Sulphur isotope composition of dissolved sulphate in the Cambrian–Vendian aquifer system in the northern part of the Baltic Artesian Basin

Chemical Geology 383 (2014) 147–154

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Sulphur isotope composition of dissolved sulphate in the Cambrian–Vendian aquifer system in the northern part of the Baltic Artesian Basin Valle Raidla a,⁎, Kalle Kirsimäe b, Jüri Ivask a, Enn Kaup a, Kay Knöller c, Andres Marandi b, Tõnu Martma a, Rein Vaikmäe a a b c

Institute of Geology at Tallinn University of Technology, Ehitajate tee 5, 19086 Tallinn, Estonia Department of Geology, University of Tartu, Ravila 14a, 50411 Tartu, Estonia UFZ Helmholtz Centre for Environmental Research, Department of Catchment Hydrology, Theodor-Lieser-Str. 4, 06120 Halle, Germany

a r t i c l e

i n f o

Article history: Received 5 March 2014 Received in revised form 13 June 2014 Accepted 17 June 2014 Available online 26 June 2014 Editor: David R. Hilton Keywords: Groundwater Bacterial activity Sulphide oxidation Sulphate reduction

a b s t r a c t The groundwater in the Cambrian–Vendian aquifer system with its δ18Owater values of about −22‰ (VSMOW) and a low radiocarbon concentration is of glacial origin from the Last Ice Age. Earlier surveys have highlighted a negative co-variance of sulphate and bicarbonate content in the groundwater of the Cambrian–Vendian aquifer system, whereas the most depleted dissolved inorganic carbon δ13C values have been measured mainly in groundwater samples with the lowest sulphate concentrations. In this paper we studied the origin of sulphate and the factors controlling the sulphur and carbon isotope geochemistry in the aquifer system. Direct sources of sulphate were not found, but relying upon δ18OSO4 measurements we suggest that the sulphate originates from oxidation of sulphide minerals whereas the δ34S of the dissolved SO2− 4 in the groundwater is more enriched than the δ34S of the surrounding rocks. We show that bacterial activity may have caused the enrichment of δ34S of sulphate. © 2014 Elsevier B.V. All rights reserved.

1. Introduction The Cambrian–Vendian (Cm–V) aquifer system is a confined waterbody in western and north-western parts of the East-European Platform. In shallowly buried marginal areas of the aquifer system, particularly in northern part of the Baltic Artesian Basin, the groundwater is fresh and widely used in public water supply. The fresh groundwater at the northern margin of the basin has the lightest known oxygen isotopic composition in Europe (δ18Owater values of around −22‰) and a low radiocarbon concentration suggestive of glacial origin of the water (e.g. Vaikmäe et al., 2001). Raidla et al. (2009) showed that the groundwater from the Cm–V aquifer system is a mixture of three end-members — fresh glacial meltwater, relict Na–Ca–Cl brine and recent meteoric water. The massbalance model of the carbonate system coupled with the modelling of radiocarbon age of the groundwater from the Cm–V aquifer system (Raidla et al., 2012) shows that the infiltration of the water occurred not earlier than 14,000 to 27,000 radiocarbon years ago, which is coeval with the advance and maximum extent of the Weichselian Glaciation in the area (Kalm, 2012).

⁎ Corresponding author. E-mail address: [email protected] (V. Raidla).

http://dx.doi.org/10.1016/j.chemgeo.2014.06.011 0009-2541/© 2014 Elsevier B.V. All rights reserved.

In this paper we study the sulphur isotope composition in the groundwater of the Cm–V aquifer system at its northernmost margin in North-Estonia. Our goal is to reveal the origin of sulphate and the factors controlling the sulphur and carbon isotope geochemistry in the aquifer system's water and rock matrix, and to test evidence of bacterial activity in the aquifer system's hydrogeochemistry. Raidla et al. (2012) put forward a hypothesis that, although the origin of the degradable organic matter and sources of sulphate in the groundwater are virtually unknown, it is possible that the isotope system has been modified by bacterial activity, most likely via sulphate reduction. 2. The study area The Cambrian–Vendian aquifer system is the shallow northern part of the Baltic Artesian Basin (BAB), which completely underlies the Baltic States and partly the border areas of Russia, Poland and Belarus. A large part of the BAB lies beneath the Baltic Sea. The Cm–V aquifer system encompasses the thick (up to 90 m) sequence of Ediacaran and Cambrian sandstones alternating with clays and silty clays (Fig. 1). The Ediacaran and Cambrian sedimentary rocks outcrop into the Gulf of Finland and to the bottom of the Baltic Sea. The aquifer system is confined by the underlying crystalline basement of the Palaeoproterozoic age and the overlying Lontova–Lükati aquitard, and the groundwater is under pressure. In the northeastern Estonia the aquifer system is divided by

148

V. Raidla et al. / Chemical Geology 383 (2014) 147–154

a

b

c

Fig. 1. A geological scheme of the northern Baltic Artesian Basin in Estonia with the positions of the studied wells (a), the West–East cross-section of northern Baltic Artesian Basin (b) and the North–South cross-section of the northern Baltic Artesian Basin (c).

Ediacaran clays of the Kotlin age into two aquifers: the upper Voronka (V2vr) and the lower Gdov aquifer (V2gd) (Fig. 1). In northern Estonia the conductivity of the water-bearing rock is 0.5 to 9.2 m d−1, with the average of 5 to 6 m d−1. Transmissivity in northeastern Estonia is 300 to 350 m2 d−1, decreasing in southerly and westerly directions (Perens and Vallner, 1997). The Lontova–Lükati aquitard, overlying the aquifer system is composed of silty clays, siltstones and clays of the Lower Cambrian age. The Lower Cambrian clays are diagenetically immature and plastic with water content of about 20 to 30% (Kirsimäe and Jørgensen, 2000; Raidla et al., 2006). The thickness of the clayey complex is 90–100 m in North-Estonia, but decreases towards the south until disappearing in South-Estonia (Fig. 1). The aquitard has a strong isolation capacity with vertical hydraulic conductivity of 10− 7 to 10− 5 m d− 1 (Perens and Vallner, 1997). The Lontova–Lükati clays are gradually replaced by interbedded clay and sandstone in western part of the area and their

vertical hydraulic conductivity is N10− 5 m d− 1 (Perens and Vallner, 1997). At the North-Estonian coastline the aquitard is incised by deep valleys filled with loamy till and glaciofluvial gravel (Tavast, 1997). These valleys serve as recharge areas where the water from the upper groundwater horizons infiltrates to the Cm–V aquifer system (Fig. 1). In the most part of northern and central Estonia, the siliciclastic rocks of the Cm–V aquifer system are covered by up to 300 m thick layer of Ordovician and Silurian marine carbonate rocks. The groundwater in northern part of the aquifer system is of Cl– HCO3–Na–Ca and Cl–HCO3–Ca–Na type with the TDS content between 0.4 and 1.0 g L−1 (Savitskaja and Viigand, 1994). In southern part of the aquifer system the groundwater is replaced by saline relict Na–Cl water with TDS values of up to 22 g L−1 (Karise, 1997). The most characteristic property of the groundwater from the Cm–V aquifer system is its stable isotope composition. δ18Owater values are between −18.5 and − 22‰ VSMOW (Vaikmäe et al., 2001), whereas the isotope

Table 1 Chemical composition of the groundwater in the northern part of the Cm–V aquifer system. Cm–V — Cambrian–Vendian aquifer; V2gd — Gdov aquifer; V2vr — Voronka aquifer; * — chemical results by Perens and Boldureva (2008). Well no

Well ID

Location

Aquifer

Depth

Date

m 3344 9997 14927 552 14146 14784 8914 10703

9 10 11 12 13 14 15 16 17 18

10714 2738 2775 3323 2388 2386 2233 2249 2198 2217

19 20

2197 2207

21 22

2470 2092

23 24

2084 3434

Haapsalu City* Risti village* Murru prison Keila City* Tallinn. Kopli port Tallinn. Kopli port Muuga port* Rakvere City Rakvere City Arkna water intake Rakvere City Kunda City Unukse village* Aseri City Aseri City Purtse village Kohtla-Nõmme City Sillamäe City Sillamäe City* Sillamäe City Sillamäe City Sillamäe City* Sillamäe City Viivikonna mining Narva-Jõesuu City Narva-Jõesuu City Narva-Jõesuu City Narva mining

Cm–V Cm–V Cm–V Cm–V Cm1ln Cm–V Cm–V V2vr V2vr Cm–V V2gd V2gd V2gd V2gd V2gd V2vr V2gd V2vr V2vr V2vr V2vr V2gd V2gd V2vr V2vr V2vr V2gd V2vr

295 280 200 214 60 65.5 75 222 222 235 268 205 207 185 200 130 255 150 124 124 151 220 220 180 101 101 211 180

13–Mar–07 13–Mar–07 1–Oct–10 19–Mar–07 22–Mar–07 22–Mar–07 27–Mar–07 25–Sep–07 1–Nov–11 1–Nov–11 1–Nov–11 29–Nov–11 12–Nov–07 16–Jan–08 29–Nov–11 29–Nov–11 8–Jun–07 1–Dec–11 7–Jun–07 1–Dec–11 01–Dec–11 07–Jun–07 01–Dec–11 1–Dec–11 24–Sep–07 3–Nov–11 3–Nov–11 3–Nov–11

El. cond. −1

Eh

°C

μS · cm

mV

10.8 9.8 9.0 9.3 9.1 8.4 9.3 11.9 12.3 11.1 11.8 10.6 11.7 10.9 10.9 8.5 14.0 9.9 10.5 10.1 10.1 12.4 11.9 11.5 9.6 9.1 10.8 11.4

292 329 551 519 306 469 728 322 443 349 586 275 1027 548 258 233 547 360 496 285 825 1041 650 339 524 410 1614 410

−108.0 −110.0 −134.0 −118.0 −79.1 −108.0 −107.8

−60.5 −96.5 −69.6 −65.7 −87.6 −35.1 −96.0 −69.3 −97.2 −156.0 −79.3 −87.2 −118.8 −86.4 −98.9

pH

δ34S ‰ VCDT

8.6 8.6 8.9 7.7 7.9 7.8 8.0 7.9

7.6 7.8 7.8 7.7 8.0 7.3 8.2 7.8 8.1 7.9 7.5 7.9 8.1 8.5 8.2 8.1 8.3

8.3 10.1 12.0 11.0 24.9 28.8 17.5 20.9

30.5 19.7 20.4 28.3 26.8 55.1 27.3 43.6 23.6 20.5 25.7

δ18OSO4

δ18Owater

δD

‰ VSMOW −5.3 2.1 1.2 −2.4

−8.2

−10.8 −9.6

6.0 8.9

0.0

9.7

−20.2 −20.4 −20.4 −21.3 −20.6 −20.7 −22.0 −20.1 −20.0 −20.5 −20.2 −21.4 −21.1 −21.3 −21.1 −20.1 −19.7 −18.1 −19.4 −19.4 −18.7 −18.6 −18.6 −18.9 −18.4 −18.3 −18.4 −19.0

δ13C ‰ VPDB

mg L 38.1 38.3 41.7 40.9 53.7 48.9 33.3 34.0 72.5 43.4 96.8 73.0 71.3 66.5 84.9 59.1 49.1 49.9 20.0 17.0 30.4 18.6 25.7 22.7 10.8 11.9 54.2 10.2

−158.6 −151.6 −136.7 −146.2 −140.4 −140.0 −143.2 −138.4 −138.0 −143.7

−12.3 −13.2 −13.9 −17.6 −18.3 −18.7 −17.5 −19.4 −14.3 −17.1 −21.5 −22.0 −15.7 −20.2 −17.8 −17.0 −17.3 −18.1 −21.7 −21.4 −17.4 −20.9 −21.2 −30.9 −19.5

Mg2+

Na+

K+

Cl−

SO2− 4

HCO− 3

13.3 9.0 9.3 11.8 20.2 19.6 10.7 16.0 20.2 14.1 23.4 25.0 21.4 17.7 21.0 18.2 20.1 22.3 2.4 5.6 11.1 6.6 8.8 7.6 4.1 3.9 20.3 2.2

86.2 73.8 65.5 57.2 30.0 31.8 21.7 98.0 81.0 72.1 92.9 140.0 116.7 110.0 104.5 92.5 140.0 259.2 197.0 157.8 205.0 388.9 395.8 157.8 206.3 143.7 461.7 145.4

6 6.0 6.8 8.6 10.8 9.4 6.7 8.5 7.4 5.8 8.5 10.0 10.5 8.5 8.5 8.6 8.0 8.6 4.5 6.1 8.0 7.0 10.7 9.0 4.0 5.7 13.8 5.7

155.3 139.3 125.3 123.0 118.1 92.5 42.2 150.0 81.0 72.1 92.9 316.3 239.3 212.0 104.5 179.6 275.5 483.0 192.2 157.8 318.4 588.9 395.8 201.5 225.1 143.7 461.7 145.4

32.1 28.0 18.0 26.4 b3.3 10.0 b3.3 0.5 1.8 b0.1 1.7 b0.1 b3.3 5.8 b0.1 b0.1 9.5 16.7 1.9 b0.1 2.7 b3.3 3.4 b0.1 12.8 b0.1 b0.1 b0.1

109.8 109.8 115.9 122.0 164.7 183.0 170.8 201.3 161.7 176.9 173.9 244.1 237.9 250.2 247.1 201.4 152.5 219.7 231.9 244.1 219.7 152.5 164.7 216.6 195.2 201.3 183.0 234.9

−1

−10.8 −12.0 −153.0

−150.8 −153.8 −151.7 −162.0

Ca2+

V. Raidla et al. / Chemical Geology 383 (2014) 147–154

1 2 3 4 5 6 7 8

Tem

149

150

V. Raidla et al. / Chemical Geology 383 (2014) 147–154

composition of the modern precipitation is − 8 to − 11‰ VSMOW (Punning et al., 1987). Low 14C activities show that the water was last exposed to atmosphere over 10,000 years ago. A sub-glacial recharge mechanism is the most probable explanation of how the depleted water reached the Cm–V aquifer system. During the Pleistocene the continental ice sheet that covered the area would have increased hydraulic pressure at the base of the glacier, which would have reversed the regional groundwater flow. The inflow of diluted glacial meltwater into the aquifer occurred along the outcrop area at the basin's margin (Raidla et al., 2009). 3. Material and methods Groundwater was sampled and analysed from water supply and observation wells during two fieldwork campaigns in 2007 to 2008 and 2010 to 2011 in North-Estonia (Fig. 1. Table 1). The amount of water collected for sulphate isotope analyses ranged from 5 to 100 L, depending on the quantity of dissolved sulphate. The samples were pre-filtrated and acidified/stirred to remove carbonates. The dissolved sulphate was precipitated using BaCl2·2H2O. The precipitated BaSO4 was collected by filtration through nitrocellulose membranes, washed to remove residual BaCl2 and dried at 50 °C. In thirteen samples collected in 2007 and 2008, sulphur isotopic compositions were measured after conversion of BaSO4 to SO2 using an elemental analyser (continuous flow flash combustion technique) coupled with an isotope ratio mass spectrometer (Delta S, ThermoFinnigan, Bremen, Germany) at the Stable Isotope Laboratory of the UFZ in Halle/Saale, Germany. Samples collected in 2011 were analysed at the Isotope Science Lab in Calgary, Canada, by using the same technique, but a different mass spectrometer/elemental analyser combination (Thermo Finnigan DeltaPLUS XL and Fison NA1500). Sulphur isotope measurements were performed with an analytical error of the measurement of better than ± 0.3‰ and results are reported in delta notation (δ34S) as part per thousand (‰) deviation relative to the Vienna Cañon Diablo Troilite (VCDT) standard. Oxygen isotope analysis of sulphate was performed only for 11 samples for which high temperature pyrolysis at 1450 °C in a TC/EA connected to a delta plus XL mass spectrometer (ThermoFinnigan, Bremen, Germany) with an analytical precision of better than ±0.5‰ was used. Results of oxygen isotope measurements are expressed in delta notation (δ18OSO4) as part per thousand (‰) deviation relative to Vienna Standard Mean Ocean Water (VSMOW). For normalizing the δ34S and δ18OSO4 data, the IAEA-distributed reference material NBS 127 (BaSO4) was used. The assigned values were + 20.3‰ (VCDT) δ34S and +8.6‰ (VSMOW) for δ18OSO4. Measurements of stable isotopes of δ18O and δD in the water samples and δ13C in precipitated carbonates were performed at the laboratory of mass spectrometry of the Institute of Geology at Tallinn University of Technology using a Thermo Fisher Scientific Delta V Advantage mass spectrometer and GasBench II. Reproducibility was better than ±0.1‰ for δ18Owater and ±0.5‰ for δ13C. The results are expressed in ‰ deviation relative to Vienna Standard Mean Ocean Water (VSMOW) and Vienna Peedee Belemnite (VPDB) for O, D and C, respectively. For normalizing the δ18Owater and δ13C data, the IAEA-distributed reference materials VSMOW, SLAP, NBS 19 and LSVEC were used. The pH of δ13C sample was set to pH N 12 by adding the required amount of concentrated CO2-free NaOH solution. Analytical grade BaCl2·2H2O in excess of expected DIC was added to precipitate DIC in the form of BaCO3. The results are expressed in ‰ deviation relative to Vienna Standard Mean Ocean Water (VSMOW), Vienna Peedee Belemnite (VPDB) and the Cañon Diablo Troilite (VCDT) standard for O, C and S, respectively. Major ions have been measured on Dionex ICS-1000 ion chromatograph at Tartu University, except for bicarbonate, which was measured in the Laboratory of Geological Survey of Estonia by the titration

method. In addition, some results from earlier studies were used (Perens and Boldureva, 2008). 4. Results and discussion 4.1. Distribution of dissolved sulphate The sulphate content in the studied samples from the Cm–V aquifer system is generally low ranging from below detection limit (0.1 mg L− 1) to 32.1 mg L− 1 (Table 1). Raidla et al. (2012) noticed that the sulphate concentration was the highest in Western-Estonia and low or almost undetectable in the central and eastern parts of the study area. δ34S values of the dissolved sulphate in the groundwater of the Cm–V aquifer system vary between 8.6 and 55.1‰ (average 24.0‰). δ 18 O SO 4 values of the dissolved sulphate range between − 10.8‰ and 9.7‰ (average − 0.2‰) while δ18Owater and δDwater values of the groundwater samples vary between − 18.4‰ and − 21.4‰ and − 136.7‰ and − 162.0‰, respectively (Table 1). 4.2. Direct sulphate sources No direct sources of SO24 −, such as gypsum or mirabilite deposits, occur in the aquifer system's rock matrix (Puura et al., 1983; Raidla et al., 2006). In our opinion, there are two possible sources of sulphate: (a) relict basinal brine or (b) oxidation of sulphide minerals. Chemical composition of the Cm–V aquifer system varies both laterally and vertically. The shallow diluted water along the northern margin of the Baltic Artesian Basin (BAB) with the TDS content between 0.4 and 1.0 g L−1 (Vaikmäe et al., 2008) changes to Cl–Na, Cl–Na–Ca and Cl–Ca– Na type groundwater with TDS contents ranging from 1 to 22 g L−1 in southern, deeper part of the Cm–V aquifer system (Karise, 1997). Deep brines in the central part of the BAB are typical sedimentary basinal Cl–Ca–Na fluids, with TDS of 60 to 140 g L−1 (Mokrik, 1997), enriched with respect to Ca2+ and depleted in Mg2+ and SO2− 4 compared to the evapo-/cryo-concentration trend of the modern seawater (e.g. Lowenstein et al., 2003). Also, Karro et al. (2004) suggested that some sulphate in the Cm–V aquifer system could have come to the system from deeper, saltier parts of the fractured crystalline rock underlying the aquifer rocks. Using the mass balance of geochemically conservative tracer components (Cl− and δ18Owater), Raidla et al. (2009) showed that the groundwater from the Cm–V aquifer system is a mixture of three major endmember sources: (a) the basinal brine, (b) diluted (meteoric) water, and (c) glacial melt-water. The results of mixing calculations (Raidla et al., 2009) suggest that in the shallow waters at the northern margin of the aquifer system the glacial component comprises N80% of the water. The content of the glacial component decreases and, accordingly, the brine component increases gradually with depth, and at ~700 m the glacial component has decreased to about 30 to 40% (Raidla et al., 2009). The brine component is enriched with respect to sulphate and in the basin's deeper southern part, in Latvia, sulphate concentrations are up to 1 to 2 g L− 1 (Kondratas, 1967). According to Raidla et al. (2009), the relict component composes a maximum of 1 to 2% of the groundwater from the Cm–V aquifer system in northern Estonia, which, assuming that the sulphate content in the original brine was 2 g L−1, accounts for approximately 20 to 40 mg L−1 in glacially diluted water that is comparable to the maximum sulphate values in the area (Table 1). Moreover, δ34S composition of the groundwater from the Cm–V aquifer system averaging at 24‰ seems to support the idea that the sulphate originates from relict seawater. Modern marine sulphate has the δ34S content of + 20.1‰, but it has varied during the Phanerozoic from about 30‰ in the Cambrian to about 10‰ at the beginning of the Triassic (Bottrell and Newton, 2006). However, the sulphate in a marine reservoir has a homogeneous and well-defined δ18OSO4 isotopic composition of about 9.5‰ (Longinelli, 1989), while the δ18OSO4 in the studied samples varies largely from

V. Raidla et al. / Chemical Geology 383 (2014) 147–154

−10.8 to 9.7‰ (average −0.2‰). This suggests that the sulphate found in the groundwater of the Cm–V aquifer system in northern Estonia originates from neither dissolution of marine sulphate minerals nor relict basinal brine. Interpretation of sulphate's oxygen isotope composition is difficult because the oxygen incorporated in the dissolved sulphate can originate from both atmospheric oxygen and/or a water molecule. Sulphate's oxygen isotope composition can be derived from and/or altered through several abiotic and biochemical processes, including disproportionation of elemental sulphur (e.g. Van Stempvoort and Krouse, 1994; Balci et al., 2012), oxidation of reduced compounds (pyrite, sphalerite etc.) (e.g. Balci et al., 2007; Heidel et al., 2013; Müller et al., 2013) and dissimilatory sulphate reduction (e.g. Böttcher et al., 2005; Brunner et al., 2005, 2012; Heidel and Tichomirowa, 2011). On the other hand, once formed, δ18OSO4 is a conservative environmental tracer as long as no biochemical processes such as bacterial sulphate reduction affect the sulphate pool. At ambient environmental temperatures the oxygen isotope exchange between sulphate and water is limited and occurs at a significant rate only in geothermal waters and/or in very low pH conditions (Krouse and Mayer, 2000; Tichomirowa and Junghans, 2009). However, Halas et al. (1993) show that equilibrium between water molecules and residual sulphate in groundwater is possible at low temperatures and in near-neutral pH (T = 9.5 °C and pH = 8) conditions at long residence times (~ 27,000 yr). This would result in covariance between δ18OSO4 and δ18Owater values. Groundwater's comparatively long residence time (N14,000 yr) in the Cm–V aquifer system would imply that partial oxygen isotope exchange might have occurred. Nevertheless, the measured δ18OSO4 and δ18Owater values do not show a strong positive trend (the correlation coefficient is 0.65; R2 = 0.42; Fig. 2) and isotope exchange at low temperatures has not yet been proven in the Cm–V aquifer system in Estonia. Raidla et al. (2012) put forward a hypothesis that the sulphate in the Cm–V aquifer system could originate from the oxidation of sulphide minerals that occurred during the intrusion of glacial meltwater. Most of the supraglacial waters appear to be undersaturated (c. 30 to 90%) with respect to atmospheric O2 (Brown et al., 1994). Normally, during infiltration to the glacier's basal part dissolved oxygen content in the melt-water decreases and reaches anoxic conditions. Still, basal melting of the glacier may receive a subglacial supply of O2 dependent on the composition and concentration of gas bubbles in the basal ice (Tranter et al., 2002). In addition, subglacial waters may be heavily enriched in dissolved gases, in case of O2 up to 50 times higher than airequilibrated water, through refreezing (McKay et al., 2003). This could mean that the oxidative meltwater penetrating into the aquifer system provoked the initial oxidation of sulphide minerals that provided most found in the groundwater from the Cm–V aquifer system. of the SO2− 4 Sulphite minerals (mainly pyrite, but occasionally sphalerite and galenite) occur in both Ediacaran and Cambrian sediments (e.g. Raidla

15

R² = 0.42

5 0

18

δ OSO4 (‰)

10

-5

et al., 2006) and in Proterozoic metamorphosed crystalline rocks confining the aquifer system from below (Petersell et al., 1991). Balci et al. (2007) show that water is the main oxygen source of sulphate during anaerobic and aerobic oxidation of pyrite, whereas abiotic/aerated experiments of pyrite oxidation at different pH levels show that 20 to 30% of the sulphate's oxygen comes from atmospheric O2 (Kohl and Bao, 2011). Heidel et al. (2013) have experimentally shown that during oxidation of sulphide mixtures the molecular oxygen is the main oxygen source of sulphate during the initial stage of oxidation, while the amount of water-derived oxygen in the sulphate increases N70% at later stages. The groundwater of the Cm–V aquifer system δ18OSO4 values vary from − 10.8 to 9.7‰, while its δ18O values stay rather uniformly at −18.5 to − 23‰ (Raidla et al., 2009), and the δ18O composition of O2 in the air is 23.5‰ SMOW (Kroopnick and Craig, 1972). During abiotic oxidation of sulphides (e.g. pyrite) the 18O is preferentially taken up into sulphate and the reported positive isotope offsets between sulphate and water in the absence of molecular O2 vary around 0 to 8‰ (e.g. Lloyd, 1968; Taylor et al., 1984; Van Everdingen and Krouse, 1985; Balci et al., 2007, 2012; Heidel et al., 2011). On average, the measured isotopic values of the sulphate oxygen in the groundwater from the Cm–V aquifer system are at least 5 to 10‰ higher than they would be, provided that the water was the only source of oxygen. However, in oxidation experiments using mixtures containing pyrite, galena and sphalerite, Heidel et al. (2013) showed high ε18OSO4–H2O values reaching 16.1 to 18.9‰, whereas the δ18OSO4–H2O values in their experiments depended on the initial δ18Owater values. These numbers were significantly higher (by 6.9 to 9.1‰) in waters with the initial δ18O value of − 17.4‰ compared to waters with the initial δ18O value of 8.0‰. An isotopic effect of this amplitude (14 to 17.6‰) is tied to the sulphite to sulphate oxidation step, which is paramount to the sulphate oxygen's isotopic composition (Müller et al., 2013). Oxygen isotope offset of this range and magnitude could explain the measured values and the variation in the δ18OSO4 values of the groundwater from the Cm–V aquifer system. Variations in the δ18OSO4 values could possibly be explained by differences in amounts of dissolved oxygen and/or ratios between different sulphide minerals. Alternatively, the oxidation of sulphides in abiogenic and anaerobic environments could be driven by Fe3+ as an oxidant (e.g. Heidel and Tichomirowa, 2011). However, the isotopic effect in this case is significantly lower (ε18OSO4–H2O = 2.3‰, Heidel and Tichomirowa, 2011) and could not explain the observed δ18OSO4 values in the groundwater of the Cm–V aquifer system. Earlier studies (e.g. Taylor et al., 1984; Van Stempvoort and Krouse, 1994; Brunner et al., 2005) suggest that the enriched δ18OSO4 values compared to those in the initial sulphate oxygen pool indicate that biological reactions occurred during the S oxidation process. However, Balci et al. (2007) show that δ18OSO4 values cannot be used conclusively to distinguish between biological and abiotic oxidation of pyrite. Moreover, a percentage of the sulphate that is brought into the bacterial cells during sulphate reduction is not reduced to sulphide completely. Instead, it undergoes isotope exchange between oxygen and water, reoxidates to sulphate, and is released back to the ambient sulphate pool (Mizutani and Rafter, 1973; Fritz et al., 1989; Brunner et al., 2005, 2012; Knöller et al., 2006; Mangalo et al., 2007, 2008; Wortmann et al., 2007; Farquhar et al., 2008; Turchyn et al., 2010). Nevertheless, explanations of this mechanism and conditions that control the process have remained speculative (e.g. Antler et al., 2013). 4.3. Sulphate δ34S

-10 -15 -25

151

-23

-21

-19

-17

-15

δ Owater (‰) 18

Fig. 2. δ18Owater and δ18OSO4 variation in the Cambrian–Vendian aquifer system.

In Estonia, the δ34S values of sulphides scatter rather widely between −30 and 22‰ in the rocks of the Cm–V aquifer system (Petersell et al., 1991). δ34S values of the rocks correspond to the Poisson distribution, the most frequent values ranging from 0 to −5‰ (Fig. 3). At the same time δ34S values of dissolved sulphate in the groundwater of the Cm–

152

V. Raidla et al. / Chemical Geology 383 (2014) 147–154

60

25

50

δ34SSO4 (‰)

frequency

20 15 10 5

40 30 20 10 0

0 20

10

0

-10

-20

-30

-40

0

10

20

30

40

SO42- (mg·L-1)

δ SSO4 (‰) 34

Fig. 3. The distribution of δ34S values of sulphides in the rocks of the Cambrian–Vendian aquifer system.

Fig. 5. δ34S and SO2− correlation in the Cambrian–Vendian aquifer system. 4

V aquifer system vary between 8.6‰ and 55.1‰ (average 24‰), suggesting that the groundwater's sulphate isotope composition generally corresponds to the sulphides' δ34S values, but on average is more enriched compared to sulphide minerals. During abiotic and/or biotic oxidation of sulphide mineral oxidation at low temperatures only a small isotope fractionation of δ34S occurs (close to 2‰) (Pisapia et al., 2007; Heidel et al., 2009). Consequently, the production of sulphate, via oxidation of sulphide minerals, yields δ34S values of sulphate that are similar to those of the source sulphides. However, bacterial reworking of the original sulphate has a specific signature expressed in the enriched δ34S values of dissolved sulphate (Strebel et al., 1990; Aharon and Fu, 2000; Spence et al., 2001). δ13C values of DIC varying from − 10.8 to − 24.3‰ also indicate bacteria's influence in groundwater from the Cm–V aquifer system (Raidla et al., 2012). In the studied samples sulphate is a minor anion and HCO3− the dominant anion with concentrations between 100 and 244 mg L− 1 (Table 1). There is a negative co-variance of sulphate and bicarbonate, whereas the most depleted δ13C values are mainly found in groundwater samples with the lowest sulphate concentrations (Fig. 4a), except in the abundant presence of isotopically depleted carbonate cement (δ13C = − 9 to − 13‰) (Petersell et al., 1991; Raidla et al., 2012). Depleted δ13C values of DIC in association with low dissolved sulphate concentrations suggest the biological origin of dissolved carbon, which is derived from anaerobic microbial oxidation of organic materials by sulphate reducing bacteria using SO24 − as the terminal electron acceptor (e.g. Jørgensen and Postgate, 1982; Wadham et al., 2004). During sulphate reduction significant isotope discrimination of sulphur isotopes in the residual sulphate modified by bacterial dissimilatory sulphate reduction has been observed both in experiments (Lloyd, 1968; Mizutani and Rafter, 1973; Fritz et al., 1989) and in natural

environments (Zak et al., 1980; Böttcher et al., 1998; Ku et al., 1999; Aharon and Fu, 2000). The enrichment of dissolved sulphate δ34S alongside the depletion of δ13C in the groundwater of the Cm–V aquifer system (Fig. 4b) provides strong evidence of oxidation of organic material through sulphate reduction by bacterial activity. Consequently, sulphate reduction leads to a simultaneous decrease in SO24 − content and sulphur's isotopic enrichment. Nevertheless, no clear trend derives from our data (Fig. 5), although the most positive δ34S values are found in samples with the lowest sulphate concentration. We speculate that this may be caused by the secondary sulphide oxidation. In addition to oxygen, pyrite may be oxidized by NO3− or Fe3+ (e.g. Bottrell and Tranter, 2002). Anoxic conditions prevail in the Cm–V aquifer system, where nitrogen oxides are missing and iron is represented in the form of Fe2+. Alternatively, Lefticariu et al. (2010) showed that radiolysis of water caused by radiation dose coupled to oxidation of sulphides could be a source of sulphate in geological environments. Indeed, the δ18OSO4 values are more similar to δ18Owater values in the Cm–V aquifer system in the areas, where the highest radon and radium contents were found. However, the effect of this phenomenon is unclear and deserves further study.

-35

5. Conclusions In the Cm–V aquifer system, northern Estonia, oxygen isotope composition of dissolved sulphate suggests that the sulphate was derived from oxidation of sulphide minerals when isotopically depleted water (δ18Owater = −18 to −23‰) and molecular oxygen contained in glacial meltwater, intruded into the aquifer system during the Last Glacial Maximum. Dissolved sulphate δ34S values are enriched compared with δ34S values of sedimentary-diagenetic and possibly hydrothermal sulphides

a

b

-30 -25 -20 -15

ity

tiv

-10

carbonate cement 13 (δ C -10‰)

-5 0 0

10

ia

ter

c ba 20

SO42- (mg·L-1)

30

40

0

10

c la

20

30

40

50

60

δ SSO4 (‰) 34

34 13 Fig. 4. δ13C correlation with SO2− 4 (a) and δ S (b) in the Cambrian–Vendian aquifer system. Isotopically depleted carbonate cement in rocks (δ C = −9 to −13‰) is marked with dotted line.

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in the aquifer system's rocks. Additionally, the enrichment of sulphate δ34S is accompanied by depleted δ13C values of the dissolved inorganic carbon in the groundwater of the Cm–V aquifer system. These phenomena can be explained by bacterial sulphate reduction that causes δ34S isotope fractionation in the residual sulphate.

Acknowledgements The activities of the present study were supported by MOBILITAS Postdoctoral Research Grant 2009 MJD17 (A.M.), Estonian Science Foundation Grants ETF9196 and IUT20-34 (K.K.) and ETF8948 (R.V.) and Doctoral School of Earth Sciences and Ecology (V.R.). The paper is a contribution to IUT19-22 and to the INQUA/UNESCO supported G@ GPS Project. The manuscript was improved by constructive comments from Prof. Simon Bottrell and Prof. Stanislaw Halas. We thank Dr. Stephan M. Weise for δ34SSO4 and δ18OSO4 analyses and Mrs. Helle Pohl-Raidla for the English revision.

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