thermodynamics and organic matter

8/17/2019 Thermodynamics and Organic Matter

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assumed that the process with the highest standard

energy yield predominates. For C limited systems, how-

ever, this concept is probably too simplistic (Postma and

Jacobsen, 1996; Peiffer, 1999; Blodau et al., 1998).

Generally, in systems open to the atmosphere the C

mineralization process is driven by the thermodynamic

instability of the organic matter. Within this process,however, a partial thermodynamic and solubility equili-

brium of the terminal electron accepting steps may

occur and control the interface between SO4 and Fe

reduction zones, as was suggested by Postma and

Jacobsen (1996).

In this study the authors examine how the S- and Fe-

turnover in acidic mine lake sediments is affected by the

occurrence of partial thermodynamic equilibria and the

presence of labile organic matter. To these ends a simple

thermodynamic model is formulated and combined with

a rate expression that connects the quality of organic

matter in the sediments with SO4 and Fe reductionrates. This resulting model is tested with empirical data

that have been presented earlier (Blodau et al., 1998,

2000; Peine et al., 2000), and its implications for the

management of highly acidic waters are discussed.

2. Theory

2.1. Biogeochemical processes

A simple thermodynamic model of the biogeochem-

ical processes involved in the Fe sulfide accumulation

process is adopted (Fig. 1).

Hydrolysis of complex organic matter and fermenta-

tion of dissolved organic matter into small organic

molecules precede the utilization of C in SO4 and Fe

reduction (Fenchel et al., 1998). It is assumed that a

thermodynamic equilibrium with respect to fermenta-

tion is not attained, since the reaction products are uti-

lized by SO4 and Fe reduction. This concept is supportedby very low concentrations of fermentation products in

sediments which usually occur on the nano- to micro-

molar scale (Lovley and Goodwin, 1988; Novelli et al.,

1988; Chapelle et al., 1995). The hydrolysis/fermentation

step is also reduction rate limiting. Initially, SO4 and

Fe(III) oxides occur at millimolar concentrations and

are usually not rate limiting in sediments of unproduc-

tive acidic mine lakes (Peine and Peiffer, 1996, 1998;

Blodau et al., 1998). The reactivity and concentration of

the organic matter is assumed to control the fermenta-

tion, and the sum of SO4 and Fe reduction rates (Blodau

et al., 2000) (Fig. 1).Ferric iron reduction, SO4 reduction and H2S oxida-

tion with Fe(III) oxides as electron acceptors are subject

to thermodynamic constraints. These processes might

approach a partial thermodynamic equilibrium (Postma

and Jacobsen, 1996). Shifts in Gibbs free energies (G)

of these processes are assumed to control the ratio of

SO4 to Fe reduction rates and the occurrence of H2S

oxidation (Fig. 1). Acid base reactions and the pre-

cipitation of FeS are kinetically fast and assumed to be

in thermodynamic equilibrium. The latter assumption is

based on SO4 reducing environments frequently being

near equilibrium with respect to FeS (Wersin et al., 1991;

Perry and Pedersen, 1993).

2.2. Thermodynamic relations

To relate the redox, acid base, and precipitation pro-

cesses that might affect the accumulation of Fe sulfides,

two general thermodynamic expressions will be derived.

Beginning with the reduction of SO4 in reaction (4)

acetic acid stands for a variety of utilizable compounds:

CH3COOH þ SO24 þ 2 H

þ ) 2H2CO3 þ H2S ð4Þ

For Fe(III) reduction the respective expression is:

CH3COOH þ 8 FeOOH ðsÞ þ 16 Hþ ) 2 H2CO3

þ 8 Fe2þ þ 12 H2O ð5Þ

H2S and Fe2+ may then precipitate as FeS (reaction 3).

Although not explicitly considered in this derivation, it

should be noted that the nature of the Fe oxides, as well

as their stoichiometric composition, varies in the investi-

gated type of sediment (Peine et al., 2000). The Gibbs free

energy G of a redox reaction can be calculated by Eq. (6)

(Langmuir, 1997):Fig. 1. Conceptual model of the biogeochemical processes and

controls in the sediments. See text for details.

26 C. Blodau, S. Peiffer / Applied Geochemistry 18 (2003) 25–36


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G ¼ G þ RT ln ð½aa ½bb ½cc ½ddÞ ð6Þ

with G: molar standard Gibbs free energy [J]; R: Gas

constant [J mol1 K1]; T: absolute temperature [K];

[a],[b],[c],[d]: reactant activities [mol l1] and a,b,c,d: stoi-

chiometric coefficients.

The G difference between two redox reactions canbe expressed as

G1 G2 ¼ G1 G

2

þ RT ln ½aa ½bb ½cc ½dd ½ee ½f f ½gg ½hh

ð8Þ

Equation (8) can be transformed to Eq. (9) by filling in

Eqs. (4) and (5), standardizing on electron equivalents,

and writing in logarithmic notation:

GFe GSO4 ¼ GFe G

SO4

þ RT2:3 log Fe2þ

þ 1=8 log SO24

1=8 log H2S½ þ 7=4 pHÞ

ð9Þ

Sulfate and Fe(III)-reducing bacteria utilize the same

range of organic substrates (Lovley and Phillips 1988;

Jørgensen, 1983; Coleman et al., 1993). Hence, organic

substrate and H2CO3 do not appear in expression (9).

The reoxidation of H2S to SO4 can be described by

8 FeOOH sð Þ þ H2S þ 14 Hþ ) 8 Fe2þ þ SO24

þ 12 H2Oð10Þ

The reoxidation reaction of H2S t o S O4 is hypo-

thetical. Only a reoxidation process consisting of at least

two steps, involving S as an intermediate product, has

been documented to the authors’ knowledge (Peiffer et al.,

1992; Thamdrup et al., 1993). As an alternative to reaction

(10), reaction (11) is thus considered and referred to when

necessary:

2 FeOOH sð Þ þ H2S þ 4 Hþ ) 2Fe2þ þ S sð Þ

þ 4 H2Oð11Þ

The reoxidation of H2S t o S O4 [reaction (10)] can, based

on electron equivalents and in logarithmic notation, bedescribed by

GS ¼ GS þ RT 2:3 log Fe

2þ

þ 1=8 log SO24

1=8 log H2S½ þ 7=4 pHÞ ð12Þ

Note that expression (12) is equivalent to expression (9).

The fast reactions, considered to be in equilibrium are

FeS sð Þ þ Hþ ¼ Fe2þ þ HS ð13Þ

HS þ Hþ ¼ H2S aqð Þ ð14Þ

The mass action expressions for these reactions are

(Stumm and Morgan, 1996)

Log K FeS ¼ log Fe2þ

þ log HS½ þ pH ¼ 2:95

ð15Þ

Log K 1H2S ¼ log HS½ pH log H2S½ ¼ 7:01

ð16Þ

Expressions (15) and (16) can be combined to

Log H2S½ ¼ log K FeS log K 1H2S log Fe2þ

2 pH

ð17Þ

and be used to control [H2S] in the thermodynamic

expressions (9) and (12). This results in a thermo-

dynamic expression under solubility equilibrium and

Fe(II)-rich conditions:

GFe GSO4 ¼ GS ¼ GS

þ RT 2:3 1=8 log K 1H2S 1=8 log K FeSð

þ 2 pH þ 9=8 log Fe2þ

þ 1=8 log SO24

ð18Þ

The thermodynamic expressions for the Gibbs free

energy difference between SO4 and Fe reduction (9) and

the Gibbs free energy of sulfide oxidation (12) are

equivalent. Hence GS describes the thermodynamic

state of both phenomena.

The oxidation of H2S to S [Eq. 11] can analogously

be derived and described by:

GS0 ¼ GS0 þ RT 2:3 1=2 log K 1H2S 1=2log K FeS

þ 3 pH þ 3=2 log Fe2þ

ð19Þ

GS and GS0 were calculated from thermodynamic

data according to Bigham et al. (1996); Langmuir (1997)

and Stumm and Morgan (1996) with GS= 60.5 KJ

eq1 for Schwertmannite (Fe8O8(OH)x(SO4) y; Bigham et

al., 1996), GS=43.5 kJ eq1 for Goethite, GS0=

54.4 kJ eq1 for Schwertmannite, and GS0=36.4

kJ eq1 for Goethite.

The stoichiometric coefficients in Eqs. (18) and (19)indicate that under the chosen assumptions GS is

mainly controlled by the pH, the activity of dissolved

Fe(II), and the nature of the Fe(III) oxides. The latter

can change the magnitude of GS and GS up to 30

kJ eq1, based on available thermodynamic data (Big-

ham et al., 1996; Stumm and Morgan, 1996). In con-

trast, the activity of SO4 is of little significance with

respect to GS.

For GS)0 a simultaneous partial thermodynamic

and solubility equilibrium is attained (Postma and

Jacobsen, 1996). Expression (18) states that the simul-

taneous partial and solubility equilibrium represents the

C. Blodau, S. Peiffer / Applied Geochemistry 18 (2003) 25–36 27


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state at which the model system will operate at the

highest possible rate of sulfide accumulation, given a

fixed supply of organic substrates. At that state both Fe

and SO4 reduction provide the same energy gain per

transferred electron equivalent. Since both Fe and SO4reducers cannot outcompete each other they should

operate at similar rates, if Fe oxides and SO4 are notrate limiting. In contrast, the reoxidation of H2S with Fe

oxides becomes infeasible at this point, since GS=0.

Hence, the model predicts that at GS=0, FeS pre-

cipitates and that the Fe sulfide accumulation rate

should be maximal.

With increasing magnitude of GS, positive or nega-

tive, conditions become unfavorable for the accumula-

tion of Fe sulfides. In the case of SO4 reduction being

thermodynamically favored (GS0), the system shifts

to sulfidic conditions, because Fe reducers are out-

competed. Hence, little or no Fe(II) will be supplied,

and if no other Fe(II) source is available, no FeS willprecipitate. In the case of Fe(III) reduction being

favored (GS0), the supply of H2S might cease (Lov-

ley and Phillips, 1988) and H2S itself will be chemically

reoxidized upon contact with Fe oxides (Peiffer et al.,

1992).

In natural systems the concentration conditions for a

simultaneous partial and solubility equilibrium are not

precisely known. This is due to the variable nature of

Fe(III) oxides which may represent a mixture of several

minerals, leaving the exact magnitude of GS and

GS0 in expressions (18) and (19) unknown. This

uncertainty requires the consideration of a GS range

so long as the prevailing type of Fe oxide is obscure.

2.3. Simulation of sulfate, iron reduction and

neutralization rates

To describe the system’s ability to oxidize organic

matter by SO4 and Fe reduction (Fig. 1) a first-order

rate expression is formulated that incorporates the ori-

gin and age of the deposited C. Under C limited condi-

tions, the sum of SO4 and Fe reduction rates is likely

constrained by the concentration of reactive AOC, and

secondarily by refractory non-AOC. The latter probably

does not sustain high reduction rates (Blodau et al.,2000). It has also to be considered that the deposited

AOC becomes increasingly recalcitrant over time (Wes-

trich and Berner, 1984; Middelburg, 1989).

The age of the sediment layers was determined from

bulk density profiles, which indicated the onset of sedi-

mentation after flooding of the open pits through a dis-

tinct bulk density jump, and assuming constant

deposition rates as described in Blodau et al. (2000).

The AOC and non-AOC concentrations in the sedi-

ments were approximated by Eq. (20) using d13C and C/

N signals, which were distinct for AOC and non-AOC

(Blodau et al., 2000).

13 OCð Þ j ¼ x j 13 AOCð Þ þ 1 xð Þ j

13 non-AOCð Þ ð20Þ

with 13(OC) j : 13 measured signature of the organic C

in sediment layer j (unit: %); x j : relative share of AOC

in layer j ; d13(AOC): 35.8%; d13 (non-AOC): 26%.

The sum of SO4 and Fe reduction rates was fitted to thefirst order expression (21) containing the concentration of

AOC and non-AOC and the age of the respective layer.

Rmodel; j ¼ k1 t1dep; j cAOC; j þ k2 cnon-AOC; j ð21Þ

with Rmodel, j : sum of SO4 and Fe reduction rate in

sediment layer j (unit: meq cm3 a1); k1, k2: transfer

coefficients (unit: a1); tdep, j : attenuation factor equal-

ing the age of a layer j in years (without unit); cAOC, j :

volumetric AOC concentration in layer j (unit: meq

cm3

); cnon-AOC, j : volumetric non-AOC concentrationin layer j (unit: meq cm3).

GS was used to identify SO4 reducing and Fe redu-

cing zones in the sediments as described in the previous

section. The redox zonation obtained from the sum of

SO4 and Fe reduction rates as generated by (21), and the

stoichiometry as presented in Eqs. (1)–(3) were used to

convert the SO4 reduction rates into neutralization

rates. The precipitation of FeS or FeS2 was assumed

and the transformation of FeS to S was not considered

because its effect on the neutralization rates is minor.

For the model it was further assumed that a simulta-

neous partial thermodynamic and solubility equilibrium

was present in a 20 kJ eq1 window around GS=0,

due to the uncertainty about the nature of the Fe oxides,

and due to the fact that a minimum energy quantum is

required to affect the competition between microorgan-

isms (Hoehler et al., 1994). Due to the absence of

empirical data this value is poorly constrained and will

require more experimental work to build up confidence. In

first approximation it was assumed that within this win-

dow SO4 and Fe reduction coexisted at equal rates, on an

electron equivalent basis, and that a reoxidation of sulfides

did not occur or was slow. At GS20 kJ eq

1 by SO4 reduction.For the generation of total inorganic reduced S

(TRIS) concentration profiles steady-state compaction

and time-invariant thermodynamic conditions since the

beginning of sedimentation were assumed, and changes

in C concentrations due to mineralization were not

considered (Blodau et al., 2000). Having parameterized

Eq. (21), the TRIS concentrations were calculated from

Eq. (22) for each sediment layer with a time step of 1

year for the time period since deposition. This time per-

iod was inferred from the estimated age profiles of the

sediments. The effect of compaction on the C concentra-

tions is accounted for by the ratio Bdi /Bd

n in expression



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(22) which reduces the measured C concentration in a

layer to the respective one at an earlier time interval.

cTRIS; j ¼Xni ¼1

k1t1dep;i cAOC; j

Bdi

Bdnþ k2 cnon-AOC; j

Bdi

Bdn

t and if Gs; j ;i 5

20 kJeq

1

ð22Þ

with cTRIS, j : concentration of TRIS in sediment layer j

(unit: meq cm3) i : age of sediment layer (unit: a); n: time

period since deposition of sediment layer j (unit: a); Bdi :

presently observed bulk density at time i (unit: g cm3);

Bdn: presently observed bulk density at time n (unit: g

cm3); t: time period for which rates are assumed to

be constant (unit: a).

From the simulated TRIS concentration, the esti-

mated age of the sediment layers, and the stoichiometry

in expressions (1)–(3), an average long-term neutraliza-

tion rate due to precipitation of FeS or FeS2 was esti-mated for each layer, integrated across the depth profile

(23) and standardized on the sediment surface area.

Albeit being based on relatively crude approxima-

tions, the generated rates and TRIS concentrations are

instructive in so far as they allow for a quantitative test

of the overall conceptual model against empirical data.

The generated TRIS profiles incorporate the thermo-

dynamic state, the influence of decreasing C reactivity

with time, the different age of individual sediment lay-

ers, and the compaction of the sediments. The combi-

nation of these factors would otherwise obscure the

interpretation of differences in TRIS concentrations

within and between sites.

3. Sites, experimental methods and sediment properties

3.1. Sites

The lakes 77, 76, 116, and Ausee are strip mining lakes

located in Brandenburg and Bavaria (Germany).

Groundwater flooding of the mining areas started in

1965–1968 (76, 77, 116) and 1982 (Ausee). The maximum

water depths of the lakes are 5 m (76), 8 m (77) 11 m (116),

and 25 m (Ausee). All lakes showed a dimictic regimeover the sampling periods, with a clinograde O2 profile in

76, 77, and 116 and an orthograde O2 profile in Ausee.

During summer stratification, the pH was between 2.8

and 3.2 in the epilimnion. The surface waters were SO4rich with concentrations ranging from 1.5 to 12 mmol

l1. Dissolved total Fe concentrations in the epilimnion

varied from 0.3 to 2.0 mmol l1 (Peine, 1998).

3.2. Methods

Sediment cores were taken with a gravity corer. The

cores were cut into segments, placed in N2

flushed bags,

and frozen prior to the solid phase analyses. Diffusion

chambers (Hoepner, 1981) were used to sample the pore

water of the sediments. Seasonal variability was recorded

by sampling the pore water of site 77 in May, August,

November, 1996, and in February, 1997. Fe2+ con-

centrations and pH were determined immediately, while

the other subsamples were frozen (18 C) and storeduntil analyzed. SO4

2 was determined by ion chromato-

graphy, and Fe2+ by the phenanthroline method

(Tamura et al., 1974). Total Fe was determined by flame

atomic absorption spectrometry after digestion of the

dried sediment with concentrated HNO3 and HCl acid

(1:1 ratio) in a microwave digester. The nature of the Fe

oxides was determined by X-ray diffraction and has

been described in detail elsewhere (Peine et al., 2000).

The content of total inorganic reduced S compounds

(TRIS: FeS2, FeS, S) were determined by the method

of Fossing and Jørgensen (1989). Frozen sediment sam-

ples were thawed under N2 and distilled with HCl (c=5mol l1) and CrCl2 (c=0.15 mol l

1). The H2S released

into the N2 stream was trapped in 50 ml of NaOH

(c=0.15 mol l1) solution. The sulfide was precipitated

by addition of zinc acetate and photometrically deter-

mined. In 2 of 3 cores from lake 116 TRIS contents were

estimated from the difference of total S, dissolved SO4and the C contents assuming a C/S ratio of 100 (Jør-

gensen, 1977). S was extracted from fresh sediment by

methanol and measured by HPLC and UV detection

(Ferdelmann et al., 1991). Carbon and N and total S

contents of the sediments were determined with a C/N/S

analyzer after drying the sediment samples. Sulfate

reduction rates were measured by the 35S-radiotracer

technique (Jørgensen, 1978). Ferric iron reduction rates

were estimated by the closed-vessel incubation technique

(Roden and Wetzel, 1996) and pore water modelling of

dissolved Fe(II) profiles (Blodau et al., 1998). Activity

coefficients for the dissolved species were estimated by

the extended law of Debye-Hueckel. 13C /12 ratios (d13C

notation, unit %) were determined by gas–mass spec-

troscopy (thermal decomposition) after homogenizing,

freeze drying and grinding of the sediment samples.

The sediment age was estimated from finely sectioned

sediment cores. The position of the former mine ground

was indicated by an abrupt change in bulk density withdepth. The approximate age of the sediments were esti-

mated from the dry weight that had been deposited

from that point on, assuming constant deposition rates.

The sediment age and the solid phase concentrations of

C, total and reactive Fe and TRIS were used to deter-

mine average deposition rates since the beginning of

flooding. The dating was checked by 137Cs dating of a

finely sectioned sediment core in lake 116 and a com-

parison to data from sediment traps which were sam-

pled in lake 77 on a regular basis for one year (Peine et

al., 2000). 137Cs activities were determined by gamma

spectrometry after freeze drying of sediment material.



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3.3. Sediment properties

Reactive Fe contents in the sediments ranged from

about 0.1 to 7 mmol g1 (dry weight). Total Fe contents

ranged from less than 0.1 to 15 mmol g1. X-ray dif-

fraction of site 77 samples showed that the upper cen-

timeters were dominated by schwertmannite and thatbelow that depth this mineral was absent or of minor

importance compared to goethite (Peine et al., 2000). At

site 116 crystalline Fe oxides probably predominated

(Blodau et al., 1998). TRIS contents (Fig. 2) ranged

from 0.002 to 2.1 mmol g1 (


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a very small fraction of the total C that had been

deposited. At site 76 C and Fe deposition rates and

TRIS accumulation rates were all relatively small. At

site 116 low Fe deposition rates, moderate C deposition

rates, and high TRIS accumulation rates coincided.

AOC accounted for a larger fraction of the total C that

had been deposited, when compared to site 77.

4. Results and discussion

4.1. Redox zonation

Site 77 was characterized by a shift from GS0 to

GS0 with increasing sediment depth (Fig. 2).

According to the adopted model the upper zone should

be Fe reducing and sulfide oxidizing, and should not

accumulate Fe sulfides. This prediction corroborates

with the empirical data (Fig. 2). In the upper zone both

TRIS and FeS contents were very low. A similar ther-

modynamic pattern prevailed in the sediments of lake

Ausee, in which no TRIS had accumulated (Fig. 2). At

greater depths at site 77, thermodynamic conditions wereclose to partial equilibrium (GS 0) (Fig. 2). According

to the model in this zone both SO4 and Fe reduction

should operate at similar rates and Fe sulfides should

accumulate. This was the case: both processes coexisted

(Fig. 4) and Fe sulfides had accumulated (Fig. 2).

In the presence of non-crystalline Fe oxides, similar

thermodynamic conditions (GS 0) may also have

prevailed in the upper layers of the sediments of lake 76

(Fig. 2). The low rates of Fe reduction, which can be

inferred from the absence of significant Fe(II) con-

centration gradients (Fig. 2), suggest that in this sedi-

ment SO4

reduction (Fig. 4) clearly dominated. This

would corroborate with the high GS values in presence

of goethite (Fig. 2).

In contrast to site 77, at site 116 the shift from

GS0 to GS0 occurred in the centimeter below the

sediment–water interface (Fig. 2). Below that depth

Fe(III) and SO4 reduction coexisted, with SO4 reduction

predominating in terms of electron equivalents, and Fe

sulfides accumulated.

As already noted, the reoxidation reaction of H2S to

SO4 is hypothetical. Only a reoxidation process consist-

ing of at least two steps, involving S as an intermediate

product, has been documented (Peiffer et al., 1992,

Thamdrup et al., 1993). The authors used Eq. (19) to

estimate GS. In the partial equilibrium zone of site 77,

GS (Eq. 18) was close to 0 kJ eq1 (4 to +12 kJ

eq1) and GS (Eq. 19) was clearly positive (+12 to

+20 kJ eq1). Hence, the inaccuracy in the thermo-

dynamic description of the sulfide oxidation process in

Eq. (18) does not compromise the analysis with respect

to the impossibility of H2S oxidation in the partial

thermodynamic equilibrium zone. This also holds true

for the fact that most of the TRIS in the sediments was

FeS2, while in the model this species does not appear(Eq. 18). It is reasonable to assume that initially FeS

precipitated and controlled the H2S activity in the pore

waters and was subsequently transformed to FeS2(Wersin et al., 1991; Peiffer et al., 1992; Perry and Ped-

erson, 1993; Rickard, 1997). Moreover, as can be seen

from the stoichiometric coefficients in Eq. (18), the effect

of the solubility product of Fe sulfides on G is minor

when compared to the effect of pH values and Fe(II)

activities. Even if the assumption that the precipitation

of FeS controlled the H2S activity were not met in the

sediments the effect of this error would be minor and

not influence the conclusions.

Fig. 3. Input data for the rate related expressions (20)–(22). These are the molar C/N ratios, organic C concentrations, d13C values,

and the estimated age of the sediments. Error bars indicate standard deviations ( n=4).



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Table 1

Deposition rates and accumulation rates of TRIS

Period Site C (g m2 a1) AOC

(g m2 a1)

Reactive Fe

(g m2 a1)

Total Fe

(g m2 a1)

TRIS

(g m2 a1)

1968–1996 77 165 (56–300) 3–6 81–98 196 0.1–1.3

76 23 – 64 64 1.6

116 96 (86–106) 5–10 57 57 8.2 (1.3–17.1)

1986–1996 Au 132 – 35 110 0.06

77 120 6–13 145 360 0.5

76 23 – – – –

116 51 10–19 20 146 12.1 (1.6–22.1)

Deposition rates of AOC were calculated by estimating AOC concentrations with Eq. (20). Values in brackets show the variability of

between replicate cores.

Fig. 4. Measured and stimulated short-term SO4 and Fe reduction rates and TRIS (in electron equivalents) in the sediments.

‘‘Dashed’’ TRIS concentration profiles represent 5 a intervals beginning after flooding of the lakes. Error bars indicate standard

deviations (n=3 or 4). For details see text.



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4.2. Sulfate, iron reduction and neutralization rates

The rate model (21) was parameterized using the data

of site 77 (k1=0.61 a1; k2=0.011 a

1) and reproduced

both the magnitude and the shape of the turnover pro-

files (Fig. 4). The parameterized model from 77 also

adequately simulated the magnitude of Fe and SO4reduction and short-term neutralization rates when

applied to site 116 (Fig. 4, Table 1). It underestimated,

however, the SO4 reduction and short-term neutraliza-

tion rates at site 76 (Fig. 4, Table 1). It has to be con-

sidered, though, that the rate calculation for site 76 was

based on C concentrations from an individual core

which had a lower sediment thickness than was

observed on average in that lake. Averaged results from

multiple cores, as used for the other sites, might have

improved the corroboration between simulated and

empirical data for site 76.

The magnitude of TRIS concentrations and theresulting long-term neutralization rates were fairly well

simulated using expression (22) (Fig. 4, Table 2). The

model produced a TRIS peak close to the sediment–

water interface at site 116 (Fig. 4), which was caused by

high concentrations of AOC at that depth. This peak

also occurred in the empirical data. In the model, the

current high deposition rates of AOC will sustain high

neutralization rates at that site if the current conditions

are projected into the future (Fig. 4, 116).

In contrast to site 116, smaller quantities and less

reactive AOC reached the SO4 reduction zone in 77. At

site 77 most of the decomposed C was ‘‘lost’’ to Fe

reduction. This is also illustrated by the discrepancy

between the measured and simulated TRIS concentrations

that accumulated in the model when the thermodynamic

partial equilibrium at site 77 was (hypothetically) exten-

ded to the sediment–water interface (Fig. 4, ‘‘77 partial

equilibrium’’). The overestimation of the simulated

TRIS concentration in the zone between 6 and 15 cm

depth at site 77 (Fig. 4) could be eliminated by assuming

yearly fluctuations of the redox zonation by 3 cm, as

presently observed at that site (Fig. 2, Fig. 4, ‘‘77 redox

fluctuation’’).

The overall adequate agreement between model and

empirical data suggests that the relative rates of Fe and

SO4 reduction and the accumulation of TRIS are con-

trolled by the presence or absence of a partial thermo-dynamic and solubility equilibrium. Once such an

equilibrium is established it can probably be fairly per-

sistent since a steady-state approach seemed to be suffi-

cient to reproduce the observed TRIS-patterns in the

sediments. In the presence of such an equilibrium, the

variation in TRIS concentrations can be explained by

variation in the concentration, age and origin of organic

C, as was demonstrated above. In the presence of a

partial thermodynamic and solubility equilibrium, rates

of neutralization can hence be modelled using a kinetic

that is solely based on changes in the availability of

organic C. The kinetic expression underlying such amodel can be formulated by using one or two transfer

coefficients, which control the rate at which organic

matter is fermented and utilized by SO4 and Fe reduc-

tion. These transfer coefficients can be explicitly esti-

mated by relating turnover measurements to organic

matter quality parameters as carried out in this study, or

by the inverse modeling of concentration profiles, using

a diagenetic model (Berner, 1980).

4.3. Implications for lake management

The previous discussion shows that the sediments can

be separated into a sulfide oxidizing and an Fe sulfide

accumulating type. These two types will respond differ-

ently to the additional deposition of organic matter

proposed as a means to accelerate the neutralization

process (Fyson et al., 1998). The Fe sulfide accumulat-

ing type (site 116) will effectively increase TRIS accu-

mulation rates, as long as the supply of SO4 from the

water column does not become transport limited. Such a

Table 2

Neutralization rates (mmol m2 a1) due to SO4 reduction and precipitation of FeS/FeS2 in the sampled lakesa

Lake Short term (in 1996) Long term (1968–1996)

Simulated rates

mmol m2 a1Measured rates

mmol m2 a1Simulated rates

mmol m2 a1Measured rates

mmol m2 a1

Au 0 n.d. 0 4

77 533 698 (481. . .902) 54b. . .132c 50 (9. . .81)

76 487 1287 (1079. . .1837) 163. . .187c 96. . .225

116 1055 897 (482. . .1401) 225. . .258c 521 (76. . .1034)

a Long term rates are based on the time period after flooding. Ranges for the measured rates are due to spatiotemporal variations.b Assuming a vertical redox fluctuation of 3 cm around the measured average depth of the sulfide-oxidizing/sulfide-accumulating

boundary.c Presence of either FeS or FeS

2.



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limitation has been reported from highly eutrophic mine

lakes (Hupfer et al., 1998).

A similar response, however, cannot be expected from

the sulfide oxidizing type (sites Ausee and 77). These

sediments are sulfide oxidizing systems which are prob-

ably internally stabilized: due to the presence of reactive

Fe, SO4 reducers are thermodynamically poorly compe-titive in the first place. The slow transformation of

schwertmannite to goethite, as occurring at site 77

(Peine et al., 2000), releases H+ (Bigham et al., 1996).

The reduction of schwertmannite does not consume H+

either (Peine et al., 2000). Thus the pH in the pore water

stays around 3 (Fig. 2, Ausee, 77). The low pH is,

according to Eq. (18), the main factor which thermo-

dynamically favors Fe reduction and sulfide oxidation,

and suppresses SO4 reduction. A mechanism that could

increase the pH is hence missing. The sediments, there-

fore, remain in an Fe reducing and sulfide oxidizing

state, as long as Fe and SO4 reduction are still C limited.Consequently, moderate fertilization measures that

avoid a strong eutrophication of the surface waters may

be ineffective. If an additional C source was added to

this type of sediment it should be less readily decom-

posable in order to reach the SO4 reduction zone after a

few years. The effectiveness of this addition will still be

comparatively low, as long as substantial amounts of

schwertmannite are present in the sediment.

If the sediments in an acidic lake are predominantly

sulfide oxidizing it might be more efficient to reduce the

supply of reactive Fe from the water column. This

would cause GS to increase more rapidly beneath the

sediment water interface, as in lake 116, and increase the

efficiency of the TRIS accumulation process (Fig. 4, site

77, ‘‘partial equilibrium’’). To these ends the dissolved

Fe input into the lake water should be decreased. This

would result in a dampening of the seasonal Fe cycling

in the surface waters, which results in the deposition of

Fe(III) after the summer stratification (Peine et al.,

2000).

This reasoning is also to some extent supported by the

Fe, C and TRIS deposition data of the investigated

lakes. The main difference in the biogeochemistry of

lake 116 and lake 77, which are otherwise similar in size,

depth, age, and surface water chemistry, is the deposi-tion of reactive Fe to the sediments. This deposition is

currently on average almost an order of magnitude

higher in lake 77 compared to lake 116 (Table 1).

Moreover, the reactive Fe deposition rate has increased

over time in lake 77, whereas it decreased in lake 116

(Table 1). Reactive Fe deposition rates to the younger

and deeper Ausee sediments were, on the other hand,

comparatively low not yet allowing for a switch to a

sulfide accumulating state. However, more survey data

on dissolved Fe concentrations in the lake water, reac-

tive Fe sedimentation rates and nature of the Fe oxides

in various lake systems are necessary if the mechanistic

understanding that is gained from this study is to be

validated on a broad basis.

Perhaps the most significant suggestion of the study is

that once the reactive Fe supply to the sediments of

highly acidic waters decreases and AOC is deposited,

the neutralization process will be accelerated by a posi-

tive feedback mechanism. This mechanism is based onthe strong dependency of GS on the pH and the fer-

rous iron activity in the pore waters of the sediments

(Eq. 18). Increasing Fe(II) activity and pH values due to

SO4 and Fe reduction, and the precipitation of Fe sul-

fides, increase GS in the sediment pore waters.

Increasing GS values encourage the establishment of a

partial thermodynamic equilibrium, which, in turn,

optimizes the coexistence between SO4 and Fe reduc-

tion, and enhances the TRIS accumulation rates in the

sediments. Higher TRIS accumulation rates imply the

elimination of acidity and dissolved Fe from the water

column. This might increase the productivity of thesurface waters, for example by increasing the avail-

ability of P in the water column (Kleeberg, 1998), and

thus accelerate the input of AOC into the sediments.

Higher AOC deposition rates will cause larger SO4, Fe

reduction and TRIS accumulation rates, closing the

positive feedback loop.

It is obvious that this biogeochemical feedback

mechanism could be quite important for the long term

development of highly acidic waters and should be

investigated in more detail in future work.

Acknowledgements

We thank D. Arneth and S. Ba ¨ r for technical assis-

tance. This investigation was supported by the German

Federal Minister of Science and Technology.

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