e rno2

en

mark

Municipality of Herlev, Department of Environment, Herlev Bygade 90, 2730 Herlev, Denmark

libria and redox processes occurring at partial equilibrium. 2007 Elsevier Ltd. All rights reserved.

e.g. marine, lacustrine and soil sediments, but due to theimportance of the advective ow in aquifers these processespresent themselves dierently. The Rm study indicated acoupling of inorganic geochemical processes with themicrobiologically mediated redox processes through the

0016-7037/$ - see front matter 2007 Elsevier Ltd. All rights reserved.

* Corresponding author.E-mail addresses: raj@er.dtu.dk (R. Jakobsen),

lisecold@ or.dk (L. Cold).

Geochimica et Cosmochimica Acta1. INTRODUCTION

Groundwater is an important source of drinking water inmany places of the world. However due to pollution fromindustry and agriculture, near-surface aquifers are in manycases abandoned, and water is obtained from deeper oftenanoxic aquifers. Anoxic redox processes are important indetermining natural water quality in these deep-seated aqui-fers (Park et al., 2006), however, detailed knowledge of thesesystems is scarceto some degree due to the large depth

which limits the available data. Several studies have focusedon the microbiology of deep-seated aquifers (e.g. Pedersen,2000; Amend and Teske, 2005), but the main focus here ison the geochemistry of Fe-oxide reduction, sulfate reductionand methanogenesis mediated by microorganisms in theaquifer and the relations between the processes and the sedi-ment. Earlier studies of the Danish anoxic Rm aquifer(Jakobsen and Postma, 1999; Hansen et al., 2001; Larsenet al., 2006) have shown it to be characterized by the samegeochemical processes found in other anaerobic systems,Received 12 May 2006; accepted in revised form 15 January 2007; available online 20 January 2007

Abstract

The study addresses a 10 m deep phreatic postglacial sandy aquifer of vertically varying lithology and horizontally varyinginltration water chemistry, displaying calcite dissolution, ion-exchange, and anaerobic redox processes. The simple varia-tions in lithology and inltration combine into a complex groundwater chemistry, showing ongoing Fe-oxide reduction,sulfate reduction and methanogenesis. Rates of sulfate reduction, methanogenesis and methane oxidation were measureddirectly using radiotracers. Maximum rates were 1.5 mM/yr for sulfate reduction, 0.3 mM/yr for methanogenesis, and only4.5 lM/yr for methane oxidation. The overlap of sulfate reduction and methanogenesis was very small. The important inter-mediates formed during the degradation of the organic matter in the sediment, formate and acetate, had concentrationsaround 2 lM in the sulfate reducing zone, increasing to 10 and 25 lM in the methanogenic part. The concentration of H2was around 0.25 nM in the Fe-reducing zone, 0.4 nM in the sulfate reducing zone, and increased to 6 nM in the methanogeniczone. Using in situ concentrations of products and reactants the available energies for a range of dierent reactions could becalculated. The results of the calculations are in accordance with the observed distribution of the ongoing redox processes,implying that the system is well described using a partial equilibrium approach. A 2D numerical PHAST model of the systembased on the partial equilibrium approach, extended by implementing specic energy yields for the microbial redox processes,could explain most of the observed groundwater geochemistry as an expression of a closely coupled system of mineral equi-Geochemistry at the sulfattransition zone in an a

equilibrium interpretation using

Rasmus Jakobsa Institute of Environment and Resources, Technical University of Den

bdoi:10.1016/j.gca.2007.01.013eductionmethanogenesisxic aquiferA partialD reactive transport modeling

a,*, Lise Cold b

, Bygningstorvet, Bygning 115, DK-2800 Kongens Lyngby, Denmark

www.elsevier.com/locate/gca

71 (2007) 19491966

on water and sediment chemistry from the Asserbo aquifer

matter. However, the Asserbo aquifer consists of shallow

1950 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966marine sands rather than the dune sands on Rm, has ahigher ow rate, a dierent inltration composition, and islithologically less homogeneous. This enables us to furtheraddress how aquifer ow and sediment characteristics as wellas microbial processes control the groundwater chemistry ofan anaerobic groundwater system, increasing the potentialfor transferring the insights from these detailed studies to lessaccessible aquifers. The main features of the Rm aquifercould be modeled using a 1D PHREEQC model (Jakobsenand Postma, 1999) applying the partial equilibriumapproach. This approach is currently only one of severalapproaches used in modeling of redox processes in the sub-surface as there is also the kinetic approach used by Hunteret al. (1998), and various combinations (Jin and Bethke,2005; Brun and Engesgaard, 2002). The latter as well asCurtis (2003) provide reviews and comparisons of some ofthese approaches.More elaborate schemes involving a closermodeling of the microbial dynamics are described byThullner et al. (2005) and Watson et al. (2005). The Asserboaquifer presented here appears to be in a quasi-steady-state,implying that the system will change as the reactivity of theminerals and the organicmatter in the system change, but be-cause reaction rates are low, changes will be slow comparedto the ow rate of the system. Furthermore, with no availableinformation on, e.g. seasonal dynamics, modeling the systemusing the simple partial equilibrium approach used by Jakob-sen and Postma (1999) seems appropriate. Due to thenon-uniform chemical composition of the inltration at theAsserbo site, the system is modeled in 2D using PHAST(Parkhurst et al., 2005). Furthermore, in order to approacha modeling of the levels of the intermediates, H2, acetateand formate, formed during the decomposition of organicmatter in the system, the simple partial equilibrium approachused by Jakobsen and Postma (1999) has beenmodied to in-clude the non-zero energy yield that the microbes mediatingthe redox processes need. The use of the numerical modelmakes it possible to quantitatively address the interactionsbetween inorganic geochemistry and the microbiologicallymediated redox processes occurring in this type of systemthat develops into a very closely coupled system of severalmineral equilibria and terminal electron accepting processescontrolling the groundwater chemistry. This quantitative insightshould be useful when interpreting other anaerobic aquifersystems, where the dataset is normallymuch less comprehensive.

2. METHODS AND SETTING

2.1. Groundwater sampling and analysis

In general sampling was carried out as described in Jakobsenand Postma (1999) and Hansen et al. (2001), the following is aon Northern Zealand, Denmark, describes results from asimilar system, in terms of age and reactivity of the organicthermodynamics of these (Jakobsen and Postma, 1994, 1999;Hansen et al., 2001) and lead to the partial equilibrium ap-proach to describing redox processes (Postma and Jakobsen,1996). The study presented here, based on a detailed data setshort summary of this. Groundwater samples were taken usingstainless steel drive point piezometers. Steel pipe (100) was driveninto the aquifer using a pneumatic hammer. The 6 cm long 50 lmstainless steel mesh screen lter tip was equipped with a check valveto allow sampling by nitrogen gas displacement, taking care tosample the middle part of the water in a pump cycle minimizingeects of degassing. Three sites spaced 10 m apart on a lineintended to coincide with the ow direction were sampled.

For hydrogen sampling, a bundle of 10 mm diameter PVCtubes equipped with a 20 mm disc-shaped 20 lm nylon screen wereinstalled in a hand-drilled well using plastic casing and shoe(Jakobsen and Postma, 1999). To obtain meaningful results, metalparts were avoided and the well was left to rest for two monthsbefore hydrogen sampling commenced (Jakobsen and Postma,1999). Hydrogen was measured in the eld using the bubblestripping method of Chapelle and McMahon (1991), modied,using lower ow, to account for the small intake area as describedin Jakobsen and Postma (1999) and Hansen et al. (2001).

Methane water samples were collected in a syringe, avoidingany air contact, and 3 ml were injected in a pre-weighed 13 mlevacuated blood sample vial, and then frozen upside down at18 C. For all other components, water samples were lteredanaerobically through 0.2 lm lters in the eld. Samples for inor-ganic anions, acetate and formate were collected in 5 mL poly-propylene vials and always frozen immediately to below 18 C inthe eld. Also in the eld, pH, O2 and electrical conductivity weremeasured in a ow cell, alkalinity determined by Gran titration andFe2+ (Stookey, 1970) and H2S (Cline, 1967) were determinedspectrophotometrically. The pH measurements from the two lastsampling positions appear awed, presumably due to a malfunc-tioning pH meter. The data from the rst data set show that calciteequilibrium is obtained at depth. Therefore the pH in the last twodata sets have been adjusted by shifting the data by 0.7 U, in bothcases resulting in an asymptotic approach to calcite equilibriummatching the development in the rst data set rather closely, sup-porting this x.

In the laboratory, cations were measured by AAS, the con-centrations of anions determined by ion chromatography, onsubsamples amended with 4& 0.05 M NaEDTA to avoid precipi-tation of oxidized Fe, and methane by gas chromatography withFID detection on the headspace of the vial after thawing, and theaqueous concentration was calculated from the volume of waterand headspace in the vial. The concentrations of acetate and for-mate were determined using ion exclusion chromatography with aDionex AS-10 column and suppressed EC detection. Samples foracetate and formate were thawed less than 12 h before analysis andboth standards and samples were amended to hold 0.2% chloro-form to minimize both loss by microbial oxidation in standards andan increase in concentrations in samples, presumably due tomicrobial decomposition of DOC (dissolved organic carbon). Thedetection limit for acetate was 0.2 lM but often unidentiedinterfering peaks were observed, leading to a higher eectivedetection limit. For other analytes, detection limits were always,apart from sulde, well below measured concentrations.

2.2. Radiotracer rate measurements

Sediment cores for sediment analyses and radiotracer ratemeasurements were taken in 50 mm ID, 1.5 mm thick, stainlesssteel tubing, with a barrel free corer (Starr and Ingleton, 1992).Core depths were adjusted for compaction (if necessary) bycomparing concentrations in core pore water with correspondingconcentrations in well water samples. Displacements were generallyless than 0.4 m. A separate core was taken for each of the radio-tracers used. After retrieval, the cores were cut in 50 cm sectionsand the ends immediately sealed with plastic stoppers and wrapped

Geochemistry at the sulfate reductionmethanogenesis transition zone 1951with aluminum-foil-tape. The core sections were then lowered in awell for 13 h to bring them back to in situ temperature.

After retrieval of the core sections from thewell, 1 mmholes weredrilled in the tubingwall and 12.525.0 lLof radiotracerwas injectedalong a line across the core at intervals of 1012 cm. The holes wereresealed with aluminum-foil-tape. Then the core was relowered intothewell for incubation. ForCO2 reduction 100150 kBq (12.5 lL) ofH14CO3

was injected, which changed the in situ total inorganiccarbon concentration by1%. For acetate 1.3 kBq of 14CH3COO-Na was injected, which changed the in situ acetate concentration by

1952 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966The piezometric heads, and sampling positions are shown inFig. 1 together with the ow vectors estimated by the PHASTmodel for the 2D cross-section.

3. RESULTS

Three sites 10 m apart (Fig. 1) along a ow line weresampled for water chemistry, and at the central site coreswere taken and geochemical sediment parameters anddirect radiotracer rate measurements were made.

3.1. Groundwater chemistry

The general groundwater chemistry at the three sites, forthe major ions that are not, or only indirectly, aected by

Fig. 1. (a) Site with sampling positions piezometers, heads, near-surface gsection with the calculated ow vectors shown as a base point with a linscales on the map and in the cross-section.redox processes, is shown in Fig. 2 together with the pHand the total alkalinity. The Na and Cl proles show adistinct maximum around a depth of 57 mbs. At least twoexplanations are possible. It could be a result of a higherlevel of dry deposition upstream due to, e.g., higher treesor a release of dissolved ions from high salinity water fromearlier inundation events retained in a layer of low perme-ability. The similar peak in sulfate (Fig. 3) suggest accumu-lation of sulfate as dry deposition or that sulfate from pyriteoxidation in the proposed low permeable layer is releasedwith the high salinity water. Analysis of water sampledfrom the piezometer wells indicate the occasional presenceof high salinity water in the upper groundwater. In thepaleosol indicated in Fig. 1, Cl reached 3.8 mM, almosttwice the highest concentration in the proles, and sulfate

eology and an indication of (b) the 2D vertical PHAST model cross-e indicating velocity and direction (right to left). Note the dierent

Fig. 2. Concentrations of major components in the Asserboaquifer as symbols and the model results as lines. Note the owdirection is from left to right.

Geochemistry at the sulfate reductionmethanogenesis transition zone 1953reached 0.74 mM, similar to what is seen in the proles.Also in the area upstream of the sampling site, high concen-trations were found in the piezometer, though not as highas those found in the groundwater. This implies that thehigh salinity groundwater, is either found deeper, whichwould imply an inundation event, or in a location not sam-pled. The owpattern indicated by the contours of thegroundwater head makes it dicult to pinpoint the areaby backtracking a ow line, but according to the owmodel the area would be outside the area covered by thepiezometers. Though there could be other explanationsfor the anomaly in the water chemistry the modeling de-scribed later assumes the proposed inundation event. Theincrease in calcium at 3.5 mbs presumably corresponds tothe level from where inorganic carbonate, mainly as shellfragments, is present in the sediment. The dissolution ofthese causes an increase in alkalinity along with Ca andMg, present in the calcite shells. This does not cause thepH to increase, because the pH already increased fromaround 5 in the top samples to 7.2 at 4 mbs, presumablyrelated to the reduction of Fe-oxides described below.The patterns of variation seen in the cations, most distinctfor Mg, indicate that the distribution of these are aectedby ion-exchange processes, presumably induced by theintroduction of the high-salinity water.

Data for components directly aected by redox process-es are shown in Fig. 3. The Fe2+ concentration shows thatthe aquifer becomes anaerobic just below the water table.The proles show most of the classical anaerobic redoxsequence; an increase in Fe2+ due to Fe-oxide reduction,followed by depletion of sulfate due to sulfate reductionand an increase in methane from methanogenesis whenthe sulfate is depleted. The upper part of the peak in the sul-fate prole coincides with the Cl prole. The decrease insulfate from 5.5 to 6.5 mbs is accompanied by an increasein sulde from 5.5 to 7 mbs. Sulde disappears or decreasesat depth presumably being xed in an Fe-sulde phase. Thedistinct increase in dissolved manganese which should oc-cur rst in a classical redox sequence appears below thezone of increasing Fe2+ concentrations indicating eitherreduction of very stable Mn-oxides or perhaps more prob-able, when noting the coincidence with the pattern observedfor Ca, a release of Mn contained in some of the carbonategoing into dissolution between 3.5 and 6 mbs.

Concentrations of the intermediates acetate, formateand H2 are shown in Fig. 4. Concentrations of formateare close to constant at site 2 and 3, while at site 1 valuesshow more scatter and an increasing tendency as the systembecomes increasingly reduced with depth. A similar thoughless erratic pattern is seen for acetate. The low concentra-tions, compared to the measured rates (Fig. 6), imply thatthese intermediates have a low residence time and are pro-duced close to the sampled site. The source is presumablyorganic macromolecules present in solution and carboncompounds bound to the sediment. The H2 prole from site2 shows low values around 0.5 nM over most of the prolewith a slight increase over depth to 6 mbs, and below this,the level increases sharply to 7 nM at 6.5 mbs.

Fig. 5. shows the DOC, the d(TIC-Ca) (the increase inTIC from one sample depth to the next corrected by the

1954 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966increase in Ca removing carbonate released from calcitedissolution) and the ammonium concentration. Only a veryminor part of DOC consists of acetate and formate (Fig. 4)and though the porewaters presumably also contain othersmall organic molecules such as lactate, propionate andbutyrate, these are probably not quantitatively important,so that the DOC is dominated by macromolecules. DOCvalues are high in the upper part and the decrease down

Fig. 3. Concentrations of redox-reactive components in theAsserbo aquifer as symbols and the model results as lines. Notethe ow direction is from left to right.to 5.5 mbs correlates with the increase in Ca from the cal-cite dissolution. This indicates that the higher Ca2+ concen-trations aects the mobility of the DOC, perhaps bychanging the adsorption characteristics of the organic mol-ecules. The generally positive values of d(TIC-Ca) indicateoxidation of organic matter over most of the prole, thefew negative values are presumably eects of ionexchange aecting the Ca concentrations. The steady in-crease in ammonium over depth also indicates the oxidationof organic matter as it presumably reects the release ofammonium to the porewater as organic matter is degradedas described from marine sediments by Caneld et al.(1993).

3.2. Measured rates

Rates of sulfate reduction measured by injecting 35S,rates of methanogenesis from the reduction of injected14CH3COOH labeled acetate, and H

14CO3 as well as the

rates of oxidation from the labeled 14CH3COOH, and ratesof methane oxidation measured using 14CH4 are plotted inFig. 6. Oxidation of organic matter appears to be focused inthe 56 mbs interval with sulfate reduction as the mostdominant redox process. The zone of sulfate reductionshows a slight overlap with the zone of methanogenesis.Though rates of methanogenesis are not as high, they are

Fig. 4. Measured concentration of intermediates used in terminalelectron acceptor processes with model results shown as lines. Notethe ow direction is from left to right.

Geochemistry at the sulfate reductionmethanogenesis transition zone 1955within the same range. Both CO2 reduction, acetate fermen-tation and oxidation show a distinct peak just around7 mbs indicating that rates are related to the reactivity ofthe organic matter in the sediment. The acetate oxidationrate is very similar to the CO2 reduction rate, while the ace-tate fermentation rate is considerably smaller. Methane oxi-dation rates are even lower, showing that though theprocess occurs the extent is very limited as only about 1%of the methane appears to be reoxidized.

3.3. Sediment geochemistry

Sediment parameters are shown in Fig. 7. The Feextractable by 6 N HCl acid is 1015 mmol/kg in the upperinterval 57 mbs, and then increases to 25 mmol/kg at7.5 mbs, part of this could be siderite. Suldes are presentin the entire 57.5 mbs interval, but below 6 mbs onlyCRS (mostly FeS2) are found while above about one thirdof the suldes are found as AVS (mostly FeS), showing thatthe conversion of FeS to FeS2 runs to completion in thissystem. For sedimentary organic matter the amounts of

Fig. 5. (Left) The d(TIC-Ca), the change from one observationpoint to the one below, derived from observations, representing theincrease in TIC not related to calcite dissolution. (Center) Theconcentration of ammonium, released from oxidizing sedimentaryorganic matter. Derived or measured values are plotted as symbols,the model results are shown as lines. (Right) The measured totaldissolved organic carbon. Note the ow direction in this plot isfrom top to bottom.ADSOC and NADSOC are about the same, except forthe peaks around 5.7 and 7 mbs where the extra organicmatter appears to be NADSOC. In the examined intervalthe sediment inorganic carbon content (SIC) is relativelyconstant around 250 mmol C/kg, with a single outlier at7 mbs with only 60 mmol/kg, possibly related to the highrate of methanogenesis observed close to this level(Fig. 6), producing extra CO2, causing or having causedextra carbonate dissolution.

4. DISCUSSION

4.1. Mineral equilibrium control

Fig. 8 shows activity plots for the most important min-eral equilibria in the system. For calcite, the plot shows thata very large part of the data plot very close to the linerepresenting calcite equilibrium. These data plot in a densecluster because the pH, the alkalinity and the Ca concentra-tion all become close to constant at depth. The points belowthe line are from the upper part of the aquifer from wherecalcite has either been leached out and is present in suchsmall amounts that equilibrium cannot be obtained withinthe time available for the groundwater owing through.That calcite equilibrium is obtained at depth ts well withthe presence of SIC in the cored interval (Fig. 7). For sider-ite many data points plot along a line representing anSIsiderite = 0.75. Equilibrium control at supersaturations ofthis magnitude, reecting a kinetic inhibition of the ongoingprecipitation that is compensated by the supersaturation,have been described for other systems, e.g. Postma (1981),Jakobsen and Postma (1999) and Jensen et al. (2002). Alsothe nearly constant Ca2+/Fe2+ activity ratio seen at depthindicates siderite equilibrium control. Again the points thatare o the line are from the upper part where carbonateactivities are low. For FeS, data points are few due tothe detection limit for sulde, but the plot, where thelog(IAPFeS) falls into a rather narrow range, indicates thatFeS precipitation is controlling the observed concentrationof sulde. This implies that there may be intervals whereFe2+ is controlled by equilibrium with both Fe-suldeand Fe-carbonate while at the same time carbonate iscontrolled by equilibrium with both Ca-carbonate andFe-carbonate, leading to an interlocked system where, e.g.the production of CO2 from methanogenesis may aectthe observed sulde concentration.

4.2. Exchange

Comparing the Na and Cl proles (Fig. 2) indicates thatNa is slightly delayed in terms of the vertical transport,indicating ion exchange processes aect the cations. Thisalso appears to aect the Ca, Mg, and K, resulting in whatappears to be a chromatographic sequence where Mg is dis-placing Ca which again is displacing K from the exchanger.The Ca is also aected by the dissolution of calcite, and thecalcite equilibrium in the lower part implies that if ionexchange releases Ca it may lead to precipitation of calcitepotentially lowering the pH. The ammonium (Fig. 5) doesnot seem to be aected by the ion exchange processes, but

1956 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966is probably, to a higher degree, controlled by the releasefrom organic matter being oxidized, implying that it maybe a cumulative indicator for the oxidation of organic mat-ter in the system. As mentioned dTIC-Ca (Fig. 5) is slightlynegative in three intervals also indicating that ion exchangeprocesses are aecting the Ca concentration.

4.3. Rates and sediment composition

There is a good coherence between where measured sul-fate reduction rates are high and where the sulde mineralsare found in the sediment. The AVS minerals are only

Fig. 6. Directly measured radiotracer rates of sulfate reduction, methanoxidation. The AVS part of the sulfate reduction rate is not shown directlrates for dierent redox processes derived from the total organic carbon aprocesses to enable the calculation, are shown as lines as indicated.

Fig. 7. Sediment content of Fe extractable by 6 N HCl acid, suldes CRsuldesmostly monosuldes) and organic C as TSOC (total sedimentaNADSOC is the dierence between the two. The plot also shows SIC (sefound in the interval from 5 to 6 mbs where the rates,and especially the AVS related rates are high showing thatthe kinetics of the transformation from AVS to CRS aremore sluggish than the sulfate reduction. In the lower partof the prole the suldes are almost exclusively found asCRS. The rates of sulfate reduction are extremely smallhere which must imply that sulfate input has been higherso that sulfate reduction took place in a larger volume ofthe sediment. This indicates that the system is in fact onlyin a quasi-steady-state. The total amount of sulde accumu-lated in the 56 mbs interval is around 7 mmol/kg of sedi-ment, which using a wet bulk density of 2 and a porosity

ogenesis from acetate oxidation and CO2 reduction, and methaney, but is the dierence between the total rate and the CRS rate. Theddition entered in the model, assuming at most two concommitant

S (chromium reducable suldemostly pyrite), AVS (acid volatilery organic carbon) and ADSOC (acid desorbable organic carbon).dimentary inorganic carbon).

alcite,ven in

Geochemistry at the sulfate reductionmethanogenesis transition zone 1957of 0.33 amounts to 42 mmol per liter of porewater. Giventhe rate of around 1 mmol/L/year observed in this intervalimplies that the suldes there could have formed in just 42years. This indicates that the zone where high rates of sul-fate reduction occurs has moved upwards within the sedi-ment. What controls this movement is not known, but itcould be that the most bioavailable Fe-oxides have beenused allowing sulfate reduction to occur at a higher levelin the aquifer. It is probably more complicated than that,because the amount of organic matter in the interval is alsorather low, around 30 mmol/kg or around 180 mmol/L ofporewater, implying that if the high rate was maintainedat this level the organic matter would have disappeared in90 years. With an age of the sediment around 1000 years,it seems there is too little organic matter and too little sul-de to match the measured rates. This implies that zones ofhigh rates are temporary. It also indicates that the decreasein DOC seen over depth reects transport of organic matterfrom the surface into the aquifer. Considering the lowamount of Fe-oxide in the system compared to the signi-cant release of FeII to solution 34 mbs (Fig. 3) could indi-cate that also Fe-oxides are supplied from the soil abovepresumably as colloids. Another apparently odd thing relat-ed to the sulfate reduction is that the measured sulde con-centration is highest, below the level where the rate ofsulfate reduction peaks. However since the sulde concen-trations according to Fig. 8 appears to be controlled by

Fig. 8. Mineral equilibria. Log plots of the activities of the ions in cin a large part of the system. The SI values (saturation indexes), gian FeS phase the high measured sulde concentrationsare a function of the low Fe2+ concentrations rather thanthe sulde production. This must imply that FeS can dis-solve, perhaps explaining the low amount accumulated,compared to the rate of sulfate reduction.

Acetate oxidation rates are relatively high in the lowerpart of the prole where there is no sulfate reduction, indi-cating another electron acceptor process of which the mostprobable would be Fe-oxide reduction, which is at least notcontradicted by the increase over depth in the amount of6 M HCl extractable Fe (Fig. 7). The acetate oxidation rateis approximately the same as the rate of methanogenesis byCO2 reduction. This could indicate that the acetate isoxidized to CO2 and H2 but as discussed in the sectionbelow this is probably not the case.

The high measured rates of sulfate reduction coincidewith the increase in Ca from 56 mbs. This was alsoseen in the similar Rm system (Jakobsen and Postma,1999) indicating a relation. Several explanations are pos-sible; the dissolving calcite itself contains reactive organ-ic matter which is released during dissolution, theincrease in Ca causes changes in the availability of theorganic carbon, or the increased ionic strength as suchchanges it.

4.4. Bioenergetics

The Gibbs energy for a range of probable microbiolog-ically mediated processes have been calculated using thethermodynamic data shown in Table 1. The values forthe central sampling site where H2 was measured are plot-ted in Fig. 9. The values for sulfate reduction are some ofthe rst values reported from a groundwater system basedon actual measured values for sulde, which was alwaysbelow the detection limit in the Rm aquifer (Jakobsenand Postma, 1999; Hansen et al., 2001). The values areclose to constant around 4.5 kJ/mol e from 4.56 mbsand decrease to 6 kJ/mol e at 6.8 mbs, and very similarto the values reported from larger Middendorf aquifer byPark et al. (2006). The calculated Gibbs energy is, howev-er, around twice as low as the 2.4 reported by Hoehleret al. (2001), which was argued to be adequate for ATPsynthesis. The lower Gibbs energies, and the decrease seenin the lower part of the prole, could be an indication of

siderite and FeS, indicating equilibrium control for calcite and FeSthe siderite and FeS plots are the values imposed in the model.a higher energy requirement when sulfate concentrationsare low. The value for methanogenesis by CO2 reductionby H2 is around 3 kJ/mol of e, in the interval wheremethanogenesis was detected by the direct rate measure-ments, similar to the values found by Park et al. (2006),but again about twice as low as the value of 1.3 kJ/mol of e from Hoehler et al. (2001). Again it could berelated to a dierence in, e.g. the PCO2 , but it could alsobe related to a dierence in the measurement method,since the H2 data from Hoehler et al. (2001) and the evenhigher values of 0.9 from Schulz and Conrad (1997) arefrom incubations of the sediment. In an overall sense theobserved values are similar to the values found for theRm setting (Hansen et al., 2001). Here in the Asserbosystem, 2.56 mbs, there is a large interval where the ener-gy available for CO2 reduction is very close to zero, whichtogether with the low, but measurable, concentration of

Fig. 9. The Gibbs energy for a range of possible microbialprocesses in the aquifer system calculated from the measuredconcentrations of reactants and products using the thermodynamicdata in Table 1.

Table 1Thermodynamic data used for calculations of in situ energy yields, andPHAST simulation obtained by adjusting the K-value for the reaction (s

Reaction aDGora(kJ/mol)

a

(

0f 2H+ + 2eMH2 17.61 Fe3+ + 0.5H2M Fe

2+ + H+ 83.12 4H2 SO42 H $ HS 4H2O 262.4 3g 4H2 HCO3 H $ CH4 3H2O 229.4 4 CH3COOH 4H2O$ 4H2 2HCO3 2H 241.85 HCOOHH2O$ H2 HCO3 H 40.26h CH3COOH SO42 $ HS 2HCO3 H 20.67i CH3COOHH2O$ CH4 HCO3 H 12.48h 4HCOOH SO42 $ HS 4HCO3 3H 101.6 9i 4HCOOHH2O$ CH4 3HCO3 3H 68.6 a All reactions calculated from values in Stumm and Morgan (1996), excb Gibbs energy of reaction at 8 C and standard conditions.c The shifted value of the Gibbs energy of reaction entered into the md The temperature corrected shifted value of the Gibbs energy of reacte The energy that has to be available from the reaction for the microof This reaction canbe seen as the reaction transferring excess electron activg Methane oxidation would be the same with the two sides of the equah Reactions 6 and 8 and the thermodynamic values are derived from thi Reactions 7 and 9 and the thermodynamic values are derived from th

1958 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966methane, 15 lM in this interval, points to the presence ofmicro-niches in which methane is produced. Darling andGoody (2006) also found methane in low concentrationsin a range of not highly reduced aquifers, indicating thatthis is a general feature. The methane leaving the micro-niches will only be oxidized as long as there is energyavailable for the microorganisms. This is not the casewhen Gibbs energies in the sampled water are close tozero as neither methane production nor oxidation can oc-

the energy yields at the temperature T = 281.15 (8 C) used in theee Table 2)

DHorkJ/mol)

bDGoTr(kJ/mol)

cDGoshr(kJ/mol)

dDGoT shr(kJ/mol)

eEnergy yield(kJ/mol)

4.18 16.3 17.6 16.3 038.5 80.5 70.6 68.7 11.8235.0 260.8 220.7 221.5 4.9237.9 229.8 206.8 208.6 2.7228.4 241.0 234.0 233.7 0.9115.1 38.8 46.8 44.9 3.1

6.6 19.8 13.3 12.2 5.819.5 11.2 27.2 25.1 3.61174.6 105.6 33.5 41.9 8.0177.5 74.6 19.6 29 5.8ept the value for acetic acid taken from Atkins and de Paula (2002).

odel converted to a logK value by dividing by 5.708.ion calculated using DHor .rganisms at the model temperature of 8 C.ity introducedbyadding theorganic carbon todissolved aqueousH2.tion swapped.e sum of reactions 2 + 4 and 2 + 4*(5).e sum of reactions 3 + 4 and 3 + 4*(5).cur. However, a principal dierence is that at this sitemethanogenesis from CO2 reduction by H2 does becomethermodynamically feasible below 6 mbs, where it is alsoobserved in the directly measured rates (Fig. 6), so thatstagnant zones or micro-niches with higher H2 contentneed not be implied here. It is also worth noting that sul-fate reduction via acetate does not seem to be feasiblewith a Gibbs energy above 1 kJ/mol e, and with only15 lM of methane above 6 mbs, where it would be ther-modynamically feasible to produce it by, e.g. acetate fer-mentation, it appears that H2 as substrate is extremelyimportant in this system. This is also indicated by theGibbs energy of H2 production from acetate (Fig. 9)which is feasible above 6 mbs where the H2 concentrationis low (Fig. 5). Still acetate could be important as a sub-strate for Fe-oxide reduction, but due to the diculties inassigning a value for the Gibbs energy of formation forthe Fe-oxide being reduced no attempt has been madein calculating the free energy for Fe-reduction reactions.However, the fact that Fe2+ is present over the entiredepth interval (Fig. 3) indicates that Fe-oxide reductionis not limited to the upper part. This is supported bythe acetate oxidation rates measured below 6 mbs of0.050.3 mM/yr (Fig. 6), where the transformation of ace-tate to H2 and CO2 is not feasible according to the Gibbsenergies in Fig. 9, due to the high H2 concentration.

4.5. Modeling

4.5.1. An extended partial equilibrium approach

The approach taken is a two-step PEA model using thetypology of Brun and Engesgaard (2002). It is an extensionof the partial equilibrium approach used by Jakobsen andPostma (1999). In order to better represent the actual sys-tem, and be able to model the observed H2 values the ener-gy needed for the microorganisms to carry out themicrobiologically mediated processes needs to be includedin the thermodynamic description of the system. This corre-sponds closely to assigning an observed saturation index,SI, to a mineral in a purely geochemical model. Just asthe SI of a mineral showing a constant value above 0.0may reect that extra energy is needed for precipitation totake place at a rate corresponding to the production of a

does, namely that if a saturated solution enters a cell withthe mineral, the mineral will dissolve until the solutionreaches the specied supersaturation, which should nothappen. However, as long as the system is maintained ina supersaturated state, which for the redox processes implythat organic matter is always being added to the system,this is not a problem.

4.5.2. PHAST, Model setup

4.5.2.1. Flow model parameters. The ow model domain hasa no-ow boundary corresponding to the groundwater di-vide, a bottom no-ow boundary corresponding to the bot-tom of the aquifer, a constant head boundary at thedownstream end of the model and a ux boundary ontop corresponding to the inltration set to 150 mm/yr. Auniform hydraulic conductivity of 8.5 105 in the hori-

C da

192:0

valu

262:3

for a

19:8 k

H2. To the

Geochemistry at the sulfate reductionmethanogenesis transition zone 1959reactants, the negative Gibbs energy for the microbiologi-cally mediated reactions reect that energy is needed, in thiscase in order to sustain the life functions of the bacteria at arate controlled by the rate of electron donor supply. In or-der to facilitate comparison with observed values, the redoxreactions in the PHREEQC database were rewritten so thatthe reductant was H2 instead of electrons (Table 2). In away this is along the lines of Hoehler (1998), who proposesthat the observed H2 level due to the hydrogenase enzymebeing present in many microorganisms, may actually reectthe internal redox state of many dierent types of anaerobicmicroorganisms, regardless of the main electron donorused. The logK value for the new reaction is calculatedfrom the DG of the rewritten reaction and then in orderto implement the energy surplus the logK value in the data-base used by PHREEQC is shifted to lower or higher val-ues, depending on the direction of the reaction in thedatabase. An example, using sulfate reduction, is shownin Table 2. The same approach was used to attempt themodeling of the observed acetate and formate values. Thiswas done by adding two reactions to the database whereC(0) as in acetate and C(2) as in formate, are transformedinto H2 and HCO3

and assigning logKs for the two reac-tions that are shifted from the thermodynamic equilibriumvalues, again to allow an energy gain for the microorgan-isms carrying out these processes. The shifting of the logKvalues for the redox reactions holds the same general prob-lem as the setting of a xed supersaturation of a mineral

Table 2An example of how redox reactions were rewritten in the PHREEQ

The reaction for sulfate reduction in the PHREEQC database is

SO42 9H 8e $ HS 4H2O with : logK 33:65=DGor

To make the value for the available energy comparable to observed

4H2 SO42 H $ HS 4H2O with : log K 45:95=DGor The logK value is then shifted (similar to imposing supersaturationavailable for the microorganisms:

4H2 SO42 H $ HS 4H2O shifted logK : 38:5=DGor 2Implying that the left side components need to be higher, e.g. more5.3 kJ/mol of electrons at 25 C. As the enthalpy of reaction is non-zerzontal direction was the result of tting the observedhydraulic gradient (Fig. 1), in the vertical direction a small-er value of 1.3 105 was used assuming that the horizontalbedding leads to lower values. The horizontal conductivityis slightly lower than the value of 1.3 104 m/s determinedby sieve analysis, but considering the uncertainties it comesclose. As shown in Fig. 1 a section of the upper ux bound-ary has a high salinity solution associated with it. Upstreamand downstream of this section, two dierent solutions withlower salinities inltrate. The inltrating solutions used arelisted in Table 4. The variation in the inltrating solutioncompositions, represent a horizontal 1D variation neededto model the features observed in the sampled proles.The width and the position of the high salinity inltrationwas chosen to match the peak in chloride at site 2(Fig. 2). The grid cells in the 750 m long and 10 m highmodel domain were 5 m long and 0.2 m high, the time-stepused was 0.02 yr. The weighting parameters used byPHAST for the space and time dierencing were set to0.32 and 0.60, respectively. This set of values is a compro-mise between calculation time, numerical oscillations andnumerical dispersion. Numerical dispersion is particularlydicult to avoid when the ow direction is not aligned withthe grid. The free water table and the recharge boundaryalong the top makes it inevitable. The modeled owdirections and velocities are shown as velocity vectors inFig. 1. Because of the numerical dispersion, no additionaldispersivity was entered. If the actual zone of high salinity

tabase so that a minimum energy yield of the reaction is available

7 kJ=mol

es based on H2 it is rewritten into

kJ=mol

mineral) so that the reaction only occurs when there is energy

J=mol

he available energy is 262.3 219.8 = 42.5 kJ/mol reaction oractual energy available will change with the temperature (see Table 1)

inltration does have sharp boundaries, the modeled Clmatches the measured Cl (Fig. 2) with the amount ofnumerical dispersion present in the used model setup.

4.5.2.2. Geochemical model parameters. The model setup isin terms of sediment geochemistry essentially a 1D system.This 1D system is very similar to the 1D PHREEQCapplication by Jakobsen and Postma (1999) for the Rmaquifer, but in this 2D PHAST model of the Asserbo site,horizontal zones, essentially layers, are associated withan initial solution given in Table 4, a number of solidphases and a rate of organic matter addition summarizedin Table 3.

For a given solid phase the imposed SI and initialamount of mineral, calcite, siderite, FeS etc., is entered.In this model, the layers have all been assigned a xed cal-cite saturation simply matching the calculated SI in theaquifer, in a way a geochemical boundary condition. Itwas chosen not to do a more elaborate matching of the re-sult of the calcite dissolution using a kinetic expression asthere are no data available on particle distribution, surfacearea etc. Because Mn peaks appears to coincide with thedissolution of calcite, and the increase in Mg also appearsrelated, the calcite used in the model were dissolution isdominant is not pure CaCO3 but Ca0.849Mg0.15Mn0.001CO3.Substitution of Ca by Mn, though much more substantial,in the form of Ca-rhodochrosites are known from the Baltic

Table 3Organic matter oxidation rates assumed for the model simulation

Depth(m)

C(0)rate mM/yr

SI Mg,Mn-calcite/calcite

SIFe(OH)3

1960 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 194919660.01.8 0.015 1.51.82.1 0.003 1.52.12.4 0.003 4.80 1.52.42.7 0.006 4.40 1.52.73.0 0.009 4.20 1.53.03.3 0.024 3.20 1.53.33.6 0.024 2.80 1.53.63.9 0.024 1.80 1.63.94.2 0.036 1.30 1.94.24.5 0.036 0.92 2.34.54.8 0.036 0.89 2.34.85.1 0.036 0.74 2.35.15.4 0.042 0.56 2.35.45.7 0.66 0.30 2.45.76.0 0.90 0.22 2.86.06.3 0.30 0.20 2.86.36.6 0.045 0.12 2.86.66.9 0.045 0.08 2.86.97.2 0.09 0.00 2.87.27.5 0.015 0.00 2.87.57.8 0.015 0.00 2.87.810.0 0.024 0.00 2.9Saturation indexes (SI) imposed on the system for Mg,Mn-calcite(down to 6.6 mbs), calcite below 6.6 mbs, and for Ferrihydrite,obtained by tting. The initial amount for these two was 10.0mol/L. The SI for amorphous FeS was set to 0.0 down to 6.9 mbsand 1.0 below, and for siderite 0.75 was used. Both values wereadjusted based on the plots in Fig. 8. the initial amount was set to0.0 for both.Sea (e.g. Jakobsen and Postma, 1989), and marine calcitegenerally contains high amounts of Mg (Mucci, 1987).The substitution inuences the solubility of the calcite. Astoichiometric solubility constant, with a logK value of8.38 was used for the MgMnCalcite (Busenberg andPlummer, 1989), neglecting the minor amount of Mn. Atdepth where PHREEQC calculations show that the pore-water is close to equilibrium with calcite, pure calcite wasused as the CaCO3 phase as this is more realistic in caseof reprecipitation. The specied SI of amorphous ferrihy-drite was used to adjust the stability of the Fe-oxide, andwith that, the tendency for FeIII to go in to solution asFe3+ and further be reduced by H2 according to Eq. (2)in Table 1. The Fe2+ concentration is thereby constrainedindirectly. The coupling implies that the specic value ofthe SI is related to the value of 25 kJ/mol chosen for theenergy available for Fe-reductiona value for which thereis currently no observations. The t was obtained bydecreasing the SI over depth. This could reect that themore unstable Fe-oxides are no longer present at depth,either due to previous reduction or due to a catalytic trans-formation as suggested by Pedersen et al. (2005). Theseuncertainties related to modeling the Fe-oxides reect theunresolved issues regarding Fe-oxides in natural systems.

As it was discussed for Fig. 8, it seems that siderite iscontrolling the Fe2+ concentration when an SI 0.75 forsiderite is reached. Based on Fig. 8, an SI of 0.0 for theFeS precipitate dened in PHREEQC by the reaction:FeS + H+ = Fe2+ + HS, has been imposed (using valuesfrom Davison (1991); FeS: logK = 2.95/Fe(HS)2:logK = 6.45), again with no initial FeS. It was chosen notto model the precipitation of FeS by a kinetic expressions,as the observations indicate that precipitation takes place ata constant supersaturation, and data on what might controlthe kinetics in this system are not available. Below the levelwhere sulde was detectable an SI = 1.0 was imposedimplying that the suldes eventually change to more stablephases.

The redox reactions in the model are driven by the inputof organic carbon, in the form of CH2O, to each cell in thesystem at a specied rate (Table 3). The organic carbon isoxidized to inorganic carbon and the increased electronactivity arising from this enters into the system of redoxequilibrium reactions (Table 1), which together with the im-posed mineral equilibria determine a new stable solutionchemistry for the given cell. The organic matter added toeach layer is put into the cells using the kinetic feature inPHREEQC, however, in this model the rate is constantover time, for a given layer. Several of the processes couldhave been tted using kinetics rather than using mineralequilibria, but the obtained parameters could not be com-pared to actual data on mineralogy, surface area etc. asthese are not available. Kinetic descriptions could poten-tially still be useful for long term modeling of the system,but not for a steady-state model. In the chosen setup, therate of organic matter oxidation together with the stabilityof the Fe-oxide are the main tting parameters. The ttedenergy yields (Table 1) for sulfate reduction and methano-genesis, and the saturation index for amorphous FeS wereall chosen very close to the observations. To produce the

observed increase in the NH4 concentration the organic

matter added needed to be given a C:N ratio of 10((CH2O)10NH3). This is above the Redeld ratio for marinealgae of 6.6 expected due to the marine origin of the sedi-ment. Well cultured soils have a C:N ratio of 10, indicatinga soil origin for the organic matter. However, the observedratio reects the organic matter being consumed, the bulkorganic matter may be quite dierent. The rate of additionof organic matter given in Table 3 is the total rate, in Fig. 6it is split up into the dierent redox processes calculatedassuming either Fe-oxide reduction with sulfate reductionor methanogenesis and plotted with the most relevantdirectly measured rates. Assuming a 1D vertical distribu-tion is a simplifying assumption implying that within themodel area the formation can be viewed as consisting ofhomogeneous layers in terms of the controlling geochemicalparameters. The tting of the model was done by compar-ing the model output to the central prole, site 2.

As the distribution of cations show that ion exchangereactions are occurring, an exchanger with a capacity of30 mmol(+)/L was specied for the entire system. Thiswas initialized by specifying the distribution of cations onthe exchanger to be in equilibrium with a solution basedon the standard exchange constants in the PHREEQCdatabase. The solution used was a constructed solution(see Table 4) made by mixing seawater with a water compo-sition from the upper part of the system and letting thismixture come into equilibrium with calcite. The solution

The model has been run for 100 years, the time it takesto almost exchange all of the water in the model with theproposed inltrating solutions, implying that what is beingmodeled is an assumed quasi-steady state situation. Thisseems to be a reasonable approximation for the redox pro-cesses where the sediment characteristics of a given layerpresumably change slowly, however, for the ion-exchangeprocesses this is a simplication.

4.5.3. Model results and discussion

Model results for individual parameters are shown in theFigs. 26 with the measured data, and will be discussed be-low. To give an impression of the 2D patterns Fig. 10adshows the simulation after 100 years of, a) a 2D cross-sec-tion of chloride, as a tracer of the ow pattern, b) sulfate, asan example of a solute being removed, c) a plot of the dis-solved FeII an example of a solute being produced and d)the amount of precipitated Fe-sulde after 100 years, anexample of a solid phase forming. Fig. 10a and b shows thatthere is some numerical oscillation occurring at the bound-ary between the inltration areas of the high and lowsalinity solutions, these oscillations are, however, smearedout upstream of the sampled area due to the numerical dis-persion. The sulfate distribution shows the relative com-plexity that arises when a very simple 1D variation ininltrating water chemistry, resulting in the pattern shownfor chloride in Fig. 10a is combined with a 1D distributionof sediment characteristics (organic matter oxidation rates

value

c Sodowinit

4.90.00.20.00.00.00.00.0f0.00.00.20.10.0

rom s

ed fo

the, 8.3om pe(OH

Geochemistry at the sulfate reductionmethanogenesis transition zone 1961composition could be seen as the result of inundating sea-water mixing with water in the system and equilibratingwith the calcite in the system.

Table 4Inltrating solutions and exchanger solution used in the model, all

Parameter(mM) (8 C)

a Solution inltrating inupstream end of model(medium TDS)

b Solution inltratingin central part ofmodel (high TDS)

pH 4.8 4.8Alkalinity 0.04 0.02Na 1.2 1.952K 0.16 0.13Mg 0.248 0.30Ca 0.186 0.756Mn 0.002 0.002FeII 0.04 0.04FeIII f0.00055 f0.00056O(0) 0.015 0.015Cl 1.35 2.00Sulfate 0.443 1.00Ammonium 0.002 0.002

C(-4) Br (tracer) 0.001

a Solution from site 2, 5.9 mbs modied to resemble the solution fand pH, and diluted to obtain an adequate ionic strength.b Solution from site 2, 5.9 mbs modied to resemble the solution us

increasing ionic strength.c Solution from site 1, 2.5 mbs, also used as the initial solution ind Initial solution below 5.1 mbs, taken as the solution from site 2e Dilute seawater with extra Ca, alkalinity and sulfate, perhaps frf Calculated by PHREEQC from equilibrium with amorphous Fg Calculated by PHREEQC from charge balance.and mineralogy). The ow pattern seen from Fig. 10a ispresumably a reasonable representation of the Asserboaquifer, but the complexity issue is worth keeping in mind

s are in mmol/L

lution inltratingnstream and upperial solution (low TDS)

d Initial solutionfor the lower partof the aquifer

e Solution usedfor initializing theexchanger

7.48 8.32 4.04 0.5757 1.02 2.0853 0.18 0.0414 0.39 0.1137 1.70 g0.48702 0.002 0.0024 0.04 0.040041 f0.00000215 0.009454 1.1 2.2035 0.008 0.302 0.261 0.001

0.337

ite 1 2.5 mbs used for the downstream part in terms of its oxic state

r the upstream part in terms of its oxic state and further modied by

model from 0 to 5.1 mbs.mbs.yrite oxidation and calcite dissolution.)3.

1962 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966when other, very probably, less simple ow systems with aless simple distribution of sediment characteristics are sam-pledgenerally at much lower sample resolution. Fig. 10d,shows, as expected, that most of the FeS precipitation takesplace where the high sulfate water intersects with the layerswith high rates of organic matter degradation.

Columns (1D) of modeled values for the three observa-tion points are plotted as lines with the observed data inFigs. 26. As described above, the tting of the modelwas done by comparing the model output to the centralprole, site 2. This means that the dierences in the chloridet for the 3 boreholes indicate either variations in the salin-ity of the high salinity water over time or space, or probablevariations in the actual ow paths in the system in responseto inltration variations and heterogeneities in the geology.The dierences in the chloride t are rather small, indicat-ing a relatively homogeneous geology on this scale, wherethe sampling sites are only 10 m apart. The small dierenceswill also show up in other parameters controlled by the in-put solution: sulfate Na, K, Mg, while other parameters,e.g. pH, Ca, FeII, alkalinity and sulde are mainly con-trolled by the imposed mineral equilibria and the coupling

Fig. 10. Model output concentrations in mM for (a) chloride, (b) sulfatreached a quasi-steady-state. Note that the ow in this gure is from left tand depth intervals of the sampling sites as indicated on the FeS plot (dto oxidation of the organic matter added to the system. Thisalso implies that the numerical dispersion aects theseparameters to a lesser degree.

The model t for the cations (Fig. 2) is not perfect in anyway, partly because the system is being modeled as if in aquasi-steady-state, while any eects of ion exchange pro-cesses imply that the system is not in a steady state. To rea-sonably model the cations, information on the actual timingof events aecting the exchanger composition, and thewater composition involved would be required, The peakscould reect changes in inltration chemistry on a timescale of years.

Comparing the shape of the sulfate and the chloridepeak shows that though the peak in the sulfate concentra-tion is maintained (due to sulfate reduction rates beinglow compared to the ux of sulfate into the system) thepresence of a layer with highly reactive organic matter cre-ates the illusion that sulfate is moving slower than chloride,while in fact it is the tip of the peak which is being removedby sulfate reduction, the good t obtained for the sulfate isdirectly related to the rate of organic matter oxidation inthe model.

e, (c) FeIIaq and (d) solid FeS after 100 years when the model haso right. The dashed lined boxes on each plot indicates the positions).

Geochemistry at the sulfate reductionmethanogenesis transition zone 1963The obtained t for the other redox related componentsis also pretty close. This is to some extent a result of adjust-ing the saturation indexes for calcite and FeS, to valuesclose to the observed values. However if the tted rate oforganic matter oxidation was slightly dierent then, be-cause of the close coupling with the mineral phases throughthe water chemistry implied in the partial equilibrium ap-proach, the t would be worse. The good t in all three pro-les indicates that, since all model sediment parametersonly vary in the vertical direction, the aquifer material, interms of the reactivity of the sedimentary organic matterand the mineral phases controlling the water chemistry,must be rather homogeneous in the horizontal plane, atleast within the scale of sampling. No attempt has beenmade to test the uniqueness of the calibrated set of param-eters, but tting the observations made it quite clear thatminor changes in the parameters have very large eectson the simulated output. The ammonium proles arematched using a single C:N ratio, but this leads to anunderestimation of the rate of organic matter oxidation be-low 6.5 mbs indicating that the C:N ratio of the organicmatter being oxidized in the lower part is higher, possiblyreecting a stripping of more N-rich parts over time.

In order to maintain a reasonable pH while the Fe(II)concentration builds up, it turned out to be necessary tohave some buering of the pH of the system. In this caseit was simply obtained by adjusting the TIC, by loweringthe pH in the inltration water, while maintaining the samelow alkalinity. The resulting TIC in the modeled water at 2mbs is around 1.3 mM, compared to the value of the upper-most groundwater, derived from measured alkalinity andpH, ranging from 1.8 mM at site 1 to only 0.3 at sites 2and 3. In the real system some of the pH buering couldbe related to Ca2+H+ exchange on the carboxylic acidgroups in the DOC which as mentioned (Fig. 4), shows anotable decrease at the top of the prole coinciding withthe increase in Ca. It could also be due to buering by pro-tons being released from surfaces in the sediment, as imple-mented in the model for the Cape Cod system (Parkhurstet al., 2003). However there are no data available to supportone or the other hypothesis, and considering that site 1 hada high TIC increasing the TIC seems reasonable.

Using SI = 0.75 for siderite leads to a model pH that istoo high in the lower part, resulting in Ca2+, becoming toolow, in the lower part. The simultaneous equilibrium withsiderite, calcite and Fe-sulde xes the Fe/Ca and theHCO3

=HS activity ratios, the end result being that themodeled Ca becomes too low and the modeled pH too high(Figs. 2 and 3). Setting the SI for siderite higher would givea better t of Ca and pH, but would imply that sideritewould not precipitate in the model.

In general, imposing all the equilibrium constraints list-ed to some extent locks the water composition in place andit could seem that the good ts shown for pH, alkalinityand Ca in Fig. 2 and the redox parameters in Fig. 3 aremerely a result of xing all these equilibria. However,CH4 and sulfate are not xed by equilibria and the sideritesaturation discussed above shows that xing these equilib-ria does not necessarily lead to closer ts. The tted rateof organic matter oxidation is close to the measured valuefor sulfate reduction (Fig. 6), and, since this is an indepen-dent measurement, the good t indicates that the descrip-tion of the system by a partial equilibrium approach is agood approximation for the upper part. In the lower partthe model rates are lower than the measured rates. Simplyincreasing the overall rate would increase the modeledammonium and the methane concentrations. The modeledammonium could be lowered by a higher C:N ratio, butto avoid too high methane concentrations requires thatthe extra added carbon is used to increase the amount ofFe-oxide being reduced, which would also t better withthe high measured rates of acetate oxidation (Fig. 6). How-ever to increase the amount of Fe-oxides being reducedwould require increasing the solubility of the Fe-oxides inthe lower part, which again due to the coupled equilibriawould lead to an even higher pH. Still, the close to constantmeasured Fe2+ concentration below 7 mbs, together withthe calculated supersaturation for siderite does presumablyconceal a concomitant reduction of Fe-oxides and metha-nogenesis as reported by Berner (1981), Jakobsen and Post-ma (1999) and Ferro and Middelburg (2003), and as theappreciable acetate oxidation rates in this zone (Fig. 6)together with the lack of available Gibbs energy for the pro-duction of H2 from acetate (Fig. 9), would indicate. It isinteresting to note that if siderite is precipitating whileFe-oxide reduction and methanogenesis proceed concomi-tantly at a ratio of 1:6 in terms of C(0) transformed, onemay have a net transformation of Fe-oxides to siderite, con-currently with methanogenesis, with no direct trace of theFe-oxide reduction in the porewater chemistryexceptfor a constant SI for siderite according to Eqs. (7)(9):

Fe-oxide reduction :

4FeOOH CH2O 3H2CO3 ! 4FeCO3 6H2O 7Methanogenesis :

6CH2O 3H2O! 3H2CO3 3CH4 8Combined reaction :

4FeOOH 7CH2O! 4FeCO3 3CH4 3H2O 9this scenario would be ideal from a microbiological systempoint of view as the energy available for the Fe-oxide reduc-tion would be maintained, and the energy available formethanogenesis would stay higher than without the precip-itation of the produced carbonate. This `deality could causethe system to approach the ideal ratio and stay in this stateas long as the kinetics of the Fe-oxide reduction could keepup. In the model, the ratio is in the interval 1:6.47.8 belowa depth of 8.4 mbs, where there is also siderite precipitationin the model taking place at a rate matching the Fe-oxidereduction, indicating a thermodynamic optimum.

Considering that the values for the threshold energyyields for sulfate reduction and methanogenesis entered inthe model were calculated from observed H2 values, onewould expect the nice t to the H2 data where these process-es occur (Fig. 4). So for these two reactions the appliedshifting of the equilibrium constants will yield a good pre-diction of the H2 concentrationsespecially if H2 measure-ments for calibrating the energy yields of the processes inthe given system are present. The model values for H2in the upper part dominated by Fe-oxide reduction are

1964 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966underestimated, but could not be increased to the measuredvalues using reasonable t parameters. If the thresholdenergy yield for Fe(III) reduction to occur is set higher,to increase the modeled H2 concentration, the Fe(III)-oxidewould have to be more soluble than ferrihydrite, in orderfor enough Fe3+ to be available for reduction. The explana-tion could be that stagnant zones play a role as discussed inSect. 4.4. The 1D dual porosity model used by Jakobsen(2002) could give a better t of the H2 concentrations thana single porosity model, but there is currently no dualporosity option in PHAST. Still, the fact that the Fe-oxidesare solids could also imply that the reactivity of the Fe-ox-ides is more limiting than the reactivity of the organic mat-ter as discussed in Larsen et al. (2006). Or, that a simplepartial equilibrium description is in fact too simple andneeds to be expanded with kinetic terms as suggested, moregenerally, by Jin and Bethke (2005), for reactions occurringat more oxic conditions than sulfate reduction, as shown tobe the case for the reduction of chlorinated solvents(Heimann and Jakobsen, 2006). Also several recent studiessuggest that Fe-oxide reduction of the solid phase oxide isfacilitated by either chelators assisting in the dissolutionof the Fe-oxides (Luu and Ramsay, 2003), electron shuttles(Nevin and Lovley, 2002), removing the need for directcontact, or nano-wires that transfer the electrons directly(Reguera et al., 2005).

For formate and acetate, the simplied model appliedhere, where acetate and formate are transformed via H2,is obviously not adequate (Fig. 5), and denitely an approx-imation as all of the redox processes in the system can,depending on the microbiology, in fact proceed with acetateas electron donor. A more comprehensive approach model-ing several electron pathways was not made, as this wouldinvolve tting numerous unknown parameters controllingthe entire microbial ecology. Another complication is thatfor example acetate may both form from, and transforminto H2 and CO2, and the simple shifting of the logK value,to increase the modeled acetate concentration, implies thatthe model may form acetate when there is actually no ener-gy available for this. As mentioned above this may work ina system where electron donor is constantly added every-where, but it is a simplication. The radiotracer measure-ments, however, indicate that at least for methanogenesisthe CO2 reduction by H2 is in fact the preferred pathway.Ideally this type of microbial reaction, that can go bothways should be implemented so that it only occurs with apositive energy yield in either direction as it was done formethanogenesis by Jakobsen (2002). Still the ability of themodel to reasonably predict the H2 level can be importantwhen evaluating the potential for reduction of pollutants,e.g. Cr(VI) reduction (Marsh and McInerney, 2001). So,though the model t is not perfect it does indicate that inmore general terms the release of inorganic carbon, Fe2+

and sulde leads to a state where the groundwater compo-sition is controlled by several mineral equilibria whichagain aect the ongoing redox processes forming a closelycoupled system. An example scenario shows the implica-tions of this: if the ux of sulfate into the system decreases,methanogenesis will occur higher up in the system and thenet production of carbonic acid from this will lead to therelease of FeII and Ca from siderite and calcite and FeIIand sulde from Fe-suldes. The higher H2 concentrationoccurring during methanogenesis and any decrease in pHthat is not neutralized by the mineral dissolution would in-crease the potential for Fe-oxide reduction, but on the otherhand the increase in Fe2+ will lower this (Eq. 2, Table 1). Ifthe net eect is an increase in Fe-oxide reduction this willlead to an increase in pH, again aecting the mineral disso-lution and eventually a new quasi-steady state is reached.

The model indicates that the long term development ofthe system will be a function of how the input of sulfate,the bioavailability of the Fe-oxides and the reactivity ofthe organic matter develop over time. Sulfate input coulddecrease due to, e.g. better pollution control, initiating thescenario described above. Sulfate input could also increasedue to more extreme weather conditions in the future, lead-ing to a higher dry deposition, which would force the zoneof methanogenesis deeper into the system again leading to acascade of eects towards a new steady-state.

5. CONCLUSIONS

5.1. A comparison with the Rm aquifer

The Asserbo aquifer is in many ways geochemically sim-ilar to the Rm aquifer, with some noteworthy dierences.(1) There is no pool of AVS below the sulfate reducing zoneindicating that there has been enough sulde available forthe conversion, which could be related to the higher mea-surable sulde concentrations in the Asserbo system com-pared to the non-measurable concentrations in the Rmaquifer. (2) The H2 level in the lower part of the system ishigh enough to sustain methanogenesis removing the needfor stagnant microniches, the reason for this is not obvious.(3) The Asserbo data on DOC clearly indicate an inux oforganic matter to the system from the soil above, presum-ably supplementing the sediment bound organic matter,which appears to be necessary for sustaining redox process-es in the system. This is probably also the case on Rm butdata on this were less clear.

5.2. Modelling

The partial equilibrium approach using shifted K-valuesfor the redox processes assuming H2 is the electron donorcomes a long way in describing the observed redox chemis-try including the H2 concentrations when sulfate reductionor methanogenesis are dominating. When going in to moredetail it is clear that in order to model the concentrations ofacetate and formate in general and the H2 concentrationwhen Fe-oxide reduction is dominating, a model with acloser description of the microbial ecology, the kinetics ofFe-oxide reduction, and perhaps able to handle dual poros-ity is needed. The relatively simple setup using just three dif-ferent inltration chemistries and a 1D variation insediment characteristics does produce most of the complexvariation and shows that in the very probable case whereinltration composition, geology and geochemistry are allless simple, quantitative interpretations of observationsbased on less densely spaced groundwater and sediment

Geochemistry at the sulfate reductionmethanogenesis transition zone 1965samples will be dicult to obtain. The overall conclusion isthat anaerobic aquifer systems may develop to a state wherethe groundwater chemistry is controlled by a set of inorgan-ic dissolution and precipitation reactions that are closelycoupled with terminal electron acceptor processes drivenby the oxidation of organic matter in the system. In sucha system, small changes in one input parameter may leadto considerable changes in the whole system.

ACKNOWLEDGMENTS

We thank the editor and two anonymous reviewers whose help-ful input has led to substantial improvement of the manuscript.

REFERENCES

Amend J. P., and Teske A. (2005) Expanding frontiers in deepsubsurface microbiology. Palaeogeogr. Palaeoclimat. Palaeo-ecol. 219(12), 131155.

Atkins P., and de Paula J. (2002) Atkins Physical Chemistry,seventh ed. Oxford University Press, New York.

Berner R. A. (1981) Authigenic mineral formation resulting fromorganic matter decomposition in modern sediments. Fortschr.Miner. 59, 117135.

Brun A., and Engesgaard P. (2002) Modelling of transport andbiogeochemical processes in pollution plumes: literature reviewand model development. J. Hydrol. 256, 211227.

Busenberg E., and Plummer L. N. (1989) Thermodynamics ofmagnesium calcite solid-solutions at 25 and 1 atm totalpressure. Geochim. Cosmochim. Acta 53, 11891208.

Caneld D. E., Thamdrup B., and Hansen J. W. (1993) Theanaerobic degradation of organic matter in Danish Coastalsediments: iron reduction, manganese reduction and sulfatereduction. Geochim. Cosmochim. Acta 57, 38673883.

Chapelle F. H., and McMahon P. B. (1991) Geochemistry ofdissolved inorganic carbon in a coastal plain aquifer. 1. Sulfatefrom conning beds as an oxidant in microbial CO2 production.J. Hydrol. 127, 85108.

Cline J. D. (1967) Spectrophotometric determination of hydrogensulde in natural waters. Limnol. Ocean. 14, 454458.

Crill P. M., and Martens C. S. (1986) Methane production frombicarbonate and acetate in an anoxic marine sediment.Geochim. Cosmochim. Acta 50, 20892097.

Curtis G. P. (2003) Comparison of approaches for simulatingreactive solute transport involving organic degradation reac-tions by multiple electron acceptors. Comp. Geosci. 29, 319329.

Darling W. G., and Goody D. C. (2006) The hydrogeochemistry ofmethane: evidence from English groundwaters. Chem. Geol.229, 293312.

Davison W. (1991) The solubility of iron suldes in synthetic andnatural-waters at ambient-temperature. Aquat. Sci. 53,309329.

Ferro I., and Middelburg J. J. (2003) Factors limiting iron oxidereduction in a tidal freshwater and a Mediterranean deep seasediment. In Ferro I. (2003) Cycling of iron and manganes infreshwater, estuarine and deep sea sediment, Ph.D. Disserta-tion, University of Groningen, Netherlands.

Hansen L. K. (1998) Biogeochemistry of methane in a shallowsandy aquifer. Ph.D. dissertation, Department of Geol. andGeotechn. Eng. Technical University of Denmark.

Hansen L. K., Jakobsen R., and Postma D. (2001) Methanogenesisin a shallow sandy aquifer, Romo, Denmark. Geochim.Cosmochim. Acta 65, 29252935.Heimann A. C., and Jakobsen R. (2006) Experimental evidence fora lack of thermodynamic control on hydrogen concentrationsduring anaerobic degradation of chlorinated ethenes. Environ.Sci. Technol. 40, 35013507.

Hoehler T. M. (1998). Thermodynamics and the role of hydrogen inanoxic sediments, Ph.D. dissertation. University of NorthCarolina, USA.

Hoehler T. M., Alperin M. J., Albert D. B., and Martens C. S.(2001) Apparent minimum free energy requirements formethanogenic Archaea and sulfate-reducing bacteria in ananoxic marine sediment. FEMS Microb. Ecol. 38, 3341.

Howarth R. W., and Merkel S. (1984) Pyrite formation andmeasurements of sulfate reduction in salt marsh sediments.Limnol. Oceanogr. 29, 598608.

Hunter K. S., Wang Y., and van Cappelen P. (1998) Kineticmodeling of microbially-driven redox chemistry of subsurfaceenvironments. J. Hydrol. 209(14), 5380.

Jakobsen R., and Postma D. (1989) Formation and solid-solutionbehavior of Ca-rhodochrosites in marine muds of the balticdeeps. Geochim. Cosmochim Acta 53, 26392648.

Jakobsen R., and Postma D. (1994) In situ rates of sulfatereduction in an aquifer (Rm, Denmark) and implicationsfor the reactivity of organic matter. Geology 23,11031106.

Jakobsen R., and Postma D. (1999) Redox zoning, rates of sulfatereduction and interactions with Fe-reduction and methanogen-esis in a shallow sandy aquifer, Rm, Denmark. Geochim.Cosmochim. Acta 63, 137151.

Jakobsen R. (2002) Including stagnant zones in order to modeloverlapping redox processes occurring in a shallow sandyaquifer. Geochim. Cosmochim. Acta 66(15A), A362, Suppl. 1Aug.

Jensen D. L., Boddum J. K., Tjell J. C., and Christensen T. H.(2002) The solubility of rhodochrosite (MnCO3) and siderite(FeCO3) in anaerobic aquatic environments. Appl. Geochem.17, 503511.

Jin Q., and Bethke C. (2005) Predicting the rate of microbialrespiration in geochemical environments, Geochim. Cosmo-chim. Acta 69, 11331143.

Kipp K. L., Jr. (1997) Guide to the revised heat and solutetransport simulator: HST3D. U.S. Geological Survey Water-Resources Investigations Report 97-4157, 149p.

Larsen O., Postma D., and Jakobsen R. (2006) The reactivity ofiron oxides towards reductive dissolution with ascorbic acid in ashallow sandy aquifer (Rm, Denmark). Geochim. Cosmo-chim. Acta 70, 427435.

Luu Y., and Ramsay J. A. (2003) Review: Microbial mechanismsof accessing insoluble Fe(III) as an energy source.W. J. Microb.Biotechnol. 19, 215225.

Marsh T. L., and McInerney M. J. (2001) Relationship of hydrogenbioavailability to chromium reduction in aquifer sediments.Appl. Environ. Microbiol. 2001, 15171521.

Nevin P. K., and Lovley D. R. (2002) Mechanisms for Fe(III) oxidereduction in sedimentary environments. Geomicrobiol. J. 19,141159.

Park J., Sanford R. A., and Bethke C. M. (2006) Geochemical andmicrobiological zonation of the Middendorf aquifer, SouthCarolina. Chem. Geol. 230, 88104.

Parkhurst D. L., Kipp K. L., Engesgaard P., Charlton S. C. (2005)PHASTa program for simulating ground-water ow, solutetransport, and multicomponent geochemical reactions. U.S.Geological Survey Techniques and Methods, 6, A8, 154p.

Parkhurst D. L., Stollenwerk K. G., Colman J. A. (2003) Reactive-transport simulation of phosphorus in the sewage plume at theMassachusetts Military Reservation, Cape Cod, Massachusetts.

U.S Geological Survey Water-Resources Investigations Report,03-4017, 33p.

Parkhurst D. L., Appelo C. A. J. (1999) Users guide to PHREEQC(Version 2), U.S Geological Survey Water-Resources Investiga-tions Report, 99-4259.

Pedersen H. D., Postma D., Jakobsen R., and Larsen O. (2005)Fast transformation of iron oxyhydroxides by the catalyticaction of aqueous Fe(II). Geochim. Cosmochim. Acta 69, 39673977.

Pedersen K. (2000) Exploration of deep intraterrestrial microbiallife: current perspectives. FEMS Microbiol. Lett. 185(1), 916.

Postma D. (1981) Formation of siderite and vivianite and the pore-water composition of a recent bog sediment in Denmark. Chem.Geol. 31, 225244.

Postma D., and Jakobsen R. (1996) Redox zonation: Equilibriumconstraints on the Fe(III)/SO4-reduction interface. Geochim.Cosmochim. Acta 60, 31693175.

Reguera G., McCarthy K. D., Mehta T., Nicoll J. S., TuominenM. T., and Lovley D. R. (2005) Extracellular electrontransfer via microbial nanowires. Nature 435, 10981101.

Schulz S., and Conrad R. (1997) Inuence on temperature onpathways to methane production in the permanently cold

profundal sediment of Lake Constance. FEMS Microb. Ecol.20, 114.

Starr R. C., and Ingleton R. A. (1992) A new method for collectingcore samples without a drilling rig: Ground Water Mon. Rev.Winter 1992, 9195.

Stookey L. L. (1970) Ferrozinea new spectrophotometric reagentfor iron. Anal. Chem. 42, 779781.

Thullner M., van Cappellen P., and Regnier P. (2005) Modeling theimpact of microbial activity on redox dynamics in porousmedia, Geochim. Cosmochim. Acta 69, 50055019.

Watson I. A., Oswald S. E., Banwart S. A., Crouch R. S., andThornton S. F. (2005) Modeling the dynamics of fermentationand respiratory processes in a groundwater plume of phenoliccontaminants interpreted from laboratory- to eld-scale. Envi-ron. Sci. Technol. 39, 88298839.

Whiticar M. J., Faber E., and Schoell M. (1986) Biogenic methaneformation in marine and freshwater environments: CO2 reduc-tion vs. acetate fermentationisotopic evidence. Geochim.Cosmochim. Acta 50, 693709.

Associate editor: Jack J. Middelburg

1966 R. Jakobsen, L. Cold / Geochimica et Cosmochimica Acta 71 (2007) 19491966

Geochemistry at the sulfate reduction-methanogenesis transition zone in an anoxic aquifer-A partial equilibrium interpretation using 2D reactive transport modelingIntroductionMethods and settingGroundwater sampling and analysisRadiotracer rate measurementsSediment parametersSite description

ResultsGroundwater chemistryMeasured ratesSediment geochemistry

DiscussionMineral equilibrium controlExchangeRates and sediment compositionBioenergeticsModelingAn extended partial equilibrium approachPHAST, Model setupFlow model parametersGeochemical model parameters

Model results and discussion

ConclusionsA comparison with the R oslash m oslash aquiferModelling

AcknowledgmentsReferences