Methanogenesis in the sediment of the acidic Lake Caviahue in Argentina

Journal of Volcanology and Geothermal Research 178 (2008) 197–204

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Journal of Volcanology and Geothermal Research j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / j vo l g e o r e s

Methanogenesis in the sediment of the acidic Lake Caviahue in Argentina Matthias Koschorreck a,⁎, Katrin Wendt-Potthoff a, Burkhard Scharf a, Hans H. Richnow b a b

Helmholtz Centre for Environmental Research – UFZ, Department of Lake Research, Brückstr. 3a, D-39114 Magdeburg, Magdeburg, Germany Helmholtz Centre for Environmental Research – UFZ, Department of Isotope Biogeochemistry, Permoserstr. 15, D-04318 Leipzig, Germany

a r t i c l e

i n f o

Article history: Accepted 2 June 2008 Available online 27 June 2008 Keywords: methanogenesis acidic lake biogeochemistry sediment carbon isotopes

a b s t r a c t The biogeochemistry of methane in the sediments of Lake Caviahue was examined by geochemical analysis, microbial activity assays and isotopic analysis. The pH in the water column was 2.6 and increased up to a pH of 6 in the deeper sediment pore waters. The carbon isotope composition of CH4 was between −65 and −70‰ which is indicative for the biological origin of the methane. The enrichment factor ε increased from −46‰ in the upper sediment column to more than −80 in the deeper sediment section suggesting a transition from acetoclastic methanogenesis to CO2 reduction with depth. In the most acidic surface layer of the sediment (pH b 4) methanogenesis is inhibited as suggested by a linear CH4 concentration profile, activity assays and MPN analysis. The CH4 activity assays and the CH4 profile indicate that methanogenesis in the sediment of Lake Caviahue was active below 40 cm depth. At that depth the pH was above 4 and sulfate reduction was sulfate limited. Methane was diffusing with a flux of 0.9 mmol m− 2 d− 1 to the sediment surface where it was probably oxidized. Methanogenesis contributed little to the sediments carbon budget and had no significant impact on lake water quality. The high biomass content of the sediment, which was probably caused by the last eruption of Copahue Volcano, supported high rates of sulfate reduction which probably raised the pH and created favorable conditions for methanogens in deeper sediment layers. © 2008 Elsevier B.V. All rights reserved.

1. Introduction An environment's pH is an important regulator of microbial processes. Usually at low pH microbial turnover rates are slower (Goodwin and Zeikus,1987). However, in some cases the inhibitory effect of acidic conditions is not so clear. In the case of microbial iron oxidation, the process is even stimulated by low pH. Methanogenesis is usually inhibited by low pH although methane production has been shown in acidic peat lands (Williams and Crawford, 1985; Goodwin and Zeikus, 1987; Dunfield et al., 1993; Hornibrooke et al., 2000; Horn et al., 2003; Bräuer et al., 2004). Since peat lands usually have a pH above 4, only limited information on methanogenesis is available for pH conditions below 4. Methanogenic activity has been shown in an acidic peat down to pH 3.8 (Kotsyurbenko et al., 2004). Some methanogens may be pH sensitive because of their restricted substrate spectrum. Acetate as a major carbon source for methanogenesis might not be available to the methanogens at low pH because non-dissociated acetate may be inhibitory to methanogenesis (Fukuzaki et al., 1990). Lower pH conditions are also suggested to reduce H2 microbial producing and consuming processes in anaerobic environments (Goodwin et al., 1988). Methanosarcina like bacteria have been enriched from a bioreactor operated at pH 4.2 (Florencio et al., 1993).

⁎ Corresponding author. UFZ, Helmholtz Centre for Environmental Research, Department of Lake Research, Brückstr. 3a, D-39114 Magdeburg, Germany. Tel.: +49 391 8109 405; fax: +49 391 8109 150. E-mail address: [email protected] (M. Koschorreck). 0377-0273/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2008.06.017

An isolate from a landfill was able to grow at pH 5 (Lapado and Barlaz, 1997). An isolate from a peat land grew at pH 5.3 but produced some methane down to pH 3.1 (Williams and Crawford, 1985). Acidotolerant (pH 4) hydrogenotrophic methanogenic consortia were enriched from a peat bog (Sizova et al., 2003). Molecular analysis of an acidic peat bog showed the presence of Methanomicrobiaceae and Methanosarcinaceae at pH 4.5 (Kotsyurbenko et al., 2004). To our best knowledge true acidophilic methanogens have not been isolated so far. Microbial sulfate reduction is also inhibited by low pH and acidophilic sulfate reducing bacteria have not been isolated so far. In acidic mining lakes sulfate reduction only occurs when the pH in the sediment is almost neutral (Meier et al., 2004). Recently, however, microbial sulfate reduction was observed in the sediments of the extremely acidic volcanic Lake Caviahue in Argentina (Koschorreck et al., 2003). This was explained by the availability of high amounts of diverse organic substrates and the absence of competing microbial processes. In the same sediment, elevated concentrations of methane were observed (Wendt-Potthoff and Koschorreck, 2002). This raised the question of whether methanogenesis takes place in the naturally acidic sediment. Sedimentary methane is commonly formed by microbial activity but in volcanic regions like Caviahue a geogenic origin of methane is also likely. The carbon isotope composition of methane allows characterization of sources and sinks of methane (Whiticar, 1999). The isotope fractionation between dissolved inorganic carbonate (DIC= CO2, HCO−3, CO2− 3 ) and CH4 illustrated by the enrichment factor (ε‰ = δ13Ccarbonate − δ13Cmethane) can be used to characterize

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Fig. 1. Vertical profiles of pH (a), dissolved gases CO2 (b) and CH4 (c) in different sediment cores from Lake Caviahue. The pH was measured in KCl extracts during the field work in 2003 (■ 2001, ● 2003, ▼ 2004).

methane formation as well as degradation processes. Low ε DIC–CH4 values (b20‰) are indicative for methane oxidation. Anaerobic methane oxidation depending on sulfate, iron or nitrate (Boetius et al., 2000; Grossman et al., 2002; Valentine, 2002; Islas-Lima et al., 2004) can lead to a significant isotope fractionation of methane as well. Therefore, aerobic and anaerobic methane oxidizing processes are difficult to characterize by isotope values alone. Enrichment factors between 10 and 50‰ may show acetogenic methane formation and high enrichment factors (N50‰) may point to CO2 reduction as the dominating process of methane formation (Whiticar et al., 1986; Botz et al., 1996; Waldron et al., 1998; Conrad et al., 1999; Whiticar 1999; Hornibrook et al., 2000; Grossman et al., 2002). For this work we use a combination of geochemical analysis, microbial activity assays and isotopic analysis in order to analyze methanogenesis under the extreme acidic conditions in the sediment of Lake Caviahue and to evaluate its role in the sedimentary carbon cycle of this extreme habitat. 2. Material and methods 2.1. Study site Lake Caviahue, Argentina, is a naturally acidified glacial lake located at 37° 53′ S; 71° 02′ W, at an altitude of 1600 masl. Morphometric characteristics are: maximum length = 5.6 km, maximum width = 2.5 km, max depth= 95 m (North basin), mean depth = 51.4 m, area = 9.22 km2, volume = 474,260 m3 (Rapacioli, 1985; Pedrozo et al., 2001). The study was conducted at the deepest part of the lake in the centre of the northern basin. The lake is fed by the acidic Agrio River (pH 1.8) originating at the Volcano Copahue (Pedrozo et al., 2001). The lake water had pH 2.6, and concentrations of iron and sulfate were 0.4 and 4.2 mmol L− 1, respectively. The last eruption of the volcano was in July 2000 (Varekamp et al., 2001). 2.2. Field work Samplings were carried out in February 2001 and March 2003 and 2004. Sediment cores were taken with a gravity corer (UWITEC, Austria) at the deepest part of the lake. This gravity corer allows taking 40 cm long cores with 60 mm in diameter. The cores were immediately sampled at the lake shore. H2S was measured by a polarographic microelectrode (UNISENSE, Aarhus, Denmark). The pH was measured by a conventional glass electrode (WTW, Weinheim, Germany) and sediment samples were filled in gas tight plastic bags and stored cool at

4 °C. In 2003 the pH values in sediment samples were determined after suspension in 1 M KCl and 1 h settling time, since direct insertion of a pH electrode within the compact and dry sediments of deeper layers was not possible. Samples for gas determination were filled in preweighed glass vials (14 ml), 2 ml H2O (pH 2) was added and the vessels were closed with a butyl rubber stopper (Ochs, Bovenden, Germany). In order to determine isotopic signatures of methane, 2 g of NaCl was filled into glass vials to form a saturated salt solution with the later added distilled water and sediment. This preparation stops microbial activity and improves the partitioning of methane into the gas phase (headspace) of the vial (Richnow, unpublished results). Vials were closed with thick butyl rubber stoppers and crimp-sealed, then the contents were mixed to dissolve NaCl and stored at room temperature until mass spectrometric analysis. During the survey in 2003, a sediment core was taken up to 6.4 m sediment depth using a sampling-platform and a piston corer (UWITEC, Mondsee, Austria, www.uwitec.at). Before the coring started, the exact water depth was determined using a gravity corer (UWITEC Mondsee, Austria). The resulting short core was also used for analyzing the sediment–water contact zone. A Livingston piston corer without liners and without core catcher equipped with thin-wall stainless steel tubes (60 mm× 2 m) was used thereafter. The piston corer was pushed into the sediment by a 20 kg weight. The steel tube penetrated into the sediment about 5 mm per each punch, in ash layers somewhat less. The sequence of the cores was 0–2 m, 2–4 m and 4–6 m. The sediment was extruded after coring by hydraulic pressure using a small water hydraulic hand pump. Unfortunately the piston coring technique did not sample the unconsolidated surface sediment at the water–sediment interface. The pH values were used to correlate the sediment layers of the surface sediments taken from the small cores with the longer piston cores used for sampling the deeper sediments. Extrapolating the pH profile of the small core (Fig. 1a) to greater depth, the pH of the uppermost layer of the long core (pH 4.2) is reached in 40 cm depth. Thus, the position of the top of the long piston core was set to 40 cm in order to obtain the real depth below the sediment surface. 2.3. Microprofiles During the survey in 2001, a sediment core (15 cm) was recovered and microprofiles of H2S, pH and O2 were measured with microsensors within 30 min after sampling in our field lab. H2S and pH were measured with microelectrodes (UNISENSE, Denmark), O2 was measured with

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Fig. 2. Calculation of selected vertical profiles of production using the software PROFILE. Black dots are measured concentrations and the dotted lines the fitted concentration profiles. The black lines show the vertical profiles of production of the particular compound modeled from the concentration profile.

microoptodes (PRESENS, Germany). The water overlying the sediment was mixed by carefully bubbling with air without disturbing the sediment surface. The core was kept in a bucket with lake water to keep the temperature as close as possible to in situ conditions. During the measurement of 3 microprofiles, however, the temperature in the water rose from 11 to 18 °C. The core was kept from light as the measurements were taken.

pared in deep microwell plates and incubated at 28 °C for 6 weeks. MPN were then calculated using the program of Klee (1993). The phosphate content of phospholipids was determined as an estimate of total viable microbial biomass in the sediments (Frostegard

2.4. Bacterial numbers Methanogenic bacteria were counted by three-tube Most Probable Number (MPN) analysis. Gas tight test tubes (15 ml) were filled with 4.5 ml of a medium containing acetate and H2 (in g L− 1): 0.2 KH2PO4, 0.25 NH4Cl, 1 NaCl, 0.4 MgCl2·6 H2O, 0.5 KCl and 0.15 CaCl2·2 H2O. The medium was supplemented with Na2S (0.12 g L− 1), Vitamin B12 (50 μg L− 1), 0.2 ml L− 1 selenite–tungstate-solution (Wendt-Potthoff and Koschorreck, 2002), 8 vitamin-solution (Pfennig,1978) (1 ml L− 1),1 ml L− 1 trace element solution SL 12B (Widdel and Pfennig, 1981), sodium acetate (0.82 g L− 1) and resazurin (0.5 μg L− 1). The medium was adjusted to pH 7.2 and gassed with 80% H2/20% CO2. We used 0.5 ml inoculum and 1:10 dilutions. Growth was counted positive when the CH4 concentration in the headspace was N100 ppmv after 13 weeks. Viable counts of aerobic iron oxidizing and thiosulfate-oxidizing bacteria and anaerobic neutrophilic and acidophilic iron reducing and sulfate reducing bacteria were determined using serial dilutions for Most Probable Number (MPN) analysis in established selective media (Wendt-Potthoff and Koschorreck, 2002). Dilution cultures were pre-

Fig. 3. Pore water concentration of CO2 (■), CH4 (●) (a) and CH4 production in batch assays (columns in b) in a long sediment core. The line (b) shows the pH measured in KCl extracts. The thin horizontal lines separate different core segments.

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Fig. 4. Isotope values of CO2 (a), CH4 (b) and enrichment factor εc (c). Enrichment factors were calculated according to Whiticar (1999) (ε = δmethane − δcarbonate). The thin horizontal lines separate different core segments.

et al., 1991). We applied the modifications suggested by Neumann (1995) and Wendt-Potthoff and Koschorreck (2002). Briefly, duplicates of freeze-dried sediment were subjected to chloroform-methanol extraction and the extracts were evaporated. The release of P from phospholipid was achieved by incubation at 510 °C for 5 h in a muffle furnace. Samples were then treated with sulfuric acid and ammonium molybdate, and 60 min after addition of malachite green solution photometric measurements at 610 nm were taken. To provide comparison with cultured bacterial populations, lipid-phosphate concentrations were converted to cells ml− 1 using sediment density and the conversion factors from Balkwill et al. (1988). 2.5. Analysis Immediately after sampling, pore water was extracted in the field laboratory by centrifugation and analyzed using test kits for Fe and SO2− 4 (Dr. Lange, Düsseldorf, Germany). Detection limits were 4 μmol L− 1 for Fe and 417 μmol L− 1 for SO2− 4 . Gas samples from the headspace of the glass vials were analyzed for CH4 and CO2 with a gas chromatograph equipped with a flame ionization detector and a methaniser. The measured values were converted to pore water concentrations (Remde and Tippmann, 1998). Water content and porosity were determined gravimetrically. Reactive Fe, including HCl-soluble Fe(II) and hydroxylamine-reducible Fe(III), was determined in triplicates using ferrozine (Lovley and Phillips, 1987). Samples were centrifuged (16,000 xg, 10 min) instead of filtration before photometric determination of iron. Total C, N and S were analyzed by high temperature combustion using an element analyzer (Vario EL, Elementar, Hanau, Germany). Stable carbon isotope analyses of CH4 and CO2 were conducted on a gas chromatography combustion isotope ratio monitoring mass spectrometer system. The system consisted of a GC (6890 Series, Agilent Technology, USA) coupled with a combustion interface (Thermo-Finnigan GC-combustion III, Bremen, Germany) and a Finnigan MAT 252 isotope ratio mass spectrometer (Thermo-Finnigan, Bremen). The methane in the GC effluent was oxidized to CO2 on a CuO/Ni/Pt catalyst held at 960 °C. A Poraplot Q column (0.32 mm × 25 m, Chrompack, The Netherlands) was used for separation of permanent gases. Helium at a flow rate of 1.5 ml min− 1 was used as carrier gas. The GC temperature program was held constant at 50 °C. Samples were injected in a hot split/ splitless (200 °C) injector with a split ratio varying between 1:1 and 1:50 according to the concentration of methane in the sample. Headspace injection volumes ranged from 0.05 to 1.0 ml based on the CH4 and CO2 concentration. The samples were acidified with HCl to pH b2 for the determination of the carbon isotope composition of dissolved inorganic

carbon (DIC). Samples were measured in at least in 3 replicates and the standard deviation was almost lower than 0.4‰. Microbial sulfate reduction was measured in small subcores in cut 5 ml syringes (Koschorreck et al., 2003) by 35S core injection (Meier et al., 2000). 2.6. Calculations Turnover rates and fluxes in different sediment layers were calculated from concentration profiles using the software PROFILE 1.0 (Berg et al., 1998). The software computes the rate of net production or consumption as a function of depth assuming molecular diffusion as the only transport process. The quality of the data is crucial for this procedure. Because of the low vertical resolution and the scatter of data a quantitative interpretation of the concentration profiles was not possible in all cases. Molecular diffusion coefficients were taken from Broeker and Peng (1974) and Li and Gregory (1974). 3. Results In 2003 the upper 6.5 cm of the sediment were light grey, most likely from volcanic ashes of the near volcano Copahue which had

Fig. 5. a) Element composition in the long sediment core. b) Total microbial cells determined from the lipid-phosphate content. The thin horizontal lines separate different core segments.

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Fig. 6. Vertical profiles of H2S (a) dissolved sulfate (b), dissolved ferrous iron (c), and dissolved ferric iron (d) in the sediment of Lake Caviahue. The H2S concentrations between 4 and 20 cm exceeded the measuring range of the H2S electrode during the field work in 2004 (■ 2001, ● 2003, ▼ 2004).

erupted in July 2000. Below this layer the sediment was greenish and organic-rich down to depth between 20 and 22 cm. The organic material was likely from dead algae. The presence of algae (mostly diatoms) was verified microscopically (M. Diaz, pers. com.). Below 22 cm the sediment became greyer with greenish layers. The pH of the sediment increased from pH 2.5 at the sediment surface to ~3.7 at 30 cm depth (Fig. 1a). The vertical distribution of dissolved gases in sediment cores in 2003 was fairly similar to the distribution monitored during the campaigns in 2001 and 2002 (Fig. 1b, c), indicating quite stable conditions in the sediment. The methane concentration increased linearly with depth in the upper 30 cm of the sediment (Fig. 1c) which indicates a source of methane below this depth. Applying a linear fit to the methane profiles we obtained a methane flux of 0.93 ± 0.45 mmol m− 2 d− 1 from the deeper parts of the sediment to the surface. Using the software PROFILE (Berg et al., 1998) we identified zones of production and consumption in the sediment (Fig. 2). Low rates of methane consumption were found in the surface layers (Fig. 2e). The production rates calculated from the vertical profiles were not significantly different from zero. The activity assays did not show a net methane production in the upper 30 cm of the sediment (Fig. 3) and methanogenic bacteria were not detected in our MPN tubes. Unfortunately we did not obtain a positive control during the investigation of the short core in Caviahue but the MPN procedure was tested in our laboratory and produced reliable data. To assess the source of methane in the sediment of Lake Caviahue in 2003 a 6.4 m long sediment core was investigated (Figs. 3, 5). The pH increased to a constant value of 5.6 in 90 cm depth. The methane concentration increased with depth from values of lower than 0.5 mmol L− 1 in the upper layer to concentrations up to 3 to 6 mmol L− 1 at 30 cm depth. Below 30 cm, the concentration decreased down to a relatively constant value of about 1 mmol L− 1 at 1 m depth (Fig. 3b). The highest methane concentrations were observed in some samples below 4 m. The activity assays showed methanogenic

activity between 70 cm and 180 cm and no methanogenic activity at the surface or at greater depth (Fig. 3). The carbon isotopic composition of methane and DIC was analyzed to characterize sinks and sources of methane (Fig. 4). The carbon isotope composition of CH4 was between −63.4 and −74.3‰ which is typical for biogenic methane (Whiticar, 1999). The isotope fractionation (ε) between DIC and methane in the upper section of the core was between −46 and −48‰ in the depth interval between 0 and 12.5 cm depths, respectively. At 12.5 to 29.5 cm depth the isotope composition of methane was between −69.3 and 66.2‰ and the isotope fractionation was between 53 and 59‰. This may indicate that methane formation was dominated by CO2 reduction with a minor contribution of methane from acetoclastic methanogenesis. ε values further decreased with depth to more or less constant values between −70 and −80‰ below 1 m depth (Fig. 4) indicating an increasing contribution of methane from microbial CO2 reduction. The δ13CH4 values of methane in deeper layers were between −63‰ and −74‰ and constant with depth (Fig. 4b). δ13DIC increased from −17‰ at the surface to constant values around 5‰ below 1 m depth (Fig. 4a). The CO2 concentrations in different cores were more heterogeneous than the methane concentrations (Fig. 1b). The heterogeneity of two cores from the same sampling date in 2001 was in the same range as the inter-annual variability. The general pattern was that CO2 increased to a relatively constant value at a depth between 5 and 20 cm. CO2 profiles were difficult to fit because of scattered data (Fig. 2d). The data indicate CO2 production throughout the upper 30 cm with lower rates at the surface and below 20 cm. CO2 showed a broad maximum between 1.5 and 3.5 m depth and very high values at 4.1 m (like CH4) and 6.3 m depth (Fig. 3a). The high gas concentrations at 4.1 m depth were accompanied by high total cell counts (Fig. 5b). We also observed high contents of C and N below 4.5 m depth (Fig. 5a). Analysis of a grab sample from the surface showed that most of the carbon was organic (63 g kg− 1 organic C versus 7 g kg− 1 inorganic C).

Table 1 Fluxes at the sediment surface and production rates in different sediment layers (mmol m− 2 d− 1) calculated with PROFILE. Positive fluxes are directed from the sediment into the water cm

O2

Flux at top 0–6.5 6.5–12.5 12.5–24 Sum a

−16.8 ± 9.8 −16.8 ± 9.8

CO2 1.1 ± 0.9 −0.12 2.1 −0.43 1.55

Not different from zero (error bars include zero).

CH4 a

−0.3 ± 0.4 −0.57 ± 0.48 − 0.96 ± 1.45a −1.46 ± 3.13a −3 ± 4.3a

H2S

S2− 4

Fe2+

H+

0.5 ± 0.3 0.4 ± 0.3 0.1 0.1 0.8

4.6 ± 2.5 5.3 ± 3.2 −0.9 ± 1.8a − 0.1 ± 0.9a 4.2 ± 2

−2.6 ± 1.2 3.4 ± 2 −0.1 ± 0.5a −0.6 ± 0.4 2.7 ± 1.2

−1.7 ± 1.9a −0.2 ± 0.7a −0.3 ±0.5a −0.6 ± 0.5 −1.0 ± 1.7a

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Fig. 7. Microprofile of O2 (black lines) and H2S (grey dotted lines) measured in a sediment core from the profundal of Lake Caviahue.

The total sulfur content of the sediment was uniformly high down to 2.6 m and fluctuated between 0 and 30 g kg-dw− 1 below this depth. The pore water concentration of H2S showed a maximum between 5 and 20 cm depth (Fig. 6a) confirming earlier observations of microbial sulfate reduction in this sediment layer (Koschorreck et al., 2003). The modeling of the H2S profiles indicates a net production of H2S in the upper 6 cm and below 25 cm (Fig. 2f). The H2S distribution may be influenced by microbial production and consumption of H2S as well as changed solubility at different pH and the variable availability of iron. The decrease of H2S below 20 cm depth was probably caused by higher precipitation rates of sulfidic minerals at higher pH. Sulfate reduction rates as measured by 35S core injection were 5.2 ± 1.5 nmol ml− 1 d− 1 in the 0–6.5 cm layer which is in the range of previously published values (Koschorreck et al., 2003). The profiles of ferrous iron as well as sulfate showed a distinct peak around 5 cm depth (Fig. 6b, c) indicating that there was a zone of ferrous iron and sulfate production in this sediment layer where iron and sulfate were diffusing upward into the water and downward into the sediment. From the zone of highest concentrations upward fluxes −2 −1 of 2.6 mmol Fe m− 2 d− 1 and 4.6 mmol SO2− d and downward 4 m −2 −1 fluxes of 0.8 mmol Fe m− 2 d− 1 and 0.7 mmol SO2− d were cal4 m culated. The calculations indicate, that ferrous iron and sulfate were produced in the upper 6 cm of the sediments with rates of 59 ± 35 and 81 ± 49 nmol ml− 1 d− 1 (Fig. 2b, c). Protons were consumed with a rate of 11.6 nmol ml− 1 d− 1 throughout the upper 20 cm of the sediment (Fig. 2a). The proton flux of 1.7 ± 1.9 mmol H+ m− 2 d− 1 was not significantly different from zero. From the pore water profiles we calculated fluxes at the sediment surface (Table 1). The vertical resolution of the pore water profiles, however, was not high enough to resolve the processes at the water– sediment interface. To get information about possible oxidative processes at the sediment surface we measured microgradients of O2 and H2S with microsensors (Fig. 7). Oxygen penetrated about 1 mm deep into the sediment. From the gradients in the diffusive boundary layer a diffusive flux of 16.8 ± 9.8 mmol O2 m− 2 d− 1 was calculated. H2S was found in the anoxic part of the sediment. There was a small overlap

with the oxic zone indicating H2S oxidation by oxygen (Fig. 7). From the linear part of the H2S microprofiles an upward flux of 0.9 ± 0.4 mmol m− 2 d− 1 was calculated. Since 2 mol of O2 are necessary to oxidize 1 mol of H2S to SO2− 4 sulfide oxidation can account for 11% of the total oxygen consumption. The vertical resolution of the methane profile was not sufficient to prove if methane was completely oxidized in the sediment or if some methane enters the water column. The very low methane concentration in the profundal water (b1 μmol L− 1) is a hint that not much methane left the sediment. If we assume that all methane was oxidized in the oxic part of the sediment, methane oxidation contributed 4% to sediment oxygen consumption. Another 4% of the oxygen may be used for the oxidation of upwardly diffusing ferrous iron. We found extremely high counts of thiosulfate-oxidizing bacteria in the uppermost centimeter (Table 2). Numbers of iron oxidizers were also highest at the sediment surface but 8 orders of magnitude lower than thiosulfate oxidizers. The numbers of iron and sulfate reducing bacteria were below 104 cells ml− 1. Acidophilic iron reducers decreased with sediment depth while sulfate reducing bacteria were most abundant between 1 and 11 cm. The total cell counts as determined from lipid-phosphate analysis (Fig. 5) decreased with depth from 1011 cells g-dw− 1 at the surface to still high values of about 109 cells g-dw− 1 in 6 m depth. One has to consider, however, that our measurements are likely to overestimate living biomass since they include not only living bacteria but also lipid phosphate derived from dead but not degraded biomass. 4. Discussion Methane in the sediment of Lake Caviahue was produced by microbial methanogenesis and not by geochemical volcanic activity. There was, however, no methanogenic activity observed in the upper part of the sediment. One possible explanation would be that the methanogens were outcompeted for electron donors by the sulfate reducers. Sulfate reduction, however, was restricted to the upper 12.5 cm of the sediment. Microbial activity in general was lower below this depth as indicated by the CO2 profiles. On the other hand, the high CO2 values, the high carbon content and the presence of viable biomass in deeper sediment layers (Fig. 5b) indicate that organic electron donors were probably available. This is also supported by the batch assays, which proved methanogenesis down to more than 1 m depth. The most probable reason for the absence of methanogenesis in the upper sediment is the low pH in this zone. This is consistent with the literature which implies a lower pH boundary for the growth of methanogens between pH 3.8 and 5 (Williams and Crawford, 1985; Lapado and Barlaz, 1997; Hornibrook et al., 2000; Bräuer et al., 2004). Rates of methanogenesis in sediments of 20 neutral lakes were between 0.7 and 100 mmol m− 2 d− 1 (Koschorreck and Tittel, 2007). In comparison the CH4 flux to the surface in Lake Caviahue was only 0.93 ± 0.45 mmol m− 2 d− 1, indicating a relatively low CH4 net production in the methanogenic zone. Compared with the total consumption of oxygen in the sediment methanogenesis was involved in less than 4% of organic matter mineralization and thus, was of minor importance for the overall carbon budget of the sediment. Potential activity and viable biomass reached rather deep into this sediment. In Lake Constance, for example, microbial activity including

Table 2 MPN counts of different bacterial groups [bacteria ml− 1] and total cell numbers calculated from lipid-phosphate content Depth (cm)

Iron oxidizers

Sulfur oxidizers

Iron reducers neutrophilic

Iron reducers acidophilic

Sulfate reducers

Total

0–1 1–6 6–11 17–22

2.4 ± 1.6 × 105 5.3 ± 2.5 × 103 2.0 ± 1.4 × 103 2.3 ± 1.6 × 103

1.6 ± 1.0 × 1013 5.3 ± 2.5 × 104 1.2 ± 0.4 × 105 3.6 ± 2.3 × 103

2.0 ± 1.0 × 102 0.9 ± 0.6 × 102 8.0 ± 4.4 × 102 1.0 ± 0.7 × 102

8.6 ± 5.8 × 103 4.3 ± 1.3 × 103 2.2 ± 1.0 × 102 1.6 ± 1.0 × 102

7.9 ± 4.1 × 102 2.2 ± 1.0 × 103 1.5 ± 1.0 × 103 1.0 ± 0.7 × 102

1.1 × 1011 1.3 × 1011 7.6 × 1011 8.6 × 1011

M. Koschorreck et al. / Journal of Volcanology and Geothermal Research 178 (2008) 197–204 Table 3 Biogeochemical zonation of the sediment of Lake Caviahue Depth

Processes

0–0.2 m

Aerobic respiration Oxidation of H2S Oxidation of CH4 Probably oxidation of Fe2+ Dissolution of iron-sulfur minerals Production of H2S = sulfate reduction Low activities Methanogenesis Low activities

0–6.5 cm 6.5–40 cm 40–ca. 150 cm below 150 cm

methanogenesis could not be detected below 50 cm sediment depth (Rothfuss et al., 1997). The carbon content of the deep sediment of Lake Constance, however, is below 1% (IGKB, 2004). At sites with high organic content like peat bogs methanogenesis may take place in greater depth (Kotsyurbenko et al., 2004). We are unaware of any spore-forming methanogens which means that the cells in the deep sediment layers must have survived in vegetative state for a long time. This is supported by the observation that autotrophic methanogens are active in deep crystalline rock aquifers with only traces of organic matter (Stevens and McKinley, 1995). In addition, functional enzymes associated with live bacteria have been detected in over 124,000 years old Mediterranean sapropels (Coolen and Overmann, 2000). Our data suggest that methane was oxidized in the oxic zone either in the sediment or in the lake water. Methane oxidation at low pH has been reported from peat bogs (Dedysh, 2002), and it could be induced in soil at pH 2.3 (Bender and Conrad, 1995). Recently a biogenic methane oxidation potential has been detected in volcanic soils at pH 1.9–2.6 (Castaldi and Tedesco, 2005). The lower limit for known pure cultures, however, is about pH 4. Thus, either the responsible organisms or their special requirements for methane oxidation at extremely low pH are still unknown and require further investigation. The carbon isotope composition of methane did not show a clear indication for methane oxidation. Methane oxidation is associated with a relatively strong enrichment of 13C regardless of aerobic or anaerobic degradation. The methane isotope composition between −70 and −80‰ characterizing deeper sediment was continuously enriched to values of about −63.4‰ in the uppermost sediment layer. However, when comparing the concentration and isotope composition in the upper layers between 0 and 30 cm the enrichment of methane carbon isotope in the upper 12 cm is only in the order of some permil 13C and a trend for continuous enrichment of carbon isotopes upwards the upper sediment column is not so clear suggesting only minor methane oxidation in that zone. Mixing with atmospheric methane with isotope values of N40‰ due to circulating lake water may also be responsible for the carbon isotope enrichment of methane and the depletion of concentration. Alternatively, the slightly heavier isotope composition of methane in the upper part and the lower fractionation factors may point to a higher contribution of acetoclastic methanogenesis. Acetate concentrations in the pore water were as high as 0.35 mmol L− 1 (Koschorreck et al., 2003). A transition from acetoclastic methanogenesis to CO2 reduction with depth reflected by the methane depleted in 13C was observed in peat bog environments (Hornibrook et al., 2000; Kotsyurbenko et al., 2004). In Lake Caviahue, this is likely because more recent labile, sedimentary biological material from planktonic sources might be present in the upper part of the sediment giving advantage to fermenting bacteria. When the labile organic material is consumed CO2-dependent methanogenesis dominates. A similar vertical distribution of acetotrophic and hydrogenotrophic methanogenesis has also been found in a neutral freshwater lake, and it was accompanied by a decrease of Methanosaetaceae and an increase of Methanomicrobiales with depth (Chan et al., 2005). Methanogenesis by CO2 reduction requires hydrogen. However, it is unknown if hydrogen in deeper sections is of volcanic origin or was supplied by fermentation processes.

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Considering all results we may divide the sediment into zones of different activities (Table 3). The vertical sequence of the microbial redox processes follows the classical sequence determined by thermodynamic energy yield of the particular reaction. This is also supported by our MPN counts which showed highest numbers of oxidizing bacteria at the surface, higher numbers of sulfate reducing bacteria in the sulfate reduction zone and generally low numbers of bacteria of the iron cycle. The zone of low activity between the sulfate reducing and the methanogenic zone might be related to the depositional environment. In phase of volcanic activities sediments with low organic material may deposit leading to lower heterotrophic microbial activity. The material in the zone of low activity was probably also deposited during an eruption of Copahue Volcano, as it contained several grey layers similar to the ash layer on top of the sediment. The upper 30 cm of the sediment, however, were rich in organic carbon and sulfate reduction was probably not limited by organic carbon (Koschorreck et al., 2003). Unfortunately we have no sulfate data from deeper sediment layers. If we extrapolate the existing sulfate profiles (Fig. 6b) we may estimate that sulfate was depleted between 25 and 30 cm depth. Below this depth sulfate reducers could not compete with methanogenes for electron donors. Although sulfate reduction consumed sulfate in the upper sediment there was a net release of SO2− 4 to the pore water. The most probable explanation is the dissolution of sulfate containing minerals in this layer. A possible scenario is that sulfate reduction raised the pH in the sediment which in turn lowered the stability of sulfur containing minerals such as melanterite (FeSO4) and led to a release of sulfate into the pore water. Melanterite could be a component of the volcanic ash. A source of sulfate was also observed in an acidic mining lake after artificially rising the pH in the sediment, in that case due to the dissolution of jarosite (Herzsprung et al., 2002). Melanterite dissolution would also explain the observed peak of Fe2+ in the pore water. Microbial Fe(III) reduction is unlikely in this sediment, since the Fe2+ peak increased from 2001 to 2003 and Fe(III) at was hardly found at this depth (Wendt-Potthoff and Koschorreck, 2002). Altogether, these observations indicate that the surface sediment was not in a steady state which may be due to the volcanic eruptions in 2000 which killed plankton in the lake (Pedrozo et al. this issue). As a result high amounts of biomass reached the sediment. By this mechanism volcanic eruptions stimulate sulfate reduction which raises the pH in the sediment and depletes sulfate, thereby creating favorable conditions for methanogenesis in deeper sediment layers. Acknowledgements Many thanks to Martin Wieprecht, Michael Herzog and Anne Müller for their help during the field trips. Matthias Gehre and Ursula Günther are acknowledged for the help in the isotope laboratory of the Department for Isotope Biogeochemistry. Special thanks to Fernando Pedrozo and his group for the excellent support. This study was financially supported by the VW-foundation (project I78/847). References Balkwill, D.L., Leach, F.R., Wilson, J.T., McNabb, J.F., White, D.C., 1988. Equivalence of microbial biomass measures based on membrane lipid and cell wall components, adenosine triphosphate, and direct counts in subsurface aquifer sediments. Microb. Ecol. 16, 73–84. Bender, M., Conrad, R., 1995. Effect of CH4 concentrations and soil conditions on the induction of CH4 oxidation activity. Soil. Biol. Biochem. 27, 1517–1527. Berg, P., Risgaard-Petersen, N., Rysgaard, S., 1998. Interpretation of measured concentration profiles in sediment pore water. Limnol. Oceanogr. 43, 1500–1510. Boetius, A., Ravenschlag, K., Schubert, C.J., Rickert, D., Widdel, F., Gieseke, A., Amann, R., Jørgensen, B.B., Witte, U., Pfannkuche, O., 2000. A marine microbial consortium apparently mediating anaerobic oxidation of methane. Nature 407, 623–626. Botz, R., Pokojski, H.D., Schmitt, M., Thomm, M., 1996. Carbon isotopic fractionation during bacterial methanogenesis by CO2 reduction. Org. Geochem. 25 (3–4), 255–262. Bräuer, S.L., Yavitt, J.B., Zinder, S.H., 2004. Methanogenesis in McLean Bog, an acidic peat bog in upstate New York: stimulation by H2/CO2 in the presence of rifampicin, or by low concentrations of acetate. Geomicrob. J. 21, 433–443.

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