thermodynamics and organic matter
TRANSCRIPT
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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.
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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
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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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