64' 8.49

5.29. 5.16

415

Es

4.96

295

355

ES

4.66

E9

._

235

E9

j :

175

E9

6.95

4.62

4.es.

4.38

4.23

. 4.32

55

degasse

115

in situ

pH

E9

cm

pH

E9

profiledepth

4ppendix: pore water composition measured in a profile at EGr (profile E9, October 1994) and at TGe (profile T12, October 1994). The pH value was measured on a fresh sample aliquot (in situ pH) as well as after sparging the samples with N2 gas for 10 min (degassed pH). Fe(III) and Fe (II) concentrations were calculated from the measured total Fe concentration and from typical Fe(lII)/Fe(II) ratio measured in other pore water profdes from the same location. Aluminum, Si, Fe, Mg, and Ca were measured using ICP-AES; DGC was measured by a Shimadzu total carbon analyzer; all others - including CT (total dissolved CO$ - were determined using ion chromatography. AC: Acetate; Italics: interpolated values.