A stable isotope study of microbial dolomite formation in the Coorong Region, South Australia

Chemical Geology 244 (2007) 155 – 174 www.elsevier.com/locate/chemgeo

A stable isotope study of microbial dolomite formation in the Coorong Region, South Australia David Wacey a,⁎, David T. Wright b , Adrian J. Boyce c a

Department of Earth Sciences, University of Oxford, Parks Road, Oxford OX1 3PR, UK Department of Geology, University of Leicester, University Road, Leicester LE1 7RH, UK Scottish Universities Environmental Research Centre, Rankine Avenue, East Kilbride G75 0QF, UK b

c

Received 9 October 2006; received in revised form 11 June 2007; accepted 11 June 2007 Editor: D. Rickard

Abstract The Coorong Region of South Australia is a classic location for the study of modern dolomite formation. A number of ephemeral lakes here contain dolomite, whilst other lakes in close proximity, contain aragonite with minor magnesite. Our previous field and experimental work has revealed that sulphate-reducing bacteria play a key role in overcoming kinetic inhibitors to dolomite formation in this area. We here use a combination of oxygen, carbon and sulphur isotopes to help to further constrain the processes leading to dolomite precipitation in modern hypersaline environments. We report isotopic data from four dolomite and three aragonite containing lakes. δ34S data for dissolved lake- and porewater sulphates reveal 34S enrichment of residual sulphate in all seven lakes, consistent with bulk closed system microbial fractionation by sulphate-reducing bacteria. Calculations using temperature-dependent oxygen isotope fractionation factors between dolomite and the solution from which it precipitated are consistent with a strong microbial component to carbonate precipitation in these lakes. A comparison of δ18O data from dolomite and aragonite sediment and from their associated lakewaters indicates seasonal mineral precipitation occurs late in the evaporative cycle of the lakes within a yoghurt-like saturated mud horizon coincident with the zone of sulphate reduction. Surprisingly, δ13C data from the carbonates do not show the characteristic wide range of values often associated with ‘organogenic’ carbonates, but instead show that these bacterially-mediated precipitates incorporate carbon primarily from the inorganic, lakewater reservoir, partially diluted by an organic component. A ‘Coorong type’ microbial model of dolomite formation may be particularly applicable to, for example, Precambrian sedimentary dolomites where benthic microbial communities, which are today restricted to extreme environments, dominated the shallow-marine ecosystem prior to the emergence of Metazoa. © 2007 Elsevier B.V. All rights reserved. Keywords: Coorong; Dolomite; Stable isotopes; Sulphate-reducing bacteria

⁎ Corresponding author. Department of Earth Sciences, University of Oxford, Parks Road, Oxford, OX1 3PR, UK. Tel.: +44 1865 272059; fax: +44 1865 272072. E-mail addresses: [email protected] (D. Wacey), [email protected] (D.T. Wright), [email protected] (A.J. Boyce). 0009-2541/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2007.06.032

1. Introduction The mineral dolomite, CaMg(CO3)2 is a common constituent of the rock record, yet a comprehensive understanding of its mechanisms of formation has yet to be achieved. The so called ‘dolomite problem’ has

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resulted from the inability of researchers to precipitate dolomite inorganically in the laboratory under earth surface temperatures and pressures, together with the enigma of the relative scarcity of modern occurrences of the mineral despite seawater being supersaturated with respect to dolomite (Garrels and Thompson, 1962). Thermodynamic considerations predict that dolomite in the marine environment should precipitate preferentially from normal seawater before other carbonate phases, and that all marine aragonite and calcite should be immediately dolomitized (Lippmann, 1973). The fact that this does not happen is attributed to three main kinetic constraints on dolomite formation: (1) the high hydration energy of the Mg2+ ion (Lippmann, 1973); (2) the low concentration and activity of CO32− (Garrels and Thompson, 1962); (3) the presence of sulphate which forms neutral ion pairs with magnesium, thereby significantly increasing its solubility (Baker and Kastner, 1981). We believe a major insight into the sedimentary dolomite problem can be obtained through detailed study of those environments where the mineral is forming at the present day under earth surface conditions. Thus our study focuses on the well-known “Coorong dolomite”, which is forming today in the ephemeral lakes of the Coorong region of South Australia. 1.1. Previous work in the Coorong The Coorong area of South Australia is an ideal location for the study of modern dolomite formation, and a relatively large literature has been built up as a result of research over the last 50 years (Alderman and Skinner, 1957; Alderman and Von der Borch, 1961, 1963; Skinner, 1963; Skinner et al., 1963; Von der Borch et al., 1964; Von der Borch, 1965a,b; Alderman, 1965; Von der Borch et al., 1975; Von der Borch, 1976; Von der Borch and Jones, 1976; Von der Borch and Lock, 1979; Lock, 1982; Botz and Von der Borch, 1984; Rosen et al., 1988; Warren, 1988; Ahmad and Hostetler, 1988; Rosen et al., 1989; Warren, 1990; Ahmad and Hostetler, 1994; Wright, 1999; Wacey, 2002; Wright and Wacey, 2004, 2005). These studies have resulted in a number of observations and theories: the dolomite found in several ephemeral lakes in the Coorong Region is fine grained (b1 μm to ∼ 20 μm) and often spherulitic when young, ageing to more rhombic forms (Von der Borch and Jones, 1976). It forms layers of structureless aphanitic micrite, often with desiccation cracks and tepees (e.g., Von der Borch and Lock, 1979). Radiocarbon studies show the age of surface deposits to be 300 ± 250 years,

increasing to 2030 ± 250 years at 33 cm depth and the average rate of deposition is ∼ 0.2 mm per year (Skinner et al., 1963; Von der Borch et al., 1964). Skinner (1963) and Von der Borch (1965a), amongst others, argued that the dolomite is primary in origin, showing no evidence for replacement of a pre-existing carbonate phase. Rosen et al. (1988, 1989), reported two ‘types’ of dolomite present, characterised by different unit cell parameters and carbon and oxygen isotopic values (see Section 1.3). It has been suggested that dolomite precipitation is favoured by elevated pH, as a result of plant growth (Alderman and Skinner, 1957; Skinner , 1963; Von der Borch et al., 1964; Alderman, 1965), and by elevated Mg/Ca ratios (e.g., Alderman, 1965). Comprehensive sampling of lake- and porewaters (Wacey, 2002) has shown that contrary to previous speculations on relationships between lakewater composition and the timing of separation of lakes from the main lagoon (e.g., Alderman, 1965), the Mg:Ca ratios and chemical compositions of the different lakes during evaporation are varied and dynamic, and appear to be related to a complex combination of evaporation, microbial mediation, organic diagenesis and mineral precipitation. In the oldest lakes (e.g. Lake McFaiden), sediment cores record the evolution of the lake from a lagoon with seawater connection (aragonite and calcite) to partially restricted lake (Mg-calcite and proto-dolomite), to isolated ephemeral lake (ordered dolomite and minor magnesite) (Von der Borch et al., 1975). It has been proposed that dolomite formation in the Coorong was associated with groundwater discharge from carbonate aquifers followed by evaporation (Von der Borch et al., 1975; Von der Borch, 1976; Von der Borch and Lock, 1979; Rosen et al., 1988, 1989; Warren, 1990) with magnesium ions supplied either locally from Mg-calcite limestone (Ahmad and Hostetler, 1988) or from an inland volcanic province, transported into the discharge area by the shallow groundwaters (Von der Borch et al., 1975). These hydrological models for the Coorong are a useful starting point but they fail to explain how the kinetic inhibitors to dolomite formation are overcome, and why evaporation of the ephemeral Coorong lakewaters in the laboratory result not in dolomite but in aragonite precipitation. Nor do they explain why other localities around the world with similar hydrological regimes fail to precipitate dolomite. Von der Borch (1976) suggested that, after emergence and evaporation, groundwaters are subjected to “as yet undefined conditions which result in dolomite formation” — we suggest that these ‘conditions’ are microbially influenced.

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1.2. Microbial dolomite The influence of microbes on dolomite formation has been highlighted recently by various authors (Kelts and McKenzie, 1982; Baker and Burns, 1985; Compton, 1988; Slaughter and Hill, 1991; Mazzullo et al., 1995; Vasconcelos et al., 1995; Vasconcelos and McKenzie, 1997; Wright, 1997, 1999; Gournay et al., 1999; Wright, 2000; Burns et al., 2000; Teal et al., 2000; Warthmann et al., 2000; Wacey, 2002; van Lith et al., 2003; Wright and Wacey, 2004, 2005). Specifically, a link between sulphate-reducing bacteria (SRB) and dolomite formation has been observed in e.g. anoxic, hypersaline coastal lagoons such as Lagoa Vermelha, Brazil (Vasconcelos et al., 1995; Vasconcelos and McKenzie, 1997), in deep sea organic rich sediments (e.g., Kelts and McKenzie, 1982; Baker and Burns, 1985; Compton, 1988; Slaughter and Hill, 1991), and in shallow-marine mudbanks such as the Cangrejo Shoals, Belize (Teal et al., 2000). Microbially-mediated dolomite formation has also been invoked for ancient dolomites, such as the Cambrian Eilean Dubh Formation of north-western Scotland (Wright, 1997), where increased amounts of dolomite correlate with degradation of oolitic organic material and microbial mats; the Pennsylvanian Paradox Formation (Gournay et al., 1999); and the Miocene Drakes Bay and Monterey Formations in California (Burns et al., 1988; Compton, 1988). Furthermore, we believe it is no coincidence that the world's most concentrated regional zinc anomaly – the Lower Carboniferous carbonate-hosted, Irish-type base-metal deposits (Singer, 1995) – is intimately associated with both microbial mud-mound build-ups (the Waulsortian Mudbanks, Lees and Miller, 1995) and extensive dolomitization, much of it undoubtedly early (Gregg et al., 2001). This zinc anomaly is present solely through the availability of an extensive, open-system Lower Carboniferous bacteriogenic sulphide reservoir: at least 90% of the giant Navan deposit has been precipitated with bacteriogenic sulphide (N 2.5 Mt of S at ore grade; Fallick et al., 2001). The critical fluids operating during this Lower Carboniferous mineralizing event are thought to be derived exclusively from contemporaneous evaporated Lower Carboniferous seawater (Banks et al., 2002). In the dolomitic Coorong lakes, our field data and laboratory experiments indicate that SRB are critical to dolomite genesis (Wacey, 2002, Wright and Wacey, 2004, 2005; Wright and Oren, 2005). Benthic microbial communities dominate these lake ecosystems during the latter stages of the evaporative cycle, when Metazoa are excluded due to high salinity, and their metabolic

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activities significantly affect the surrounding microenvironment, overcoming the common kinetic inhibitors to dolomite precipitation by raising pH, carbonate ion and magnesium ion concentrations, and by reducing sulphate ion concentrations (Wright, 1999; Wacey, 2002; Wright and Wacey, 2004). Experiments performed using natural populations of SRB cultured from Coorong lake sediments have revealed precipitation of nano-grains of dolomite on quartz and aragonite substrates in the laboratory, and have thus provided unequivocal evidence for the microbial mediation of dolomite precipitation in this environment (Wacey, 2002; Wright and Wacey, 2004, 2005). The microbes may also be actively involved in dolomite precipitation — the nano-grains of dolomite precipitated in the above experiments are spherical to elliptical in shape, and many are darker in the centre than around their edges (Wright and Wacey, 2004), suggesting that microbes acted as nuclei for dolomite precipitation. Other experimental studies (Folk, 1993a,b; Vasconcelos et al., 1995; Vasconcelos and McKenzie, 1997; Warthmann et al., 2000) have also shown nucleation of carbonates on bacteria, nano-bacteria, or extra-cellular polymeric substances. It has also been suggested (Beveridge, 1989; Folk, 1993a) that bacteria promote mineral precipitation by creating micro-environments conducive to nucleation due to cation attraction to their negatively charged cell walls. 1.3. Isotopic studies of Coorong dolomite There have been a number of previous studies using carbon and oxygen isotopes to attempt to gain a better insight into dolomite formation in the Coorong (Degens and Epstein, 1964; Clayton et al., 1968; Botz and Von der Borch, 1984; Rosen et al., 1988, 1989). Botz and Von der Borch (1984) presented the first comprehensive carbon and oxygen isotope data for the Coorong lakes. Using this, they concluded that two types of dolomite occurred. One was fine grained and relatively isotopically light (δ13C = − 1 to − 2‰; δ18O = +3 to +5‰) and the other was coarser grained, contained excess Mg and was isotopically heavier (δ13C = +3 to +4‰; δ18O = +5 to + 6‰). The “light” dolomite was inferred to be a primary precipitate from evaporatively modified continental water, whilst the “heavy” dolomite was inferred to form from waters that were close to equilibrium with atmospheric CO2, possibly by the dolomitization of aragonite. Rosen et al. (1988) then investigated Pellet Lake and showed that the upper dolomite unit had rather isotopically heavy oxygen (+ 7.55‰ average) and carbon (+ 4.1‰ average) isotope values. The heavy

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Fig. 1. Location of study lakes, Coorong Region, South Australia.

carbon values could be attributed to brine evaporation, or to organic processes such as methanogenesis; the results were thus equivocal, although a wider variation in δ13C might have been expected if organic processes were dominant (Rosen et al., 1988, but see Wright and Wacey, 2005 and this paper). The oxygen values are clearly controlled by evaporation, and three upwardlightening oxygen isotope cycles were identified by Rosen et al. (1988) which were interpreted as due to increased importance of rainwater dilution of the brine when the lake was at its shallowest. Rosen et al. (1989) extended this work to five Coorong lakes (although two, Halite and North Stromatolite, do not precipitate dolomite today) and differentiated two types of dolomite based on their isotopic signatures, unit cell dimensions, TEM signa-

tures and bulk composition. Type A, a Mg-rich dolomite, was seen to occur in the centre of the larger lakes, has heavy oxygen (+ 7.6‰ average) and carbon (+ 3.5‰ average) isotope signals and was interpreted to precipitate rapidly from evaporating brines with slightly elevated Mg/Ca ratios. In contrast, Type B dolomite has lower oxygen (+ 6.4‰ average) and carbon (− 1.2‰ average) isotope values, is either Ca-rich or almost stoichiometric, and was interpreted to form from relatively less evaporated brines, with a contribution from light meteoric or soil CO2. Dolomitization of aragonite was not invoked by Rosen et al. for any of the dolomite. We here report additional oxygen and carbon isotope results from the uppermost 10 cm of carbonate sediments of seven Coorong lakes together with oxygen

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Table 1 Chemical analyses and SRB populations for each of the Coorong ephemeral lakes Lake

Ion (mg l− 1) Mg2+

Mini Dolomite Milne McFaiden Pellet Dolomite South Stromatolite North Stromatolite Halite Seawater

Ca2+

8719 554 8914 105 8923 264 5634 176 Lake had already dried up 15,453 177 12,848 166 12,938 226 1313 400

Mg/Ca

CO2− 3

SO2− 4

pH

SRB count ml− 1

Mineralogy

16 85 34 32 at this time 87 77 57 3.2

125 132 84 1184

17,939 19,665 24,304 9726

8.25 7.65 7.46 8.38

767 942 108 16.2

9670 28,311 26,955 2688

8.42 6.87 7.51 ∼7.8

3.74 × 106 9.22 × 105 1.72 × 106 8.63 × 105 × 3.23 × 105 4 × 104 1.06 × 103

Dolomite Dolomite Dolomite Dolomite Dolomite Aragonite Aragonite Aragonite

Seawater values are given for comparison.

isotope data from evaporating lakewater, and sulphur isotope data from lakewater sulphate. 2. Study area The Coorong lies approximately 200 km south east of Adelaide where a series of ephemeral, alkaline lakes occur a few hundred metres inland from the Coorong Lagoon itself, an offshoot of the Southern Ocean. The ephemeral lakes are situated in an interdune corridor that is c.1 m above sea level (Von der Borch, 1976), flanked on one side by Holocene dunes, and on the other by Pleistocene coast parallel dunes (Schwebel, 1983). The relative drop in sea level during the Holocene due to sediment accretion has resulted in the Coorong lakes evolving from a dominantly marine setting to that of distal ephemeral lakes (Von der Borch, 1976) (Fig. 1). Today, the lakes are inundated in winter by meteoric water accompanied by the elevation of the water table, and dry up in summer due to intense evaporation. Two major aquifers underlie the main zone of primary dolomite formation; an upper unconfined aquifer made up of late Tertiary and Quaternary bryozoan limestones and calcareous sands, and a deeper, confined aquifer comprised of Eocene siliceous sands (Floegel, 1972). These two units are separated by an aquiclude made up of clays and lignites (O'Driscoll, 1960). The upper aquifer is fed by local rainfall recharge and infiltration from freshwater swamps in the Mt Gambier region to the south east (O'Driscoll, 1960); the underlying confined aquifer originates in a widespread freshwater swamp recharge area just north of Mt Gambier and has been shown to leak through the aquiclude into the upper aquifer (Floegel, 1972). Groundwater approaches the coastal discharge zone as freshwater lenses; as they reach the coastal zone they override denser interstitial waters of oceanic and lagoonal origin (Von der Borch et al., 1975).

A diffuse and complex zone of mixing of continental and marine waters therefore results along the coast in the vicinity of the dolomitic and non-dolomitic lakes. The source of ions for dolomite formation has been much debated (e.g., Von der Borch et al., 1975; Von der Borch, 1976; Von der Borch and Lock, 1979). Calcium, magnesium and bicarbonate ions may be made available through dissolution of skeletal components from within the shallow aquifer (Von der Borch and Lock, 1979). Calcium levels are low throughout the lake cycle and relatively constant throughout the study period so that local leaching from the surrounding carbonate dunes appears to be negligible. A significant source of Mg ions may come from the leaching of ash layers that are intercalated with coastal-plain sediments over a wide area in the east of the province where much of the northwest flowing surface water originates (Von der Borch et al., 1975; Von der Borch, 1976). This ash has lost 60–80% of its original Ca2+, Mg2+ and Na+, being leached into groundwaters and contributed to runoff waters (Hutton, 1974). Eventually this will recharge the shallow aquifer that provides the groundwater for the dolomitic lakes. Within the lakes themselves, we have observed cyanobacteria, Ruppia grass and Samphire. These may all be a source of additional magnesium, as well as contributing to the organic carbon pool of the lakes during desiccation (Gebelein and Hoffman, 1973; Wright, 1999). The lakes range in size from around 300 m 2 (Dolomite Lake) to over 1 km2 (Milne Lake) and all are very shallow, rarely exceeding 1 m water depth. The lake bottoms comprise a decimetre scale layer of unconsolidated sediment of yoghurt-like consistency. The density of this layer leads to stable stratification of the lowermost part of the water column during late spring to early summer (Wacey, 2002). Mixing with the oxygenated waters of the upper water column is prevented by density stratification, leading to a largely

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(Wright, 1999; Wacey, 2002). The lakes are hypersaline, alkaline, and contain magnesium, carbonate and sulphate concentrations that are all highly elevated compared to seawater (Table 1). SRB are abundant in all lake sediments, with the highest numbers (up to 3.74 × 106/ml) in the dolomitic lakes (Table 1): dolomite lakes contain at least twice the SRB populations of the non-dolomitic lakes. The primary dolomite is fine grained (b 1 μm to ∼ 20 μm) and often spherulitic when young, ageing to more rhombic forms (Von der Borch and Jones, 1976; Wacey, 2002). As well as ordered dolomite, a Ca:Mg carbonate identical to dolomite but with weak ordering reflections, with stoichiometries ranging from Ca-rich through Ca50: Mg50 to Mg-rich is also present in several lakes. Despite intense evaporation no solid sulphates are found in these lakes; this has been attributed to seeping groundwaters and wind (e.g., Von der Borch, 1965b), and more recently to bacterial sulphate reduction (Wright, 1999, 2000; Wright and Wacey, 2004, 2005). 3. Methods Fig. 2. Evolution of δ18Olakewater during evaporation and desiccation of Coorong lakes.

anoxic environment at the sediment–water interface. In addition, the high concentration of suspended material in the water column prevents significant light penetration down to this layer. This combination of factors creates an environment that, during the latter stages of the evaporative cycle, is hostile to most plant and animal life. The establishment of anoxia at, or just below, the sediment–water interface is a key feature of these lakes allowing anaerobic bacteria, including SRB, to flourish at this time. These distal ephemeral lakes provide an excellent natural laboratory for the study of dolomite precipitation. In this extraordinary setting, lakes with a modern infill of dolomite or ‘proto-dolomite’ lie in close proximity to those completely lacking modern dolomite, all contained within the same hydrological and geological setting. We chose five lakes where dolomite was actively forming; Milne, Pellet, Mini Dolomite, McFaiden and Dolomite (the latter provided limited data because it was dry at the time of sampling), and compared them with three lakes where dolomite is not precipitating today; Halite, North Stromatolite and South Stromatolite. All lakes are characterised by dynamically changing chemistries with large shifts in ionic concentrations over time-spans of less than a week

3.1. Field and laboratory methods Sediment and water sampling took place in early summer (late October to early December) in two successive years and was timed to coincide with the latter stages of lake desiccation when physical, biological and chemical changes in the lakes will be greatest. It was known (Wright, 1999; Wacey, 2002) that this was the stage in the annual cycle where sulphate reduction became the dominant biogeochemical process and microbial dolomite precipitation would most readily occur. Porewater samples were extracted from 1 cm core intervals by high speed centrifugation. Carbonate sediment was collected using

Fig. 3. δ18Ocarbonate and δ13Ccarbonate from sediments within the Coorong lakes. Note that the dolomitic and non-dolomitic lakes plot in distinctly different fields.

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±0.3‰. All oxygen isotope compositions are reported in standard δ notation relative to V-SMOW unless otherwise stated. 3.2.2. Oxygen isotopes from carbonate sediment and lakewaters Oxygen isotope analyses of carbonates and lakewaters were prepared and analyzed using apparatus from Analytical Precision (AP2003) consisting of a fully automated sample preparation module and a sample processing module, interfaced to a dedicated Isotope Ratio Mass Spectrometer. The internal Mab2b standard was used (δ18O = − 2.40‰ V-PDB; 28.4‰ VSMOW) with a reproducibility of ± 0.2‰. 3.2.3. Carbon isotopes from carbonate sediment Carbon isotope analyses of carbonates were prepared and analyzed using the same apparatus as for oxygen above. The internal Mab2b standard was again used (δ13C = +2.48‰ V-PDB) with a reproducibility of ±0.1‰. 3.2.4. Sulphur isotopes from residual sulphate BaSO4 samples were processed following the procedure of Coleman and Moore (1978), with evolved Fig. 4. δ34Ssulphate-residual from each of the Coorong study lakes during the month of November. Note seawater value of 22‰ for comparison. All lakes show depletion of the light 32S isotope in the residual sulphate consistent with fractionation by sulphate reducing bacteria.

5 cm3 open-ended syringes. Samples were fixed with glutaraldehyde to prevent further microbial activity, sealed with butyl rubber stoppers to prevent oxidation and refrigerated immediately. Sulphate for use in sulphur and oxygen isotopic analysis was precipitated directly from lakeand porewater samples using barium chloride (Vogel, 1978). SRB population sizes were assessed using most probable number statistical dilutions (Hurley and Roscoe, 1983) after adding 1 cm3 of sediment to a Postgate B culture media (Postgate, 1984). The total numbers of bacteria (SRB + others) present in a sample were calculated using acridine orange direct counting (Kirchman et al., 1988; Fry, 1988, 1990). Total organic carbon was measured on a Strohlein Coulomat 702 machine using part of the same core used for SRB population measurements. 3.2. Isotopic analyses 3.2.1. Oxygen isotopes from residual sulphate Sulphate samples were processed following the procedure of Hall et al. (1991) with evolved CO2 analyzed on a VG SIRA 10 mass spectrometer. Reproducibility of the standard (NBS 127 with δ18O = 9.3‰ V-SMOW) was

Fig. 5. δ34Ssulphate-residual from both the lake- and porewaters of each of the Coorong lakes. Note seawater value for comparison (see Section 5.4 for discussion).

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Fig. 6. Relationship between δ34Ssulphate-residual and δ18Osulphate-residual in the Coorong lakes. A much stronger correlation is seen in the dolomitic lakes with an increase in δ34S being accompanied by an increase in δ18O. No correlation is observed in the non-dolomitic lakes (R2 = 0.0037).

SO2 isotopically analysed on a VG SIRA II mass spectrometer. Reproducibility based on repeat analyses of standards (CP1 = − 4.6‰; IAEA-S-3 = − 31‰; NBS 123 = 17.1‰ V-CDT) was better than ± 0.2‰. Precipitated BaSO4 samples showed similarly good internal consistency and reproducibility, with each lake having a small and characteristic range of values. No co-existing sulphides could be analysed as the sediments are iron poor and H2S gas escapes rapidly from the sediment before it can be mineralised. All sulphur isotope compositions are reported in standard δ notation relative to V-CDT.

4. Results 4.1. Oxygen isotopes from residual aqueous sulphate δ18Osulphate-residual in the dolomitic lakes ranges from 19.9 to 28.1‰ with a mean of 24.3‰, whilst the range for the non-dolomitic lakes is from 21.9 to 26.7‰ with a mean of 23.9‰. These values deviate significantly from the open ocean value of 9.5‰ (Lloyd, 1967; Longinelli and Craig, 1967) and from our measurements from coastal seawater (average 10.8‰). No other sulphate oxygen data have been published for these lakes and

Fig. 7. Relationship between δ18Osulphate-residual and δ18Olakewater in Coorong dolomitic lakes. A strong positive correlation is observed indicating recycling of sulphate within a lake together with isotope exchange, proceeding via oxygen exchange between sulphate–enzyme complexes and ambient water.

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Fig. 8. Paths taken during evaporation of water under various conditions. (A) Evaporation of seawater under humid conditions and (B) Evaporation of seawater under arid conditions (Knauth and Beeunas, 1986); (C) Evaporation of seawater to ×10 concentration (Holser, 1979); (D) Evaporation of seawater from Pierre et al. (1984). The dashed line represents our best estimate of the evaporation path taken by water in Mini Dolomite Lake.

indeed very little δ18O sulphate data has been reported for any saline lake settings. Longinelli and Craig (1967) report enrichments of 7–23‰ in a variety of saline lakes, the highest being from Searles Lake (+ 23.16‰) where the sulphate concentration was 48,700 ppm, the same magnitude as observed in the Coorong lakes. 4.2. Oxygen isotopes from carbonates and lakewater δ18O values for lake sediments have a narrow range between 2.8‰ V-PDB (33.7‰ V-SMOW) and 5.4‰ VPDB (36.4‰ V-SMOW), lying at the higher end in the field of normal sedimentary rocks (Hoefs, 1987), similar to values reported by Botz and Von der Borch (1984), and a little lighter than those reported by Rosen et al. (1988, 1989) for Coorong dolomite. δ18O analyses for water taken from the study lakes at several times during the sampling season are presented in Fig. 2. δ18O values span a large range of 5.7 to 12.3‰ V-SMOW, these values being isotopically heavy when compared to seawater which by definition, on the V-SMOW scale, is approximately zero. 4.3. Carbon isotopes δ13C values for lake sediments are shown in Fig. 3 (plotted against δ 18 O) and range from − 1.19 to + 3.22‰ V-PDB for the dolomitic lakes, whereas the non-dolomitic lakes have a more closely spaced δ13 C range from + 1.12 to + 3.07‰ V-PDB. These values are very similar to those reported previously for these

lakes (Botz and Von der Borch, 1984; Rosen et al., 1988, 1989). 4.4. Sulphur isotopes δ34Ssulphate-residual data from the lakes are shown in Fig. 4. Each lake has a characteristic set of δ34Ssulphateresidual values that are tightly grouped; there is no significant variation between the start and end of the sampling season. All lakes show significant modification from the standard seawater source sulphate value of + 21‰ (Longinelli, 1989; Rees et al., 1978; we measured local, shallow seawater to be + 22‰ during sample collection, Fig. 4) with sulphate in each lake enriched in 34S. Milne Lake (dolomitic) shows the maximum 34S enrichment of over 8‰, whilst the minimum enrichment (b 2‰) is in non-dolomitic North Stromatolite Lake. However, there are no patterns in the δ34S data alone that enable a discrimination to be made between the dolomitic and non-dolomitic lakes. Furthermore, geographical location has no significant influence on these data. Only minor variation in δ34Ssulphate-residual versus depth is observed (Fig. 5), with each lake having a fairly compact and characteristic set of values, again always significantly in excess of seawater δ34Ssulphate. The porewaters of Halite Lake, McFaiden Lake and to a certain extent Milne Lake show a general trend of increase in δ34Ssulphate-residual with depth from 10 mm but only by relatively small amounts (from b 1 to 1.5‰; Fig. 5).

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5. Discussion 5.1. Oxygen isotopes from dissolved sulphate The equilibrium fractionation factor for isotopic exchange between oceanic water (δ18 O ∼ 0‰ VSMOW) and sulphate ions as determined by Mitzutani (1972) is 28.3‰ at 25 °C increasing to 33.1‰ at 5 °C. Yet sulphate ions in seawater have a relatively constant δ18O of only 9.5‰ (Lloyd, 1967; Longinelli and Craig, 1967). Hence, it is clear that the oxygen in the sulphate ion does not obtain isotopic equilibrium with the oxygen in seawater, and instead largely reflects the source of the sulphate.

The δ18O of sulphate ions is strongly influenced by bacterial sulphate reduction and by disproportionation and oxidation of reduced sulphur compounds. The exact mechanism by which residual sulphate is enriched in 18O during these reactions has recently been debated (Brunner et al., 2005). Brunner et al. (2005) reviewed both the traditionally accepted mechanism of kinetic isotope exchange (Rees, 1973) and the mechanism proposed by Fritz et al. (1989) that oxygen isotope exchange between intermediate sulphur compounds and the ambient water dominated δ18Osulphate values. Using detailed experimental work, Brunner et al. (2005) were able to independently verify the results of Fritz et al. (1989). They went on to show that the equilibrium oxygen isotope

Fig. 9. Graphs of temperature versus δ18Olakewater, showing the calculated range of precipitation temperatures of the carbonate minerals obtained from measured values of lakewater and mineral, assuming equilibrium. (a) Contoured in units of δ18O-dolomite using Eq. (1a); (b) Contoured in units of δ18O-dolomite using Eq. (1b); (c) Contoured in units of δ18O-‘proto-dolomite’ using Eq. (2); (d) Contoured in units of δ18O-aragonite using Eq. (3). Halite Lake is here taken as an example of how each graph is constructed: The surface sediment from Halite Lake has a measured δ18O value of 34.62‰ (this study) and so this is drawn as a contour on the diagram (thick purple line in (d)). The measured δ18O lakewater values from this study are between 5.7 and 7‰ and so the precipitation temperatures to which these two isotopic values correspond is read off the x-axis of the diagram (two thin purple lines in (d)).

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exchange reaction could better account for the observations of δ18Osulphate-residual and of δ18Osulphate-residual vs δ34Ssulphate-residual seen in natural settings. They do not discount a kinetic isotope effect but merely conclude that it is not essential to explain the data obtained so far. Using the model developed by Brunner et al. (2005), the extent of δ18Osulphate fractionation will be variable and largely dependent on the δ18OH2O of the ambient water. For example, in the experiments of Brunner et al. (2005) δ18Osulphate-residual rose by over 30‰ when the water had δ18O of + 85‰ but only by 2‰ when the water had δ18O of − 10‰. In natural marine sediments a maximum enrichment of 17‰ has been observed (Böttcher et al., 1999; Ku et al., 1999). The linear slope of the δ 18 Osulphate-residual vs 34 δ Ssulphate-residual relationship has often been used in the past to calculate δ 18 Osulphate enrichment. A slope ratio of 1:4 was calculated by Mitzutani and Rafter (1969) and this has been widely quoted in the literature. Hence, if one knew the sulphur isotope enrichment, the equivalent oxygen enrichment could be calculated. However, the experiments of Brunner et al. (2005) show that there is no standard ratio for δ18Osulphate-residual vs δ34Ssulphate-residual enrichment during sulphate reduction — it varies depending on the isotopic composition of the starting water. Nevertheless, the equilibrium oxygen isotope exchange reactions of Brunner et al. (2005) still result in a linear covariance of δ18Osulphate-residual and δ34Ssulphate-residual when sulphate concentrations are high (i.e. equilibrium has not been reached). For the Coorong lakes, where starting sulphate concentrations are extremely high, this indicates that the process of sulphate reduction would lead to covariance of sulphur and oxygen isotopes in the residual sulphate. Moreover, as Ku et al. (1999) and Moreira et al. (2004) point out, some of the sulphate reduced may be recycled within the lake- and porewaters via bacterial oxidation of upwardly diffusing H2S and disproportionation reactions, before being reduced once more. This means that an original seawater sulphate ion is more likely to inherit the δ18O value associated with the fluids in the area it is re-oxidised (i.e. the ambient lakewater). A scatter plot of δ18Osulphate-residual vs δ34Ssulphate-residual for the sulphate analysed in this study is shown in Fig. 6. The dolomite-producing lakes show a highly significant correlation (R2 = 0.815 at the 99% confidence limit), compared to the essentially random correlation in non-dolomitic lakes (R 2 = 0.004), with an increase in δ 34 S being accompanied by an increase in δ18 O. The greater correlation in the dolomitic lakes, combined with the much larger numbers of SRB present in the dolomitic lakes (for example Mini Dolo-

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mite Lake has 3500 times more SRB than Halite Lake, Table 1) suggests that bacterial sulphate reduction has a stronger influence on lakewater/porewater chemistry in the dolomitic lakes. Our data are also consistent with the oxygen isotope exchange model of Fritz et al. (1989) and Brunner et al. (2005). In the dolomite containing lakes the high δ 18 Osulphate-residual values (up to 28‰) correlate (R2 = 0.96) with high ambient lakewater δ18OH2O (up to 12‰) (Fig. 7), and the fractionation factors (mean ▵18Osulphate-water = 15.3‰) are well within those permitted by the oxygen isotope exchange model (▵18 Osulphate-water ∼ 25‰). Our mean ▵18Osulphate-water of 15.3‰ is not as great as the 25‰ calculated by Fritz et al. (1989) and this can easily be explained. Fritz et al. (1989) used a closed system and calculated their ▵ 18 Osulphate-water from experiments where b 2% of the original sulphate remained. In the case of the Coorong lakes, a higher percentage of sulphate remains, therefore the lakes are further away from equilibrium. In addition, the signal may be complicated by inputs from sea spray (δ 18 O ∼ 10‰) and by the fact that different species and concentrations of vital components exist in these lakes when compared with the Fritz et al. (1989) experiments. 5.2. Oxygen isotopes from the lakewaters and carbonates 5.2.1. Brine evolution The lakes are situated in a seasonal seawater– groundwater mixing zone and, in addition, are often refilled directly with meteoric water. During spring and summer, all of these lakes undergo intense evaporation, which concentrates 18O in the residual lakewater. This enrichment effect was substantial here because samples were collected near the end of the annual evaporative cycle of the lakes. The driest lakes, at the time of sampling (McFaiden and Mini Dolomite), show the greatest 18O enrichment up to the week ending 8th November (Fig. 2). Beyond this date, three of the lakes (Halite, Pellet and North Stromatolite) show increased 18 O enrichment with increased evaporation, as expected. The remaining four, however, (significantly the driest) show the reverse pattern, indicating that an opposing process is taking place. An input of meteoric water can be discounted as there was no substantial rainfall in this period and different lakes responded in different manners. Craig and Gordon (1965) modeled the process of δ18O enrichment of evaporating low salinity waters showing that δ18O increases until the final stages of drying when a

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steady state is reached due to the back transport of water molecules from the atmosphere. The steady state δ18O value in that situation depended upon relative humidity (Major et al., 1992) and the δ18O value of the atmospheric moisture in the area. A rather different situation is seen in water bodies of high salinity such as those in the Coorong. Laboratory experiments by Lloyd (1966), reinforced by observations in natural settings, show that the heavy isotope content of the residual brine increases at first but then reverses, without a steady isotope composition being reached. For example, brines in the Persian Gulf were shown to increase to δ18O = +5‰ at a salinity of 150‰ but then decrease slowly at higher salinities (Lloyd, 1966). Lloyd concluded that a maximum δ18O enrichment of about +6‰ could be reached by the evaporation of coastal seawater brines and that further enrichment would be retarded by back exchange of atmospheric moisture with the evaporating fluid. Knauth and Beeunas (1986) summarize the δ18O evolution of a residual seawater brine, extending its evolution to extreme ‘halite facies’ evaporation and beyond producing, when plotted against hydrogen isotopes (δD), a hooked trajectory (Fig. 8). As salinity increases, there is an increase in the relative numbers of water molecules tied up in hydration spheres around cations. Isotopic exchange between these molecules, unbound water, water leaving the liquid–air interface and atmospheric water is believed to cause the hooked effect (Holser, 1979; Holser et al., 1979; Knauth and Beeunas, 1986) with its exact shape depending upon climatic variables. Evaporating meteoric water will follow a path similar to that for seawater except that the starting point will be somewhere on the meteoric water line. Because the concentrations of dissolved constituents are much less in meteoric water, evaporation may have to proceed to greater concentrations and greater 18O enrichments before the hooked trajectory appears (Knauth and Beeunas, 1986). Thus, 18 O enrichments in excess of 30‰ relative to V-SMOW have been demonstrated for evaporating meteoric waters in arid environments (Fontes and Gonfiantini, 1967). δ18O enrichments in all of the Coorong lakes are greater than the 6‰ predicted by Lloyd as being the upper limit for the evaporation of seawater but less than the 30‰ demonstrated for purely meteoric water. This is entirely as expected because the lakes are fed not only by seawater but also by groundwater and meteoric water. In the absence of δD data, we can only estimate the path that the Coorong lakewaters may evolve along. However, the δ18O data for the four driest lakes is consistent with the pattern discussed above (an estimate of the path Mini Dolomite lakewater may take is superimposed on Fig. 8).

Clearly, observations over a longer timescale together with δD data are necessary to constrain the evolution of the water more tightly. Even then, similar isotopic data can be obtained from both evaporating marine and meteoric waters and so chemical data combined with isotopic data is necessary to fully understand the source and evolution of brines. Our data present a complex picture of potential interaction between a number of possible fluids and intense evaporation, nonetheless they indicate that each lake undergoes its own individual evaporation cycle, resulting in different maximum enrichments in 18O at slightly different times of the year. 5.2.2. Mineral–water interactions and timing of precipitation One factor not yet considered that could also result in the drop in δ18O-lakewater in the driest four lakes is mineral precipitation. As the lakes desiccate and saturation states are reached, minerals will precipitate so long as kinetic inhibitors are overcome (Wright, 1999; Wright and Wacey 2004, 2005); each mineral has an isotopic fractionation factor governing the way isotopes in the formation water will behave during precipitation. Water–mineral interactions result in the shift of δ18O values of the mineral and the water away from their initial values. If the oxygen reservoir in the water is much greater than that in the mineral/rock, the mineral/rock values will shift but the water values will remain essentially constant, e.g. seawater values remains constant when calcite precipitates because the seawater oxygen reservoir is so large. However, the converse situation may occur when the water reservoir is relatively small such as in the case of the Coorong lakes. Carbonate mineral phases in the driest four lakes (Milne, McFaiden, Mini Dolomite and South Stromatolite) as determined by X-ray diffraction are dominantly dolomite (with some ‘proto-dolomite’) and aragonite (e.g., Von der Borch, 1976; Wacey, 2002); commonly used mineral–water oxygen isotope fractionation equations for these minerals, and their sources, are as follows (T in Kelvin): 1000lnaðdolomite–waterÞ ¼ 3:20  106 T 2  1:50

ð1aÞ

(extrapolated from high temperature studies; Northrop and Clayton, 1966; Friedman and O'Neil, 1977) 1000lnaðdolomite–waterÞ ¼ 2:73  106 T 2 þ 0:26

ð1bÞ

(Until recently it had been impossible to experimentally calibrate the temperature-dependent oxygen isotope

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fractionation factor between dolomite and the solution it precipitated from at low temperatures due to the inability to experimentally precipitate dolomite at these low temperatures. Instead one had to extrapolate the fractionation factor from higher temperature studies (using Eq. (1a)). However, Vasconcelos et al. (2005) published a low temperature oxygen isotope fractionation equation (Eq. (1b)), calculated directly from microbially catalysed, laboratory precipitated low temperature dolomite.) 1000lnaðdprotodolomiteT–waterÞ ¼ 2:78  106 T 2 þ 0:11

ð2Þ

(Fritz and Smith, 1970) 1000lnaðaragonite–waterÞ ¼ 18:45  103 T 1  32:54

ð3Þ

(Bohm et al., 2000) This gives enrichment in the mineral phase of 18O with respect to the formation water, when in equilibrium at 25 °C (298°K) with that water, as follows: (1a) Dolomite = 34.5‰ (1b) Dolomite = 31‰ (2) ‘Proto-dolomite’ = 31.4‰ (3) Aragonite = 29.4‰ Such large isotope fractionations in favour of the heavy isotope entering the mineral might easily account for associated decreases in δ18O observed in the residual pools of water in the lakes. However, in the Coorong lakes, mineral precipitation does not appear to be the mechanism by which the δ18O of the fluid is lowered. If this was so, the minerals would have a much heavier δ18O (∼ 40 to 45‰ V-SMOW), reflecting precipitation from a brine with δ18O = 10 to 12‰ V-SMOW. Instead it appears that the drop in δ18O-lakewater in the four driest lakes is indeed due to the ‘hook effect’ from intense evaporation (Fig. 8; Knauth and Beeunas, 1986). We can then take our observations one step further. Using our δ18O-lakewater and δ18O-carbonate data (Figs. 2 and 3), the dynamically changing δ18O of the lakewater may be used to investigate the time and location of carbonate mineral precipitation. As temperature increases, the isotopic fractionation factor between the mineral and the water from which it precipitates decreases, and can be calculated precisely, assuming equilibrium. The average ambient temperature of the lakes is also known (20–30 °C; Wacey, 2002), so using Eqs. (1a), (1b), (2) and (3) it is possible to deduce whether the sediment precipitated in equilibrium with

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the measured lakewaters. Fig. 9a–d plot fluid δ18O versus temperature and are contoured in units of δ18Osediment for dolomite (using Eq. (1a)), dolomite (using Eq. (1b)), ‘proto-dolomite’ (using Eq. (2)) and aragonite (using Eq. (3)) respectively. Comparing Fig. 9a with Fig. 9b, it can be seen that Eq. (1b) gives temperatures of precipitation closer to ambient temperatures for the dolomite in the Coorong lakes than does Eq. (1a). Using the equation formulated by extrapolation from high temperature studies (Eq. (1a)), temperatures ranging from 53 to 92 °C are necessary for the dolomite to precipitate in equilibrium with these measured lakewaters. However, using the low temperature microbial equation (Eq. (1b)), equilibrium precipitation temperatures are as low as 37 °C in Mini Dolomite Lake and around 45 °C in the other three lakes, suggesting that microbial mediation is a more realistic mechanism. Using the equation based on precipitating less ordered ‘proto-dolomite’ gives equilibrium temperatures (39–79°C), close to those using Eq. (1b). The temperature of the lake- and porewaters was measured throughout the field season with lakewaters ranging from 20 °C up to 30 °C and porewaters as high as 35 °C. These temperatures are lower than the calculated precipitation temperatures of carbonate minerals (e.g. around 37 °C to 78 °C for dolomite lakes using Eq. (1b). For the dolomite to have precipitated at the ambient temperatures of the lakewaters (e.g., 20–30 °C), the δ 18 O range for the lakewaters would have to be much lower, between about 2 and 7‰ V-SMOW (arrowed values on Fig. 9b). The logical conclusion is that the dolomite must precipitate from the lakewaters at some other time in the evaporitic cycle of the lakes. Based on the oxygen data and our observations of the lakes, the preferred time of dolomite precipitation is a few weeks later in the evaporative cycle of the lakes. At this time, hyper-salinity is well established, killing off any remaining macrofauna that may graze bacterial communities. The organic components of the lake (Ruppia and Samphire grass) were observed to be dying (as salinity increased) and starting to degrade during our sampling season. Complete death and degradation of this organic matter in the following weeks would provide large amounts of magnesium for incorporation into dolomite, and metabolism of the decaying organic matter by SRB would further increase the pH of the lake- and porewaters. The evaporative ‘hook effect’ would lower δ18O values in the lakewater, and temperatures may also increase by a degree or two, so that the ‘low temperature oxygen isotope fractionation equation (Eq. (1b))’ between dolomite and

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water is then satisfied. We therefore suggest that dolomite precipitation is seasonal and occurs around December each year in the larger lakes (although local climatic extremes may cause some small variation) but a little earlier (as early as October) in the smaller lakes, such as Dolomite Lake, which dry up more rapidly. In the absence of δ18O values for the porewater, we may only speculate as to whether dolomite may be precipitating within the sediment column. The dolomite is a primary precipitate (Skinner, 1963; Von der Borch, 1965a; Von der Borch and Jones, 1976; Wright, 1999; Wacey, 2002) and it is entirely plausible that it can precipitate in the waterlogged yoghurt-like zone just below the sediment–water interface. The micro-spherular nature of the youngest dolomite (Von der Borch and Jones; Wacey, 2002) may support this interpretation. The yoghurt-like zone immediately below the sediment–water interface is anoxic (Wacey, 2002), and thus SRB are found right at this interface (Wacey, 2002). Both field (Vasconcelos and McKenzie, 1997) and laboratory (Folk, 1993a; Vasconcelos et al., 1995; Warthmann et al., 2000; Wacey 2002, Wright and Wacey 2004, 2005) studies have identified the in situ mineralization of bacteria and have suggested that the bacteria play an active role in the nucleation of carbonates. This means that dolomite likely precipitates within the unconsolidated and water-saturated yoghurtlike mud horizon (upper 1–10 cm), within or just below the upwardly-migrating sulphate reduction zone. As the lakes undergo the final stages of desiccation, this yoghurt-like horizon is the last area to retain water and so at this stage lakewater and porewater effectively become one in the same. In Halite Lake, aragonite precipitation temperatures of 27–33 °C as shown in Fig. 9d correspond to observed temperatures within the lake and upper porewaters indicating aragonite precipitation during November in this lake. The situation in North and South Stromatolite Lakes is more similar to the dolomitic lakes with aragonite precipitation likely being in December. The timing of precipitation of these minerals appears at first to be somewhat different to that found by Rosen et al. (1995) for a hypersaline lake in Western Australia (Lake Hayward). They found that precipitation of carbonate phases occurred in late summer and early autumn. This is not unexpected, however, because Lake Hayward is a much deeper lake (minimum 2 m water depth in summer, Rosen et al., 1995), hence it is not ephemeral. The carbonate phases in Lake Hayward are still precipitating at the time of maximum evaporative concentration of the lakewater, just as they do in the Coorong, when conditions are most hostile to Metazoa

and when benthic microbial communities (BMC) dominate the lake ecosystem. Indeed Rosen et al. (1995) attribute carbonate precipitation in Lake Hayward to microbial mediation and state that aragonite precipitation could be at the water–BMC interface. Dolomite formation at 60–70 cm depth in Lake Hayward is also attributed to interaction with porewaters that have been modified by sulphate reduction (Rosen and Coshell, 1992). 5.3. Carbon isotopes Previous studies have used carbon isotopes to help to formulate a mechanism for dolomite formation. Botz and Von der Borch (1984) reported two different types of dolomite based on δ13C values. They suggested that isotopically light dolomite (∼ −1‰ V-PDB) formed from evaporitically modified continental groundwater, whereas isotopically heavy dolomite (up to +4‰ VPDB) formed in equilibrium with atmospheric CO2 by the dolomitization of aragonite. This interpretation appears flawed. No aragonite co-occurs in the upper sediments of dolomitic lakes, so the dolomite is very unlikely to be forming from alteration of aragonite. As mentioned earlier, Rosen et al. (1989) also differentiated two types of dolomite based partly on their isotopic signatures. The dolomite with the heavy carbon isotope signal (+ 3.5‰ average), Type A, was interpreted to precipitate rapidly from evaporating brines with slightly elevated Mg/Ca ratios. In contrast, Type B with the lighter carbon isotope values (− 1.2‰ average) was interpreted to form from less evaporitic brines, with a contribution from light meteoric or soil CO2. Rosen et al. (1988) were the only Coorong authors to mention the possibility of an organic component to this carbon but the evidence proved equivocal, and evaporation was eventually favoured. We have shown that dolomite precipitation in these lakes is bacterially-mediated (Wacey, 2002; Wright and Wacey, 2004; Wright and Wacey, 2005), in systems with an abundance of SRB, and yet it does not possess a characteristic ‘organic’ negative δ13C signal. Only two of the lakes exhibit a light δ13C signature (McFaiden Lake and Dolomite Lake, Fig. 3) and even here the values are much heavier than would be found for purely organically-derived carbon (typically lower than − 20‰, Schidlowski, 2001). Whilst strongly negative δ13C values traditionally suggests the incorporation of carbon from oxidized organic matter, slightly negative to slightly positive values, such as those found here, clearly do not rule out a contribution from bacterial processes.

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The δ13C signal found in the Coorong carbonates can be explained by the interplay of organic and inorganic carbon reservoirs (cf. Compton, 1988). Organogenic dolomite has been shown to posses a wide range of δ13C values, from around −15‰ to +20‰ (e.g., Spotts and Silverman, 1966; Murata et al., 1969; Pisciotto and Mahoney, 1981; Compton, 1988) which depend upon the relative contributions of organic carbon from sulphate reduction, methanogenesis and fermentation, plus the amount of inorganic carbon from seawater/lakewater bicarbonate and possibly from dissolved aragonite/calcite. Geochemical modeling of the Coorong lakes (Wright, 1999; Wright and Wacey, 2004) indicates that the carbonate ions being incorporated into these sediments originate mostly from inorganic bicarbonate already present in the lake- and porewater, but the availability of these carbonate ions is controlled by bacterial processes. Bacterial sulphate reduction results in pH elevation and a shift of the equilibrium between inorganic bicarbonate and carbonate ions, favouring the carbonate ion. Thus, the metabolism of SRB results in biochemical changes to the environment. SRB use sulphate as their terminal electron acceptor and consume cyanobacterial and other organic matter in the anoxic environment. They thereby liberate ammonia, which is highly soluble. Ammonia dissolving in the lake or porewater significantly increases alkalinity, raises pH and shifts the carbonate equilibrium towards the carbonate ion. − NH3 þ H2 O→NHþ 4 þ OH

OH− þ HCO−3 →H2 O þ CO2− 3 (Slaughter and Hill, 1991) If the carbon in the dolomite/aragonite was derived solely from inorganic species in equilibrium with dissolved CO2 in the lakewater, then the δ13C values should reflect equilibrium fractionation from the atmospheric CO2 value of −7‰ (Hoefs, 1987). Hence δ13C values for the aragonite precipitates should be around +3‰, with the dolomites around +5‰. It is clear that the majority of lakes are not precipitating carbonate in equilibrium with atmospheric CO2 and that there is some input from isotopically light carbon from degradation of organic material by bacterial sulphate reduction (note that methanogenesis would drive the δ13C signal to heavier rather than lighter values). In McFaiden and Dolomite lakes this bacterial sulphate reduction input is significant. Thus, SRB act as catalysts for dolomite precipitation and the δ13C value is derived from a mixture of inorganic carbonate carbon from the lakewater (probably marine in origin) and some organic carbon

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from decaying plant and microbial material. The variation in carbon isotope signature between the lakes can be explained by precipitation of carbonate minerals using varying amounts of heavy carbon from the inorganic reservoir and light carbon from the organic reservoir. These data add to the considerable database on organogenic carbonates and provide further evidence that bacterially-mediated dolomite, and indeed other microbial carbonates in the rock record, need not posses a strongly negative (purely organically derived) δ13C signal. 5.4. Sulphur isotopes Seawater spray is the most likely sulphate source (δ34 S ∼ 21‰), because the lakes are adjacent to the coast and the prevailing wind is onshore from the Southern Ocean (Von der Borch, 1976; Wacey, 2002). A second possible sulphate source could be from groundwater discharge. However, the lithology in the area where the groundwater originates is a carbonate sequence with no evidence of evaporites from which high concentrations of sulphate could be dissolved (Ludbrook, 1961). Moreover, measurements made in the field show no significant dissolved sulphate in stream water passing into the area. Rainwater sulphate will have δ34S values inherited from its source and because the weather patterns in this area mostly bring rain bearing moisture onshore from the sea, seawater will again be the dominant source. In low temperature restricted environments such as in the Coorong Lakes, a 34S enrichment of residual sulphate, as seen here, can be confidently attributed to bacterial sulphate reduction because other processes that may be operating in the lakes (e.g. precipitation of sedimentary sulphate or inorganic oxidation of sulphide to sulphate) have negligible isotopic effects. The limited variation of isotopic signal with time and depth in each lake (Figs. 4 and 5) is somewhat surprising given both the large cell numbers of SRB present (Wacey, 2002; Wright and Wacey, 2004, 2005) and the highly active nature of sulphate reduction at this time of year (Wright, 1999, 2000). Although all lakes show enrichment of residual sulphate in δ34S over seawater sulphate (by as much as 8‰) indicating that each lake overall acts as a bulk closed system with respect to sulphate, we hypothesise that the relatively small shifts in δ34S values with time and depth within each lake may be due to one or more of the following: 1) Lakewater to porewater behaving as an open system; 2) Species specific and type of organic matter effects; 3) Oxidation of H2S followed by further fractionations and transport of sulphate.

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If these lakes were acting as completely closed systems, then as bacterial sulphate reduction proceeds in the upper few centimetres of the sediments, sulphate concentration will decrease and the residual sulphate will become increasingly enriched in 34S according to a Rayleigh distillation model. Sulphate concentrations almost invariably decrease with depth in all of these lakes; on average between 46% and 69% of sulphate is reduced in the top 50 mm of the sediment (Wright and Wacey, 2004) and yet δ34S only increases with depth in some of the lakes, and even then only by very small amounts, despite the high cell numbers of SRB in the sediment. It could be concluded from this that the lakes are not behaving as closed systems, i.e. the porewater is not isolated and is not undergoing Rayleigh distillation. The very high initial concentrations of sulphate in the lakes (up to 75,000 mg/l versus 2440 mg/l for normal marine water) may help to explain the partial open system behaviour of the lakes. If there were only small amounts of sulphate present, this would soon be exhausted by bacterial sulphate reduction, and the last residual pool of aqueous sulphate would show a strong isotopic signal (e.g., Habicht and Canfield, 2001). In the Coorong lakes the source pool of sulphate is so large, even intense sulphate reduction does not reduce the concentrations to the very low levels seen in Rayleigh distillation models. Even at 4 or 5 cm depth our sulphate concentrations were still in excess of seawater. Although the isotopic signal produced by bacterial sulphate reduction is overprinted somewhat by the initial signal from the large sulphate pool, substantial bulk sulphate reduction must be occurring on a lake scale to fractionate δ34S values by as much as 8‰ (Milne Lake) compared to the source seawater sulphate. Species specific isotope effects as discussed by Detmers et al. (2001) may also be important in determining the δ34S of each lake. Detmers et al. (2001) found that the fractionation factors of different SRB species and genera growing under their optimal conditions could be vastly different, reporting a wide range between 2 and 42‰. Strains isolated from hypersaline mats, a similar environment to that of these lakes, showed some of the smallest fractionation factors (Desulfovirio halophilus = 2‰, Desulfovibrio oxyclinae = 4.5‰) whilst a strain which has been identified in McFaiden Lake, Desulfosarcina variabilis (Wacey, 2002), showed the lowest fractionation factor (15‰) of the complete oxidizing species tested. SRB species have been cultured and identified from two of the lakes (Milne and McFaiden) and have been shown to be unique to each lake (Wacey, 2002). We found no correlation

between the numbers of SRB and the δ34S signal or the amount of organic carbon (%TOC) and the δ34S which leads us to believe that species specific effects are more important in these lakes, in that each lake has developed a unique collection of bacteria which in turn give it its characteristic isotopic signal. The type of organic matter available for the SRB to degrade may also be important; Canfield (2001) and Kleikemper et al. (2004) found that widely variable fractionation factors could be obtained using the same strains of bacteria by changing their physiological pathways using different growth substrates. Thus, even if some lakes were to contain similar bacterial populations and similar amounts of TOC, the lakes may inherit different isotopic signals due to the type of organic matter available to the bacteria. Microbially-mediated transformations of sulphurcontaining species by means other than sulphate reduction may have an important influence on the sulphur isotope systematics in each lake. Jorgensen (1982) states that, commonly, 80–95% of H2S produced from sulphate reduction is re-oxidized. There have also been recent reports of sulphide oxidation in the dolomite containing lakes of S.E. Brazil (Moreira et al., 2004) with surface sulphate possibly inheriting the sulphur isotope signature of isotopically light H2S, and reports of very efficient sulphide oxidation in organic rich shelf carbonates with high rates of sulphate reduction (Ku et al., 1999). In the Coorong lakes we know that there is a certain volume of H2S escaping into the lakewater, as attested by the pungent rotten egg-like smell as one walks through the lakes and the complete absence of sulphide precipitates in the uppermost sediment (Wacey, 2002). This H2S can easily be re-oxidised in the Coorong lakes as the lakewater itself is fairly well oxygenated (Wacey, 2002); even in the absence of oxygen, nitrate and/or nitrite can be used by bacteria to oxidize H2S to intermediate sulphur-containing species such as elemental sulphur. Once intermediate sulphur species are formed these can undergo further bacterial reactions (e.g., Canfield and Thamdrup, 1994). This, together with similar disproportionation reactions for sulphite and thiosulphate may balance much of the heavy isotope enrichment caused by sulphate reduction, resulting in each lake obtaining a relatively steady isotopic state. This idea is supported by oxygen isotopes in the lakewater sulphate suggesting oxidation of sulphur and isotope exchange between sulphur species and the ambient lakewater, and by the immensely high concentrations of sulphate in the lakewaters (Wright and Wacey, 2005) suggesting sulphate recycling. It is unfortunate that sulphide minerals do not co-exist with the carbonate sediments in the uppermost yoghurt-

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like horizons of these lakes. Combined δ34S from sulphate and sulphide would clearly further enhance our understanding of the sulphur cycle in these lakes, particularly in light of recent experimental work (Brunner and Bernasconi, 2005) suggesting that almost all naturally occurring fractionations of sulphur isotopes between sulphate and sulphide in low temperature environments can be explained in terms of bacterial sulphate reduction. 6. Conclusion and wider implications This new isotopic data builds upon earlier work in the Coorong and adds supporting evidence for a microbial model of dolomite formation in the distal ephemeral Coorong lakes. Whilst our understanding of the complex interplay of processes is evolving, it is no accident that primary dolomite precipitation is occurring in the presence of intense bacterial sulphate reduction. Co-evolution of δ34Ssulphate-residual and δ18Osulphate-residual is seen only in the dolomite precipitating lakes indicating the important contribution of microbial sulphur cycling within these lakes. The application of the Vasconcelos et al. (2005) oxygen isotope fractionation factor equation to our data results in significantly improved and more realistic estimates of the temperature of precipitation of dolomite in the Coorong lakes over those obtained using an inorganic precipitation mechanism, again suggesting microbially catalysed precipitation of dolomite in this environment. δ18O values for lakewater combined with those of the carbonate phases show that, in the larger lakes, dolomite does not precipitate at this particular time (November) in the annual evaporative cycle. Instead, it precipitates slightly later (December) when macrofauna is entirely eliminated from the lakes and further degradation of organic matter provides abundant magnesium and carbonate alkalinity, and intense evaporation starts to lower δ18O lakewater. At this time, the last remaining water is held in the yoghurt-like mud horizon and so dolomite likely precipitates at shallow depth within the zone of sulphate reduction in this horizon, using SRB and other microbes as nucleation sites. Conversely, in Halite Lake, aragonite appears to precipitate during November where observed and calculated temperatures of precipitation overlap. Our δ13C data provide more evidence that microbiallymediated carbonates can possess a wide range of δ13C values, and hence δ13C alone cannot characterize a microbial carbonate in the rock record. In the Coorong lakes, the reason dolomite does or does not precipitate appears to be a delicate balance between

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the amount of SRB (dolomitic lakes have between 2.5 and 3500 times more SRB than the non-dolomitic ones), the amount of sulphate (and amount of sulphate reduced), the amount of magnesium, and the amount of organic matter degraded. These 4 factors are all partially or totally biogenically controlled. Clearly in South Stromatolite lake, where there are large amounts of SRB but no dolomite, the microbial push is present but not quite great enough to favour dolomite over aragonite — this may change with a very small change in the lake ecology (e.g., more (or a different type) of organic matter for the SRB to degrade; different species of SRB developing that are more efficient at reducing sulphate). A microbial model for widespread dolomite formation may also be applicable for massive dolomite formations. For example, during the Precambrian, benthic microbial communities dominated the ocean environment because Metazoa, which today graze the microbes and restrict them to rather extreme and inhospitable environments, had not evolved. Consequently, bacterial sulphate reduction was a commonplace and widespread process and the capability to overcome the kinetic inhibitors to dolomite formation in seawater would be greatly enhanced. In addition, the model may be applied to ancient Coorong-type settings such as the Skillogalee Dolomite of South Australia where desiccation cracks, tepee structures and an aphanitic texture lacking associated evaporite minerals (Von der Borch, and Lock, 1979) are directly comparable to the modern Coorong ephemeral lakes. Acknowledgements We acknowledge with gratitude the support and advice of Chris von der Borch and Nick McClure of Flinders University, South Australia. For training and analytical support, we wish to thank Tony Fallick at SUERC, Kevin Sharkey at Leicester University, and Steve Wyatt at Oxford University. This work was supported by NERC grants GT04/98/222/ES to DW, and IP/595/0499 to Martin Brasier. SUERC is supported by NERC and the Consortium of Scottish Universities. AJB is funded by NERC support of the Isotope Communities Support Facility at SUERC. References Ahmad, R., Hostetler, P.B., 1988. Recent advances in the study of Holocene dolomitic carbonate sedimentation in the Coorong area of South Australia. Aust. Geol. Soc. Ann. Meeting, pp. 40–41. Ahmad, R., Hostetler, P.B., 1994. Hydrogeochemical controls on the formation of primary dolomite in some ephemeral lakes in the

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