connections between sulfur cycle evolution, sulfur isotopes, sediments,and base metal sulfide...

8/11/2019 Connections Between Sulfur Cycle Evolution, Sulfur Isotopes, Sediments,And Base Metal Sulfide Deposits. FARQUHAR Et Al, 2010.

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Sedimentary and Geobiological Perspectiveson Sulfur Cycle Evolution

SEDIMENTOLOGISTS , geochemists, and geobiologists have longtried to address critical questions about sulfur cycle evolution. What is the origin of oceanic sulfate? What is the timing of theonset of bacterial sulfate reduction? How are sulfur speciestransformed and transferred between different reservoirs?Some workers have used geologic approaches to trace evidenceof formation of sulfur-bearing minerals in sedimentary settings(e.g., Grotzinger and Kasting, 1993; Pope and Grotzinger,2003; Schrder et al., 2008). Others have developed methodsthat link sulfur, iron, oxygen, and trace metal geochemistry (e.g., Anbar and Knoll, 2002; Algeo and Lyons, 2006; Canfieldet al., 2008; Scott et al., 2008). Still others have calibrated seawater geochemistry and sulfate concentrations using fluid in-clusions (e.g., Horita et al., 1991, 1996, 2002; Lowenstein etal., 2005). Some have used biological methods to documentthe presence of different types of sulfur-utilizing organisms inancient environments (Brocks and Schaeffer, 2008), and oth-ers have developed and applied methods that rely on mea-surements of the stable isotopes of sulfur and oxygen to tracechanges in the cycling of sulfur and the types of transforma-tions that occurred at different times in Earth history (Mon-ster et al., 1979; Cameron, 1982, 1983; Canfield, 2001).

Changes in the sulfur cycle have been attributed to the evo-lution of Earth surface chemistry, evolving chemistry of theatmosphere and oceans, and to accompanying changes inecology of sulfur-utilizing organisms. There is a generally, butnot universally, accepted model for sulfur cycle evolution.This model includes an early stage (Archean and earliest Pro-terozoic) with low sulfate oceans (less than 100200 M), a

component of inferred atmospheric origin that reflects low oxygen content, and a small role for sulfate reduction. Thismodel calls for a change in the sulfur cycle to slightly higheroceanic sulfate concentrations (millimolar) as a result of greaterproduction of sulfate by oxidative weathering of continentalsulfide starting at ~2.4 Ga and extending for almost the entireProterozoic. This change also included more significant recy-

cling of sulfide by oxidative weathering and ultimately led toa change in ocean chemistry by the Mesoproterozoic. A con-sensus on the sulfur cycle in the Mesoproterozoic has notbeen reached because there appears to be evidence for sul-fidic conditions in some settings and evidence for nonsulfidicconditions in others. Debate presently focuses on the extentof sulfidic and nonsulfidic ocean domains. Sulfate concentra-tions are thought to have risen to levels of 20 to 30 mM by theend of the Proterozoic (present-day concentrations), but tohave declined twice in the Phanerozoic to levels of a few mM(in the latest Cambrian to Early Devonian, and in the Juras-sic). Nevertheless, sulfate levels throughout the Phanerozoic were higher than those in the Proterozoic. The controls onsulfate levels reflect the supply of sulfate via oxidative weath-

ering and the efficiency of pyrite burial, and possibly organicsulfur, as a sulfide sink.

The Isotopic Record of Sulfur Cycle EvolutionThe isotopic record is the most continuous archive of the evo-

lution of the sedimentary sulfur cycle. Historically, it has fo-cused on the study of 34S (34S = (34S/ 32S)sample /(34S/ 32S)reference 1). All of these values are given in units of per mil () andthe factor 1,000 is, therefore, not included in the definition.More recently it has been recognized that additional infor-mation can be provided by also considering variations amongother sulfur isotopes, particularly using 33S, and 36S ( 33S

Connections between Sulfur Cycle Evolution, Sulfur Isotopes, Sediments,and Base Metal Sulfide Deposits

JAMES FARQUHAR,1,2, NANPING W U,1 DONALD E. CANFIELD ,2 AND HARRYODURO11Department of Geology and Earth System Science Interdisciplinary Center (ESSIC), University of Maryland,College Park, Maryland 20740

2 NordCEE and the Institute of Biology, Syddansk Universitetet, 5020 Odense C., Denmark

AbstractSignificant links exist between the sulfur cycle, sulfur geochemistry of sedimentary systems, and ore deposits

over the course of Earth history. A picture emerges of an Archean and Paleoproterozoic stage of the sulfur cyclethat has much lower levels of sulfate (


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= (33S/ 32S)sample /(33S/ 32S)reference-[(34S/ 32S)sample /(34S/ 32S)reference]0.515and 36S = (36S/ 32S)sample /(36S/ 32S)reference-[(34S/ 32S)sample /(34S/ 32S)reference]1.9; see Appendix 1 for information on mass-depen-dent and mass-independent sulfur isotope geochemistry).

Isotopic record before ~2400 MaThe historical interpretations of 34S of sedimentary pyrite

and sedimentary sulfate were made using a record similar tothe updated version presented in Figure 1. It was recognizedthat this record preserves variations in the range of34S, thedifference between the seawater sulfate 34S and the minimaof pyrite 34S, over geologic time (cf. Monster et al., 1979;Cameron, 1982, 1983; Canfield, 2001). One of the times where a change was noted is approximately 2400 Ma. Themean 34S of sedimentary pyrite before 2400 Ma is approxi-mately 2 per mil and the range of variability is small (Fig. 1B),suggesting high degrees of pyrite burial and a diminished rolefor processes that produce highly fractionated sulfide sulfur.

The change at ~2400 Ma (Fig. 1) is attributed to an ecosys-tem response that is related to how metabolisms of sulfate-re-ducing bacteria respond to different sulfate concentrations(e.g., Cameron, 1982, 1983; Canfield, 2001). Early work(Harrison and Thode, 1958) demonstrated that the discrimi-nation among sulfur isotopes generated during the process of sulfate reduction by bacteria was smaller when sulfate con-centrations were very low. More recent work has confirmedthis and has added a dimension to the interpretation of the

change at 2400 Ma by identifying and calibrating a depen-dence on ecosystem level expression of these isotope fraction-ations that results from changes in sulfur transport and sulfatereduction rates within unbioturbated sediment at low sulfateconcentrations (Canfield et al., 2000; Habicht et al., 2002).The results of Habicht et al. (2002) suggest that sulfate trans-port within the sediment is correlated with sulfate concentra-tions in the overlying water column. This limits the preserva-tion of the fractionations produced by sulfate-reducingbacteria in sedimentary pyrite and in the geologic record, plac-ing limits on oceanic sulfate concentrations at less than 200 M before ~2.4 Ga, and greater than 200 M after this time.

The record of 33S and 36S provide additional insight intothe evolution of the sulfur cycle at this time. The variations of 33S and 36S in rocks older than ~2.42 Ga have been attrib-uted to mass-independent frationation of sulfur isotopes(MIF-S) in atmospheric source reactions. This signal hasbeen interpreted to reflect atmospheric and surface weather-ing environments with significantly lower oxygen levels forthe principal reason that it is a prerequisite for the produc-tion, transfer, and preservation of MIF-S in surface sulfurpools (Farquhar et al., 2000, 2007a; Kasting, 2001; Farquharand Wing, 2003, 2005). These observations support a modelof a pre-2.42 Ga sulfur cycle with a diminished role for bio-logical reduction of sulfate and for diminished oxidation of sulfide because active cycling would be a process that wouldhomogenize and remove the MIF-S signal (Farquhar et al.,

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FIG. 1. A. Compilation of34S vs. time gray circle symbols (Canfield, 1998; Prokoph et al., 2007; Canfield and Farquhar,2009) with running average for sulfide (gray line) and inference about seawater sulfate (black line) that we fit using a cubicspline to sulfate data younger than 2450 Ma and constrained by literature estimates for periods prior to 2450 Ma. The curveis similar to previous curves for the Phanerozoic. The positive values for seawater sulfate and average sulfide indicate an im-balance in the sulfur cycle record extending for almost 2000 m.y. until ~500 Ma. B. Histograms of the frequency of observations of sulfide (gray filled bars), and sulfate from the black curve (hatched bars), for five intervals in geologic time. Theheight of each histogram is normalized to the total number of samples in order to put them on a similar aspect ratio. Meansfor sulfide data are also plotted. Uncertainties on the mean are determined using a monte carlo resampling technique.


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2000). Table 1 synthesizes observations from the integratedrecord of sulfur isotopes (34S, 33S, and 36S), as well as pro- viding implications from these observations. There are sev-eral issues with the characteristics of the MIF-S signal thatare important in shaping our understanding of the cycling of sulfur in the Archean because they relate to the controls onthe manifestation of the 33S (34S and 36S) signal in thesedimentary rock record.

First, there is a change in the magnitude of the range of 33S (34S and 36S) over the course of the Archean. There ismore significant variability for 33S in the Neoarchean, adamped 33S signal in the Mesoarchean, and an intermediatemagnitude 33S signal in the Paleoarchean (Fig. 2).

Second, there is clear evidence for small-scale high fre-quency variations in sedimentary rocks, mostly shales, at thethin-section and hand-sample scale, but there are also indica-tions that underlying longer timescale correlations exist be-tween cores of roughly equivalent successions. Profiles arecollected for three roughly equivalent sections through theNaute Shale of the Transvaal Supergroup and two successionsthrough the equivalent Australian Mount McRae Shale (Fig.

3). These successions can be roughly correlated using twomarker beds: Brunos band in the deeper parts and the tran-sition to Kuruman and Brockman banded iron formation inthe upper parts. These data suggest the presence of positiveexcursions for 33S in samples from the lower parts of thesesuccessions. Variations in the correlation between 34S and 33S (and 36S) also appear to be shared between these sec-tions (Kaufman et al., 2007; Ono et al., 2009b; Fig. 4).

Third, a number of studies that used microanalysis of sulfurisotopes (Mojzsis et al., 2003; Papineau et al., 2005, 2007; Catesand Mojzsis, 2006; Kamber and Whitehouse, 2007; Philippotet al., 2007; Ono et al., 2009a) have demonstrated homo-geneity rather than heterogeneity for 33S on a thin-sectionscale. Of the studies that demonstrated heterogeneity, thereappears to be a relationship between variations for 33S andthe type of pyrite that is present in the thin sections. For twostudies of the same late Neoarchean successions describedabove from the Naute Shale, evidence of significant hetero-geneity was also demonstrated (Kamber and Whitehouse,2007; Ono et al., 2009a). The fine-grained pyrite preservedthe most significant 33S signals, whereas larger spheroidal

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TABLE 1. Observations and Implications from the Multiple Sulfur Isotope Record

Older than ~2.42 Ga The presence of Archean MIF-S implies processes (photochemical) in addition to microbial sulfate reduction played an important role in the sulfur

cycle Covariation between 34S, 33S, and 36S tells us that sulfur cycle processes were not efficient enough to homogenize MIF-S once it was present Grain scale to thin section scale heterogeneity for 33S (e.g., Ono et al., 2003; Kamber and Whitehouse, 2007) implies mixing of sulfur with different

MIF-S compositions, suggesting more than one sulfur pool (e.g., sulfate and sulfide-polysulfides-sulfur) or short residence times (low concentrations) forsulfur

Coherent relationships between 36S and 33S at the formation level (Fig. 4) (e.g., Farquhar et al., 2007; Kaufman et al., 2007) imply gross changes inMIF-S source reactions and/or ecology that vary in time and possibly with geography

Observation of MIF-S signals in hydrothermal VMS systems (Farquhar and Wing, 2005; Jamieson et al., 2006) and in marine sediments subject to diage-netic processes implies the existence of multiple MIF-S sulfur cycle pools

A fractionation of ~20 is observed for34S of the most 34S-depleted sulfides and seawater sulfate (Fig. 1)

~2.42 Ga to ~600 Ma Samples of the ~2.42 Ga Duitschland Fm. capture the transition from a MIF-S world to a MDF-S world (Bekker et al., 2004; Guo et al., 2009) The transition from a MIF-S world to a MDF-S world appears to be correlated with one of the Early Proterozoic carbon isotope excursions The average 34S of Proterozoic sedimentary pyrite appears to be near zero, (if not slightly positive) and implies high fractions of pyrite burial and a

missing34S-depleted sulfur pool Studies of basins have shown more negative34S is observed in deeper parts of sedimentary basins (Shen et al., 2002) implying the missing34S-depleted

sulfur pool may have formed and been buried in deeper water settings (Logan et al., 1995), possibly subducted (Canfield, 2004) Deep- to shallow-water variation in the34S of the pyrite pool may reflect distillation of sulfur by sulfate reduction in Proterozoic oceans, reflecting low

sulfate concentrations (Johnston et al., 2006) A relationship between 34S and 33S is seen in some sections that implies an isotopic reservoir (Rayleigh) effect (Johnston et al., 2005, 2006) Significant (short time scale, high frequency) variations of34S for carbonate associated sulfate (a proxy for seawater sulfate) have been interpreted to in-

dicate low (mM) sulfate concentrations (Kah et al., 2004; Lyons et al., 2004)

The 33

S of seawater sulfate rose during the middle Proterozoic (between ~1.6 and ~1.3 Ga) (Johnston et al., 2005a) and this observation is interpretedto be related to a change in the ecology of the ocean-sediment system to favor expression of fractionations of sulfur disproportionators A fractionation of ~40 is observed for34S of the most 34S-depleted sulfides and seawater sulfate (Fig. 1)

Since ~600 Ma The variability with time of34S for seawater sulfate generally diminishes (e.g., Kampschulte and Strauss, 2004; Claypool et al. 1980), but some evidence

for higher frequency oscillations are still present (e.g., Paytan et al., 1998, 2004; Kurtz et al., 2003; Lyons et al., 2004); this has been interpreted to re-flect growth of the oceanic sulfate pool to a size large enough to dampen, but not completely eliminate, the variability in34S caused by changes in therate of or fractionations associated with pyrite burial

A fractionation of ~60 is observed for34S between the most 34S-depleted sulfides and seawater sulfate (Fig. 1) The magnitude of the fractionation between buried pyrite and sulfate appear to have changed from ~25 before ~550 Ma to ~35 between ~550 and

~300 Ma before increasing to ~43 by about 250 Ma (Wu et al., 2010)


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FIG. 2. Plot of 33S versus sample age from a compilation of values reported in (Farquhar et al., 2000, 2002, 2007a; Huet al., 2003; Mojzsis et al., 2003; Ono et al., 2003, 2006a, 2007; Bekker et al., 2004; Farquhar and Wing, 2005; Johnston etal., 2005a, 2006, 2008; Papineau et al., 2005, 2007; Cates and Mojzsis, 2006; Ohmoto et al., 2006; Kamber and Whitehouse,2007; Domagol-Goldman et al., 2008; Partridge et al., 2008; Thomazo et al., 2009). Gray filled circle symbols are samplesolder than 2.45 Ga and white filled circle symbols are for samples younger than 2.45 Ga. Ancient samples reflect mass inde-pendent reactions and preservation of isotopic signals in the rock record. Samples younger than 2.45 Ga are interpreted tocarry a mass conservation-type signal associated with mass-dependent fractionation by BSR and other sulfur cycle processes(see inset for sulfate).

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FIG. 3. Plot of 33S as a function of depth in five drill cores that intersect roughly equivalent successions of the NauteShale (A, B, C GKP01, GKF01, and AD-5) and the Mount McRae Formation (D, E ABDP-9 and AUS 493). The upperdark line represents the top of the formation and is taken as a time boundary. The lower dark line represents the bottom of the formation (corresponding to its contact with a BIF horizon referred to as Brunos band). Where the lines (bands) are gray instead of black, the position of this horizon is more uncertain. For panel C (AD-5) the top of the gray band represents thebottom of the drill hole, and the bottom of the gray band is 78 m lower than the top of the formation which is similar to thatseen in cores GKP01 and GKF01. The data and the positions of the horizons are replotted from Ono et al. (2003, 2009b)using information about the thickness of the Mount McRae Formation and Brunos band from Kakegawa et al. (1998) andinformation about the GKP01 and GKF01 cores from Knoll and Beukes (2009) and Kaufman et al. (2007) using informationfrom Beukes et al. (1990) for the stratigraphy of AD-5 core.


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grains and concentrations of pyrite in the same sections pre-served near zero or even negative 33S. These variations wereinterpreted to reflect different generations of pyrite and for-mation in a gradient, or stratified, system, including a strati-fied water column (Kamber and Whitehouse, 2007).

A number of models have been proposed to explain 33S variations in the geologic record. They include suggestionsthat the magnitude reflects variations in the proportion of at-mospheric sulfur with a much larger initial 33S (e.g., Far-quhar et al., 2000, 2001), variations in 33S of the atmos-

pheric components and therefore directly reflect a change inthe atmospheric source reactions (Ono et al., 2003; Domagol-Goldman et al., 2008; Farquhar et al., 2007a), or the natureand residence time of the sulfur pools in early Earth environ-ments (Farquhar and Wing, 2005; Ono et al., 2009a, b). It would appear that components of each of these suggestions would be required to account for the three observationsabove. A hybrid model that describes controls on the sulfurcycle is thus favored (Fig. 5).

A recurring question in the literature is whether there is anunderlying systematic variation in the magnitude of the mass-independent signal with respect to locality and or time, an

important issue for understanding the connections with thesystematics of sulfur in Archean ore deposits. There is clearevidence for heterogeneity on a hand-sample scale (e.g.,Kamber and Whitehouse, 2007; Ono et al., 2009a), but thereare also similarities in the magnitude of the mass-indepen-dent signals that appear to be preserved at least on a basinscale (Fig. 3). Resolution of this issue may be provided in partby considering the model (Fig. 5) in the context of pyrite for-mation and the atmospheric lifetime of volcanic sulfur.

Pyrite formation and burial in the present-day sulfur cycleis not uniformly distributed in oceanic sediments. Estimatesof sulfate reduction rates suggest that approximately 70 to 80percent of the sulfate reduction occurs at depths less than 200m (Turchyn and Schrag, 2004, 2006; Canfield et al., 2005).This number climbs to greater than 90 percent for depths lessthan 800 m, which represent approximately 40 percent of thetotal ocean area. Pyrite formation fixes sulfur ultimately de-rived from the sulfate pool in these areas. In the Archean,pyrite may also have acquired part of its signal from elemen-tal sulfur that is inferred to have had large 33S enrichments,but there may be a problem with invoking elemental sulfur asthe source of the large positive 33S values for pyrite. Whereas sulfate may accumulate in the oceans and be deliv-ered from seawater to the loci of pyrite formation, no accu-mulation mechanism has been described for elemental sulfur.Such a mechanism is needed to establish a distinct and largestanding pool of elemental sulfur for pyrite formation. It isalso possible that the deposition of elemental sulfur was notuniform on a global scale, and was instead controlled by thelocations of volcanic and perhaps biogenic sources of atmos-pheric sulfur gases, atmospheric residence times, and atmos-pheric transport. The second possibility is supported by calculations of the lifetime of sulfur dioxide before it is pho-tolyzed, which for sulfur gases in a low oxygen atmosphere is

less than one day (Farquhar et al., 2001). This is short com-pared to the global atmospheric timescales for mixing andrainout. If this were the case, then local deposition rates foratmospheric sulfur compounds without significant positive 33S values might come within an order of magnitude of thereduced sulfur produced by sulfate reduction in regions prox-imal to volcanic sources. The extent of these regions is notknown, but evidence to suggest that they may have been quitelarge comes from the general structure of 33S records for theHamersley, Transvaal, and Griqualand West basins (Fig. 3).

The change in the range of 33S through the Archean also,in part, reflects a change in the local fluxes of sulfur gases.The reason for such a change is unclear. It needs to be deter-mined whether large changes can be seen outside of the

Hamersley, Transvaal, and Griqualand West basins. It may re-flect changes in volcanic fluxes, or a new contribution frombiogenic reduced sulfur gases, or possibly a change in the sul-fur content of the volcanic gases resulting from the develop-ment of a pool that is recycled through subduction. The dif-ference between these Neoarchean signals and signals thatpreceded them remains an important focus for future work.

An alternative possibility (Watanabe et al., 2009) suggeststhat reactions other than those that occur in a gas phase may produce anomalous sulfur isotope signals. These authors pre-sent the results of sulfate reduction experiments using variousamino acids at temperatures between 150 and 200C. They


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Mt McRae UpperMt McRae LowerGamohaan UpperGamohaan Lower

FIG. 4. Plot of 33S versus 36S for pyrite samples from the Neoarcheanof South Africa (Gamohaan Fm.) and Western Australia (Mt McRae Fm.)(Kaufman et al., 2007). These two successions are considered to be timeequivalent and can be correlated via marker beds. The upper and lower partsof the two formations exhibit similar changes in the relationship between 33S and 36S. The lower parts of these successions are more linear, follow-ing 36S ~ 0.85 33S, and the upper portions of these successions show morescatter and 36S >0.85 33S. Slopes on lines in plot are 0.85 and 1.37. Thisdifference has been interpreted to reflect a change in mass independentsource reactions and also an overprint by cycling of sulfur in the water col-umn in the upper parts of these two successions (Kaufman et al., 2007). It hasrecently been suggested that the upper interval of the Mount McRae For-mation may reflect local sulfidic conditions (Reinhard et al., 2009). In thiscontext, it is possible that a significant part of the change in 33S vs. 36S may reflect stronger expression of sulfur cycling by sulfate reducers in a sulfidic water column. This would not require a change in sulfate concentration, but

would imply more efficient expression of fractionations associated with sul-fate reduction in the water column and would also be consistent with asser-tions for water column sulfate reduction by Kamber and Whitehouse (2007).


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suggest that mass-independent fractionation under such con-ditions may involve a magnetic isotope effect (MIE) (e.g.,Turro, 1983; Buchachenko, 1995: see App. 1) or a surface cat-alyzed reaction (e.g., Lasaga et al., 2008), but further work will be needed to resolve this. This proposal might be takenfurther to suggest that atmospheric reactions are not thesource of the mass-independent signal, and that mass-inde-pendent fractionation (MIF) is a result of processes thatoccur in sediments. We do not think that this is presently adefendable suggestion because several issues need to be ad-dressed relating to the following: (1) the nature of the signal,such as the tight coupling of 33S and 36S in many Archeansamples (Fig. 4); (2) the production and transfer of the signalthrough the wide range of Archean sedimentary, metamor-

phic, and igneous rocks; and (3) the apparent geographic andtemporal coherence of the signal over the course of theArchean and its abrupt disappearance in the Paleoprotero-zoic. Anomalous sulfur isotope fractionation arising from suchreactions may be relevant in some natural environments where thermochemical sulfate reduction occurs, although ev-idence has yet to be presented that this is the case.

Isotopic record since ~2400 MaThe sulfur cycle since ~2400 Ma is characterized by a larger

range in the signal of34S and a lack of significant mass-inde-pendent signals in sedimentary rocks, suggesting an end to

the period of MIF-S and a transition to a sulfur cycle morestrongly influenced by biological processes (Bekker et al.,2004; Papineau et al., 2005; Guo et al., 2009). The upperparts of the Duitschland Formation, South Africa, de-posited between 2450 and 2320 Ma, capture this transitionfor both 34S and 33S. They also show a correspondencebetween 34S, 33S, and 13C, which reinforces inferencesabout a link between the carbon and sulfur cycles at thetime of the rise of oxygen. High13C values for carbonatesreflect carbon burial events that were associated with sig-nificant oxygen production (Karhu and Holland, 1996). Thecoincidence of these signals links sulfate concentration, at-mospheric oxygen levels, and the origin-disappearance of MIF-S.

The larger range of 34

S observed in the Proterozoic sug-gests sulfate concentrations higher than 200 M (Habicht etal., 2002). Sulfate was sustained at levels sufficient to allow the widespread expression of a biological signal (14 mM),but not at levels as high as those in the Phanerozoic (~530mM: Horita et al., 2002; Shen et al., 2002, 2003; Canfield,2004; Kah et al., 2004; Gellatly and Lyons, 2005; Johnston etal., 2006). In addition, sulfate concentrations varied withinthe oceans over Proterozoic time (Logan et al., 1995; Hurtgenet al., 2002, 2005; Shen et al., 2002, 2003; Johnston et al. 2006;Halverson and Hurtgen, 2007), and these variable sulfate con-centrations reflected drawdown in shallow or restricted basins

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primitive mantle( 33 S = 0)

r e c y c l e d s

e d i m e n t s

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atmospheric sulfur chemistryProduction of MIF

(negative 33 S)(positive 33S)

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33 S volcanic SO 2 and H 2S( 33S = 0?)

Proportion of atmospheric sulfur increasesas a result of proximity to sulfur sources

or to changes in sulfur flux(low frequency variations

sedimentary Py from S 0(positive 33S)

Sulfate reductionulfate reductiondilutesilutes 333S signal signalSulfate reduction

dilutes 33S signal

Atmosphere: atmospheric depositionof sulfur with 33S

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i t y t o

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Oceanic sulfideceanic sulfide(positivepositive 333S))

Oceanic sulfide(positive 33 S)

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Ocean: Two sulfur poolscean: Two sulfur pools

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where sulfate reduction rates were high and resupply by cir-culation was low. The average34S of sedimentary pyrite forPaleoproterzoic rocks remains similar to the value seen in theArchean (i.e., 2), implying abundant pyrite burial. The av-erage 34S of sedimentary pyrite for Mesoproterozoic andNeoproterzoic rocks increased to higher values (~4-7; Fig.1B), which implies that both sulfide and sulfate are34S en-riched relative to the long-term sulfur inputs. This indicatesthat a pool of 34S-depleted sulfur is missing from the knownsedimentary pyrite record. Such a pool may have been an off-shore pool that was lost to deeper water, sedimentary facies,and/or sulfide deposits. It would imply that sulfur isotopiccompositions and concentrations were heterogeneous in theProterozoic oceans. An interesting suggestion made by Can-field (2004) is that the size of the exogenic sulfur reservoir was affected by loss of a deep-water sulfide pool to rocks that were ultimately subducted.

The record of four isotopes since ~600 Ma supports a sta-bilization in isotopic variability for seawater sulfate that is re-lated to generally higher (430 mM) oceanic sulfate concen-trations and an average 34S that declines to negative values(Fig. 1B) for the first time, implying a less significant role forpyrite burial. This interpretation is consistent with other di-rect types of evidence for higher sulfate levels that come fromfluid inclusions (e.g., Horita et al., 1991, 1996, 2002; Lowen-stein et al., 2005) and suggests high sulfate concentrations atleast as early as the latest Neoproterozoic. (Horita et al., 2002;Lowenstein et al., 2005).

It has been suggested that the rise in sulfate concentrationsince the latest Neoproterozoic and the change in the fractionof sulfur lost to pyrite burial are related to oxygenation of shelf environments and the establishment of conditions fa- vorable for animal radiation and bioturbation (Canfield andFarquhar, 2009). Furthermore, changes in the magnitude of

the fractionation associated with buried pyrite in thePhanerozoic are related to changes in oxidation pathways, which may have similar links to the evolution of open diage-netic systems on a global scale (e.g., Wu et al., 2010). Thechange in the magnitude of the fractionation between sulfateand buried pyrite at ~540 Ma could be related to bioturbationbecause of the way that it affects recycling of sulfur from sul-fate reduction (Fig. 6), and also because of links between bio-turbation, sulfate concentration, and the magnitude of frac-tionations produced by sulfate-reducing bacteria (SRB)(Canfield and Farquhar, 2009). This presents an interestingquestion related to the observed change in the range of 34Sin the Neoproterozoic (e.g., ~800 Ma in Canfield and Teske,1996), the change in the fractionation between sulfate and

buried pyrite (Fig. 1), and the change in 33

S starting at ~1.4Ga reflective of a change in oxidation pathways and dispro-portionation (Johnston et al., 2005a). If we accept that thechange in 33S reflects the onset of disproportionation andthe change in the fractionation between sulfate and buriedpyrite is associated with bioturbation and the rise of oceanicsulfate, then we are left with the question of what is the sig-nificance of the change in the range of34S at ~800 Ma. Can-field and Teske (1996) suggested that this change is the ex-pression of disproportionation in the sedimentary record.This may be the case if the signal seen by Johnston et al.(2005a) was for water column disproportionation, where the

sulfide pool was large and relatively homogenous. This wouldimply a retreat of the sulfur ecology, with disproportionation,into sedimentary diagenetic environments at ~800 Ma.

The Record of Sulfur Cycle Evolutionfrom Sulfide-Bearing Ore Deposits

Several types of sulfide-bearing ore deposits can be linkedto oceanic sulfur chemistry and/or to the isotope variationsobserved in the surface sulfur cycle. These include vol-canogenic massive sulfide deposits, clastic-dominated Pb-Zndeposits, sediment-hosted copper deposits, and Mississippi Valley-type (MVT) deposits.

Volcanogenic massive sulfide (VMS) deposits Volcanogenic massive sulfide (VMS) deposits occur through-

out geologic history, with examples identified in the Archean,Proterozoic, and Phanerozoic. The VMS deposits are associ-ated with volocanogenic settings and hosted in volcano sedi-mentary piles (see Huston et al., 2010). These deposits aremetal-sulfide deposits, principally Fe, Cu, Pb, and Zn, associ-ated with large-scale hydrothermal circulation systems thatinvolve seawater and may involve sulfur leached from igneousand other crustal rocks or from reduction of seawater sulfate.Ono et al. (2007) show for sulfide in modern black smokersfrom the East Pacific Rise, that 73 to 89 percent of the sulfideoriginates from leaching of juvenile sources and 11 to 27 per-cent originates from reduction of seawater sulfate. Peters etal. (2010) apply similar techniques to the Mid-Atlantic Ridgeand find slightly higher proportions of sulfide derived fromthe reduction of seawater sulfate that they attribute to thedeeper circulation of the hydrothermal system. These sys-tems are thought to be present-day analogs of ancient VMSsystems.

Present-day hydrothermal systems in active extensional set-

tings are also thought to be analogs for ancient VMS depositformation. These deposits generally grade from sulfide stock- work zones into overlying strata-bound massive sulfide andexhalites. Sulfide minerals may form in the water column as aresult of a variety of processes, including mixing or cooling of metal- and sulfide-bearing fluids, and may depend on factorssuch as the metal/sulfur ratio and the presence of sulfidic bot-tom waters (e.g., Goodfellow and Peter, 1996; Tornos et al.,1998). The recharge zones where seawater infiltrates the sys-tem may also be sites for precipitation of sulfate minerals,such as anhydrite that precipitates by retrograde solubility along flow paths that progress up-temperature. Much of theanhydrite is not preserved in older deposits because it is lostonce these systems cool. Other sulfate minerals such as barite

form when solutions containing sulfate mix with barium-richsolutions. The isotopic composition of sulfur in these depositsrecords information about the state of the oceanic sulfurcycle.

SEDEX CD-Pb-Zn sulfide depositsSedimentary exhalative (SEDEX) deposits are a type of

sedimentary rock-hosted sulfide-bearing deposit that forms when metal and sulfide-bearing fluids are mixed, cool, orevolve to a different pH in syngenetic to diagenetic sedimen-tary environments. The fluids that carry metals have migratedpredominantly from underlying clastic sedimentary rocks and


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fluid flow is driven by high geothermal gradients and focusedalong syndepositional extensional faults (Sangster, 2002; Yanget al., 2004). Formation of ore minerals may occur below thesea floor in permeable zones, but some SEDEX deposition isassociated with discharge of fluids into seawater to yield a

strata-bound deposit. Synsedimentary features are preservedin some SEDEX deposits, indicating a connection betweenthe sulfide minerals and the sediments that make up these se-quences. The connection with a vent is not always clear andevidence suggests that, in some cases, the brines migrate

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FIG. 6. A model system (box model representation) to illustrate how bioturbation may change the degree of openness of a diagenetic setting and alter the magnitude of the fractionation preserved between buried pyrite and seawater sulfate. B.Schematic plot showing relationship of fractionation between sulfate and pyrite and fraction of sulfur recycled to the oceanicpool. Another direct influence on fractionation related to bioturbation is coupled to the possibility of changes in sulfate con-

centration that feed back to the dependence of isotope discrimination by sulfate-reducing bacteria on sulfate concentrations(e.g., Canfield and Farquhar, 2009). C. Similar model system that incorporates disproportionation and D. Calculated fieldsfor overlying sulfate pool on plot of 33S versus34S. Calculated using fractionations for sulfate reduction from Farquhar andJohnston (2008) from experiments with natural populations of sulfate reducers and disproportionation data from Johnston etal. (2005b). Dark line is for sulfate reduction only. Dashed line includes reoxidation and disproportionation. Other fraction-ations yield different size fields, but the general shapes are preserved.


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considerable distances from the discharge site without mixing with seawater (Sangster, 2002).

Leach et al. (2010) argue that the term clastic-dominatedlead-zinc (CD Pb-Zn) deposits is more appropriate because itincludes SEDEX deposits, but provides a descriptive classifi-cation that does not imply a genetic link to exhalative processes.They note that many of the deposits lumped with SEDEX de-posits in other studies lack clear evidence for an exhalative ori-gin, despite many forming in the sediments (Leach et al.,2005). They suggest a further qualification that uses a desig-nation to describe CD Pb-Zn deposits in terms of the tectonicsetting in which they form, such as passive margin, continen-tal rift, continental sag basin, and back-arc basins. In thispaper, we will use the Leach et al. (2010) terminology.

MVT sulfide depositsMVT deposits are base metal (Fe, Zn, Pb) sulfide deposits

hosted in carbonate-dominated sedimentary rocks. The MVTdeposits are epigenetic deposits that are generally stratabound, but commonly not stratiform, with sulfide mineralsreplacing carbonates and filling open spaces in the host rocks(Sangster, 1990). Most MVT deposits are thought to form asa result of the mixing between sulfide-bearing and metal-bearing basinal brines, much like modern oil field brines thatare mobilized by tectonic events (Leach et al., 2001, 2005;Paradis and Nelson, 2007). Leach et al. (2010) suggest thatsulfate may be stored in these basins until remobilized by atectonic trigger, and the appearance of MVT deposits in thePhanerozoic implies a link to a threshold for establishing apool of stored sulfate. The ultimate source of sulfide isthought to be seawater sulfate because of its isotopic compo-sition and, in most cases, is thought to be the product of bio-logical or thermochemical sulfate reduction. The epigeneticprocesses by which most MVT deposits are thought to form

make them inherently difficult to date and also remove theirchemistry several steps from ocean chemistry and the surfacesulfur cycle, but links to the sulfur cycle remain and are ex-plored forthwith.

The MVT deposits are not recognized before ~2.2 Ga (thePering and Bushy Park deposits) and peak at times whenocean sulfate concentrations were highest (Kesler and Reich,2006). Leach et al. (2001) have shown that the best dated of the Phanerozoic deposits cluster during periods of tectono-thermally related flow of brines almost exclusively associated with closure of ocean basins. Two principal episodes of Phanerozoic MVT formation can be defined at ~400 to 300Ma and ~120 to 60 Ma. These two episodes correspond to pe-riods when seawater sulfate concentrations (e.g., Horita et al.,

2002; Lowenstein et al., 2003) were high, but not at theirmaxima (Fig. 7). They also correspond with times of tecton-ism and the link between tectonism and basinal brine migra-tion is the primary reason for their formation, although suffi-cient sulfate is also required (Kesler et al., 1995; Leach et al.,2001, 2005, 2010). The source of these basinal brines variesfrom deposit to deposit. It has been inferred on the basis of fluid inclusion Cl/Br ratios that the dominant source of someof these brines is evaporated seawater, rather than the disso-lution of evaporites (Kesler et al., 1995; Leach et al., 2001). Inother cases, there appears to be evidence for brines that werederived from the dissolution of evaporites. In the case of the

most ancient MVT deposits, at Bushy Park and Pering inSouth Africa, the mineralization appears to have occurred atdeeper crustal levels (2.84.8 km) than at other MVT deposits(Huizenga et al., 2006) and the34S, 33S, and 36S of the sul-fur (Schaefer; 2002; Kim et al., 2009) appear to be moreclosely related to sulfur from host rocks than from contempo-raneous or older seawater sulfate sources that are inferred to

have zero or negative33

S, respectively.Sediment-hosted copper deposits

The distribution of sediment-hosted copper deposits (Hitz-man et al., 2010) is also associated with the evolution of thesulfur and oxygen cycles. These are stratiform deposits of copper sulfides in siliciclastic and/or dolomitic sedimentary rocks. They first appear in the Paleoproterozoic after the riseof oxygen and sulfate, and have formed through the presentday (Brown, 1997). Their formation is attributed to the reac-tion of oxidized solutions containing copper, which are gener-ated during diagenetic reddening reactions of sediments(redbeds), with sulfide-containing solutions produced at, orintroduced to, the site of deposition (Hitzman et al., 2005;

Brown, 2005, 2009). Sulfur sources for the sulfide-bearing so-lutions include evaporites, basinal brines, and sulfidic hydro-carbons. Connections between sediment-hosted copper de-posits and the oceanic sulfur cycle are indirect because thesedeposits typically form in continental rift settings and arethought to be more common when rift settings are at low lat-itude, allowing evaporative concentration of sulfate (Brown,1997). This implies a connection to sulfate concentration, butalso calls for a concentrating mechanism. For example, sulfurisotope studies of the Central African Copperbelt (Cailteux etal., 2005) suggest reduction of seawater sulfate in rift basinsettings that are at times isolated from the oceans and provide


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FIG. 7. Comparison between histogram of frequency of MVT mineraliza-tion events (using data in Leach et al., 2001) and models of the evolution of

seawater sulfate concentrations from Lowenstein et al. (2005) (thick blackline) and Horita et al. (2002) (thick gray line with estimates of uncertainty given by thin gray lines). The figure shows a correspondence between thefrequency of MVT deposit forming events and tectonic processes, ratherthan a direct relationship with sulfate concentrations. It is not clear whethersulfate lags mineralization events or vice versa.


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stratigraphic evidence of control on isotopic composition by communication between basin sulfate and oceanic sulfate.

Connections between Seawater Sulfate andSulfide Ore Deposits

The shared connection through hydrothermal circulationsystems to seawater for both VMS and CD Pb-Zn depositsprovides a link to seawater sulfate. The connections betweenthe brines that form MVT deposits, as well as the sediment-hosted copper deposits, and preexisting evaporites and/or seawater sources provide a link between these deposits and thesulfur cycle, but as we have noted above, their epigenetic ori-gin adds complexity to the interpretation. The similarity in theisotopic compositions of sulfide and sulfate for VMS, CD Pb-Zn, and MVT deposits relative to that for data compiled fromthe sedimentary record is illustrated in Figure 8A. Note thatpyrite may have a much broader range in values relative to oresulfides such as sphalerite and galena (cf. Kelley et al., 2004;Leach et al., 2005). The common features of these two recordsinclude the small range of 34S for sulfide and sulfate in sam-ples of Archean age, an expanded range of34S in Proterozoicage samples that is biased to positive34S values, and a furtherexpanded range of 34S in Phanerozoic age samples. This cor-respondence illustrates a broad link that exists between thesulfur in these sulfide-bearing ore systems and sulfur fromother parts of the exogenic sulfur cycle. However, the parallelis not perfect. Whereas ore sulfides broadly overlap with thetotal range of observed sedimentary pyrite and sulfate, oresulfides are generally more enriched in34S (higher 34S) thansedimentary pyrite. This reflects a stronger link to sulfaterather than sulfide in surficial environments. The nature of theconnections is different for different classes of ore depositsand also for the different parts of the ore deposits, such asstockworks compared to exhalative lenses in VMS deposits.

The processing of sulfur in the ore depositsfor example, sul-fate in barite capsreflects Rayleigh fractionation that con-tributes to the variability. Changes in the distribution of34Sfor the three types of sulfide deposits can be noted over Earthhistory (Figs. 9, 10). The characteristics of these distributionsdiffer from those of the sedimentary record insofar as themean values are generally positive for both sulfate and sulfidefor all deposit types, which reflects the role played by sulfateand leaching of igneous sulfur, ultimately of juvenile origin, assources of sulfur rather than of pyrite produced from reducedsulfur contributed by sulfate reducers. As noted above, iso-topic studies of modern hydrothermal systems at the East Pa-cific Rise and Mid-Atlantic Ridge (Ono et al., 2007; Peters etal., 2010) suggest that the fraction of sulfur derived from sul-

fate is less than ~40 percent of the total sulfur and the re-mainder is derived from leaching of igneous sulfides. The dif-ference in 34S for sulfate and sulfide when they coexist inthese different types of deposits reflects a combination of mix-ing and fractionation processes in the ore-forming processes.The VMS deposits form at the highest temperatures (usually >300C), whereas the CD Pb-Zn and MVT deposits typically form at lower temperatures (


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sulfate and to these particular sedimentary rock-hosted VMSdeposits.The connections that are illustrated in Figure 8A are not,

however, straightforward because ore-forming processes alsoinfluence the 34S of the sulfide and sulfate in these deposits.Sulfate minerals from these deposits possess isotopic compo-sitions that generally extend to34S-enriched compositions andin some cases also include 34S-depleted sulfate (Fig. 8B).These variations can be attributed to reservoir effects associ-ated with removal of32S-enriched sulfide produced by sulfatereduction (Rayleigh fractionation and related effects: e.g.,Goodfellow, 1987) and the 34S-depleted sulfate may reflectoxidation of32S-enriched sulfide. Note that reservoir effectssuch as Rayleigh fractionation will be expressed in the water

column: the fraction of sulfate consumed in the water columnis limited by the availability of nutrients, export of sulfides by sedimentation, and eddy diffusion of sulfate and sulfide, so asignificant proportion of these may be associated with pore- waters. Johnson et al. (2009) demonstrated covariance for sul-fur and oxygen isotopes that evolve from compositions similarto seawater sulfate to higher 34S and generally higher 18O.The variation they report for34S is similar to that reported inother studies (e.g., Fig. 10) and they attribute this variation tothe processes associated with sulfate reduction in the region where barite precipitates. Broadbent et al. (1998) describe aprocess at the Century CD Pb-Zn deposits of the McArthur

Basin that involves an increase in the34

S/ 32

S of approximately 20 per mil with the evolution of the deposit. They interpretthis to reflect drawdown of the sulfate sulfur pool as a resultof sulfide sinks outcompeting sulfate resupply. This observa-tion also may be related to the observation of high34S in Pro-terozoic CD Pb-Zn deposits (e.g., Lyons et al., 2004). Thissuggestion implies that the fractional drawdown of sulfate inthese systems is greater when sulfate concentrations arelower. Although this link has not been documented (but seeJohnson and Emsbo, 2005), it would presumably draw on aninferred correlation between concentration and a functionthat relates the amount of sulfur given by the deposit size,1/(fractional drawdown of sulfate), and 1/(fluid flux). Johnsonand Emsbo (2005) suggest that the drawdown of basinal or

global oceanic sulfate would be minimal, implying that distil-lation occurs in the sediments and not in the water column.The persistence of high 34S in Phanerozoic CD Pb-Zn de-posits (Fig. 10) suggests that the record preserved by thesedeposits may not uniquely support this suggestion, althoughthey would not disprove it, either.

Insights from sulfide ore deposits for the Archean S-cycleThere are some features of VMS deposits that appear to

provide important information about the state of the oceanicsulfur cycle. Some of the oldest (34003200 Ma) VMS de-posits are preserved in Western Australia; the Kangaroo

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_ X= 23.6 10.4

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FIG. 10. Histograms of the frequency of observations of sulfide (gray filled bars), and sulfate (hatched bars), for Phanero-zoic MVT (A), Proterozoic MVT (B), Phanerozoic CD-Pb-Zn (C), and Proterozoic CD-Pb-Zn (D) deposits. The heights of each histogram are normalized to the total number of samples in order to put them on a similar aspect ratio. Mean and stan-dard deviation given on plots. Data sources listed in Figure 8.


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Caves and the Sulphur Springs deposits (Vearncombe et al.,1995; Golding and Young, 2005). They are characterized by mineralogical and textural features that are similar to younger VMS deposits with Cu, Pb, and Zn sulfide mineralization andappreciable barite. The deposits formed in a back-arc settingat a water depth greater than ~1,000 m and at or just below the sea floor (Vearncombe et al., 1995; Buick et al., 2002).They are contemporaneous with sulfate-bearing shallow- water deposits from other localities in Western Australia andSouth Africa (Heinrichs and Reimer, 1977; Buick and Dun-lop, 1990).

The presence of barite in the Sulphur Springs and KangarooCaves deposits indicates that barium was mobilized in thesubsea-floor hydrothermal system and that sufficient sulfate was present at the site of deposition to precipitate barite(Vearncombe et al., 1995; Huston et al., 2002). It can be ar-gued on the basis of 33S that the source of sulfate is marine(Farquhar et al., 2000; Golding et al., 2006), which is consis-tent with suggestions from other evidence (Vearncombe et al.,1995; Huston et al., 2002). These arguments also suggest thatthe early oceans may have had high sulfate concentrations andthat this requires an oxygenated early Earth. We do not sub-scribe to this suggestion because an oxygenated atmosphere isnot required for production of at least some sulfate in the ex-ogenic cycle (e.g., Walker and Brimblecombe, 1985; Kasting,2001; Farquhar et al., 2000, 2001). Sulfate will be producedby photochemical pathways from volcanic emissions even onan anoxic early Earth (Vanderwood and Thiemens, 1980; Walker and Brimblecombe, 1985; Farquhar et al., 2001; Kast-ing, 2001), and evidence for these reactions is present in the 33S of the sulfate in these deposits and elsewhere (Hoering,1989; Bao et al., 2007). The precipitation of sulfate requiresonly that the solubility product for barite be exceeded.

Whereas the amounts of sulfate present in these deposits

requires a sufficient supply of sulfate, the presence of baritemay not imply oceanic sulfate concentrations >200 M. Hus-ton and Logan (2004) propose that the oceans may have beenbarium enriched and sulfate limited. These authors present anumber of arguments, including many listed above, to makethe point that evidence does not exist to support sulfate-richoceans at this time. It is valuable to realize, however, that thepoint at which the solubility product of BaSO4 is balanced be-tween Ba2+ and SO42- is in the range of tens of M, which isbelow, but not significantly, the solubility of Ba2+ predictedfor an ocean equilibrated with 0.1 bar of CO2 . One-tenth of a bar of carbon dioxide is a level contained in Archean atmos-pheric energy balance models (e.g., Domagol-Goldman et al.,2008; Haqq-Misra et al., 2008). This relationship between

CO2, Ba2+

, and SO42

suggests that both species coexisted atlevels of a few tens to a few hundred M and that barite pre-cipitation occurred when these levels were exceeded.

Two other observations bear on the nature of the oceanicsulfate pool at this time in early Earth history. First, baritesamples from other Paleoarchean localities in Australia,South Africa, and India (Hoering, 1989; Farquhar et al., 2000;Bao et al., 2007; Ueno et al., 2008, Shen et al., 2009) alsocarry a 33S signature that implies the sulfate was derivedfrom seawater. In particular, the 33S signal is consistently negative and has been inferred to represent the seawater sul-fate pool. The 34S of this sulfate is not strongly positive

(~38) and thus the isotopic composition of seawater sul-fate was not strongly fractionated by global-scale burial of pyrite. Second, Foriel et al. (2004) use the Cl/Br of fluid in-clusions from the quartz-carbonate pods of the 3.49 GaDresser Formation, a shallow-water facies that is slightly older than the Kangaroo Caves and Sulphur Springs deposi-tional environments, to argue for mixing of Ba-rich hy-drothermal fluids with low sulfate seawater that had beensubject to evaporative concentration, consistent with otherconstraints on the nature of the local depositional environ-ment. The fluid inclusion data support an upper limit of 8mM for seawater sulfide. This is not as low as the estimatesmade using other lines of evidence; it is also the upper detec-tion limit of the analytical technique.

A number of Neoarchean VMS deposits do not appear tocontain similar barite mineralization. This reflects lower sul-fate concentrations or the nature of the hydrothermal circu-lation systems such that the size of the events in theNeoarchean may have been unusually large and that they may have influenced the chemistry of the overlying water, thusmaking barite precipitation unlikely. Both of these sugges-tions are difficult to test, and each may carry consequencesfor the sulfur cycle. For the first suggestion of lower globalseawater sulfate concentrations, the reasoning is that duringthe Neoarchean there was a return to more anoxic conditionsfrom an earlier Archean, more oxygenated environment. Asdescribed above, we do not believe that the evidence for anoxygenated Paleoarchean is compelling and another reasonfor a change in sulfate concentrations would be required. It ispossible, but not proven, that oceanic sulfate concentrationsmay have declined in the Neoarchean because of the estab-lishment of more effective sinks for sulfate, perhaps associ-ated with the establishment of a niche for sulfate reducers.This is difficult to test, but evidence for the hypothesis should

be present in the 34

S of seawater sulfate. Some isotopic datafor pyrite in carbonates, sulfide from marine sediments, andsulfides and rare barite in VMS deposits suggest that the34Sof Neoarchean sulfate had higher values (~1015; Thodeet al., 1991; Ono et al., 2003; Farquhar and Wing, 2005; Kauf-man et al., 2007). This may signal the onset of bacterial sul-fate reduction in the oceanic sulfur cycle and would imply agrowing sink of pyrite burial and larger fractionations associ-ated with this sink.

Cameron and Hattori (1987) report 34S analyses of pyriteand sulfate (barite and anhydrite) from Hemlo and relatedgold-rich possible VMS deposits in the Superior province thatprovide an exception to these more general observations of alack of sulfate and a lack of significant sulfur isotope fraction-

ation in Neoarchean deposits. They interpret these depositsto be the result of metal deposition in an environment with lo-cally high sulfate concentrations that indicate an Archeanocean with isolated sulfate-rich basins, despite a relatively sul-fate poor overall concentration. The range of variability for34S approaches 30 per mil and is interpreted to be a result of abiological processes. Thode et al. (1991) note that the frac-tionation of 34S/ 32S between coexisting barite and pyrite is uni-form, despite compositions that vary from sample to sample.They also note the observed fractionations are remarkably similar to the predicted fractionation on the basis of equi-librium isotope exchange at postdepositional metamorphic


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temperatures at Hemlo and that the compositions correlate with mass balance in a way suggestive of local reequilibration.A further future test of this hypothesis could be providedusing an approach similar to that of Jamieson et al. (2006), who relied on a combination of 33S and 34S to evaluate theapproach to a single-step equilibrium isotope effect. Never-theless, Thode et al. (1991) describe a relationship betweenthe isotopic composition of barite and pyrite versus modalpyrite-barite that is interpreted in terms of an initial baritecomposition (~10). In summary, Hemlo provides a rare window into the depositional environment of a deposit thathas been subjected to a rather complex metamorphic history (e.g., Muir, 2002; Heiligman et al., 2008).

Insights from sulfide ore deposits for the Proterozoic S-cycleInformation provided by Proterozoic VMS deposits is of a

different type than provided by Archean VMS deposits. Rel-atively fewer VMS deposits are observed in Proterozoic rocks,but this is not inferred to reflect a change in the oceans or inthe exogenic element cycles. Rather, it is inferred to reflectchanges in the geotectonic conditions conducive to the for-mation and preservation of VMS deposits. Where they occur,Proterozoic VMS are as informative about the nature of theoceanic environments as are Archean VMS, and they providedifferent information than that from CD Pb-Zn deposits andsedimentary geochemistry. For example, Slack et al. (2007)have argued that a lack of Ce anomalies and the presence of hematite in jasper associated with the deep-water (>1,000 m),late Paleoproterozoic (~1740 Ma) Jerome VMS deposit, Ari-zona, indicate the oceans were locally oxygenated in the sta-bility field of hematite and did not support the same globalMn oxide sink for Ce that is present today. They also infer, on

the basis of seven other similar deposits ranging in age from1792 to 1241 Ma, that these observations can be extrapolatedover a much broader time interval. Using these constraintsand log( O2)-pH (Eh-pH) diagrams (Fig. 11), it is inferredthat oxygen in the Paleoproterozoic is very low (anoxic), butstill variable over many orders of magnitude. The constraintson the presence of oxygen, therefore, do not appear to be re-lated to the stability of hematite, but to the mass balance con-straints for hematite formation. The absence of free sulfide in water overlying Jerome and similar deposits has been arguedon the basis of the absence of pyrite (Slack et al., 2007). How-ever, in at least one present-day, iron-rich setting (LakeMatano, Indonesia), iron (hydr)oxides have been shown to co-exist with very low levels of free sulfide (


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(e.g., Sangster, 1990; Lott, 1999; Cooke et al., 2000; Lyons etal., 2004). Leach et al. (2010) note that CD deposits in conti-nental rift and continental sag basins are only present in thePaleoproterozoic and Mesoproterozoic, whereas passive mar-gin-type CD Pb-Zn deposits are hosted in rocks from the Pa-leoproterozoic to the Phanerozoic. Models for the formationof these deposits invoke mixing of sulfidic and metalliferoussolutions, in many cases with links to the presence of sulfide,either in the sediments or in the overlying seawater (Good-fellow and Jonasson, 1984; Goodfellow, 1987; Shanks et al.,1987; Goodfellow et al., 1993). The sulfide from the McArthurbasin deposits has been inferred to originate either within thesediments through sulfate reduction or from bottom watersthat may have been sulfidic (Lyons et al., 2004). Both biolog-ical and thermochemical sulfate reduction are feasible inthese environments (Lyons et al., 2004; Huston et al., 2006).It is interesting to consider whether the isotopic signals pro-duced by thermochemical sulfate reduction (TSR) may pro-duce anomalous 33S. We know of one published analysis thatmay be attributable to TSR, which is for the Urquhart Shaleof the Mt Isa inlier (PPRG 1284) with a measured 33S withinerror of zero. The distribution of CD-Pb-Zn deposits may alsobe related to the presence of euxinic deep waters (e.g., Turner,1992; Goodfellow et al., 1993; Large et al., 1998; Canet et al.,2004; Lyons et al., 2004; Huston et al., 2006), although Slacket al. (2004) provide an alternative interpretation for the giantMississippian Red Dog deposit, Alaska, which is based on thelow Mo/Ti ratio in host shales.

The chemistry of the Earths middle Proterozoic (latest Pa-leoproterozoic, Mesoproterozoic, and early Neoproterozoic)oceans has been a topic of a debate centered on the questionof whether they were sulfidic in their entirety, or in part (Hol-land, 1984, 2005, 2006; Canfield, 1998). Holland (1984, 2006)suggested that ferrous iron was removed from the deep ocean

and replaced by oxygen at ~1.9 to 1.8 Ga because of greateroxygen availability in the aftermath of the Great OxidationEvent (GOE). Slack and Cannon (2009) suggested that theoceans changed to a nonsulfidic state with ~1 M dissolvedoxygen, because of the Sudbury impact, but it is not clear what feedbacks would have stabilized the system once thetransition took place. The situation that developed at this time would have been mainly the same as the modern, oxygenateddeep ocean, perhaps with slightly lower oxygen concentra-tions. Canfield (1998) proposed that a ferruginous deepocean was not replaced by an oxygenated deep ocean duringthe Proterozoic, but by a sulfidic deep ocean as a result of anexcess of sulfide over both iron and oxygen. According toCanfield (1998), the establishment of an oxygenated deep

ocean did not occur until significantly later in geologic history,likely closer to 600 Ma. The switch that triggered the transi-tion from ferrous to sulfidic was a change in sulfide supply asa result of higher sulfate delivery to the oceans. It was arguedthat if phosphorus availability was similar to that of today,then the sinks for oxygen in the deep ocean would exceed oxygen-supplied by circulation and ventilation of the deepoceans would not occur. Instead, anoxic oceans would result,and sulfate delivery and reduction would result in sulfidetitrating iron and replacing it as the redox control in the deepoceans. More recently, Holland (2006) argued that evidencefor phosphorous in Mesoproterozoic ocean sediments is

lacking and, therefore, there would not have been a largeenough flux of sinking organic matter to consume all oxygenand to support sulfide production sufficient to establish sul-fidic conditions.

Each of these suggestions explains some, but not all fea-tures of the sulfur record preserved by CD Pb-Zn depositsand oceanic sediments. Hollands (1984) suggestion does notexplain the evidence from deep sedimentary successions(e.g., Poulton et al., 2004) or from Mo concentrations andMo/TOC ratios (Arnold et al., 2004; Scott et al., 2008), or ev-idence for biomarkers from phototrophic sulfide oxidizingbacteria, which would imply sulfide in some surface waters(Brocks et al., 2004, 2005). It also doesnt explain the appar-ent missing pool or 34S-depleted sulfide evidenced by thepaucity of sedimentary sulfur with negative34S and a mean value for34S of both sulfide and sulfate that has a long-termenrichment relative to the long-term source of juvenile sulfur with near zero 34S (Logan et al., 1995; Lyons et al., 2004;Fig, 1). Canfields (1998) suggestion does not explain someobservations from ore deposits, such as the presence of ironoxide rather than iron sulfide in distal parts of exhalites (Slacket al., 2007), and nonsulfide Zn deposits that appear later inthe Proterozoic (Holland, 2005).

A hybrid model of the Proterozoic oceanA hybrid model has been proposed that attempts to recon-

cile these observations (Holland, 2005) and is rooted in theobservations from both sedimentary rocks and ore deposits.This model invokes a combination of sulfidic and nonsulfidicanoxic deep-water pools. There is some debate about how these pools may have been distributed in the oceans, and alsoas to how the oceans evolved following the reorganization at~1.9 Ga. This model calls for a situation in which sulfidic andnonsulfidic anoxic waters exist in different parts of the ocean.

Holland (2005) argues that it was not just the phosphorusavailability in the Proterozoic ocean that controlled the levelsof oxygen in downwelling waters, but also the flux of sinkingorganic matter that determines the strength of the deep- water oxygen sinks. He proposes that a slower rate of sinkingof organic matter from the surface ocean would lessen theoxygen sink and allow for a delicate balance to be met be-tween oxygen, sulfide, and ferrous iron that favored oxygen,but at very low levels. Levels were low enough such that Mnoxide was not present in high enough abundance on a globalocean scale to scavenge elements such as Mo and Ce. This hy-brid model that includes deep ocean sulfidic and nonsulfidicanoxic pools is the presently accepted model for the Protero-zoic oceans and is necessary to provide an explanation consis-

tent with the geologic and geochemical record (Table 2).The location of the sulfidic parts of the Proterozoic ocean would be related to sites where sulfide supply exceeded thesupply of ferrous iron and oxygen. Where the sulfate supply was nonlimiting, the supply of sulfide would be controlled by sulfate reduction rates and would depend on the availability of electron donors, principally organic carbon supplied fromthe upper parts of the Proterozoic water column. In present-day systems, sulfidic waters develop above sediments thatproduce more sulfide than they consume by pyrite formationand/or sulfide oxidation, and where additional production of sulfide in the water column extends the depth interval of


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sulfidic waters (e.g., Yakushev and Neretin, 1997; Srensenand Canfield, 2004). In todays environments, sulfidic watersextend up to a chemocline where they are removed by oxida-tion reactions and other Fe-O-S chemistry. In at least onepresent-day system, sulfidic waters have developed in the water column without a supply of sulfide from underlyingsediments. Crowe et al. (2008) documented a system in LakeMatano, Indonesia, in which oxic waters overlie sulfidic wa-ters that in turn overlie iron-rich waters. The controls on thissystem appear to be the delivery of organic carbon and theavailability of water column sulfate and iron. Both of these

systems provide insights into the distribution of sulfidic, oxy-genated, and ferrous waters in the Proterozoic oceans.The surface waters of the Proterozoic oceans are thought to

have been oxygenated as a result of oxygenic photosynthesis.Johnston et al. (2009) suggest that a sulfide-rich oxygen min-imum zone would have played a critical role in limiting oxy-gen production in the surface waters of the Proterozoic oceanand a stabilizing feedback that ultimately determined the at-mospheric oxidation state and the redox structure of theoceans. Oxygen limitation in oceanic shelf settings where thesupply of organic electron donors for sulfate reduction was high would lead to sulfide production. If the sulfide production

exceeded water column oxygen and iron supply, then euxinicconditions would have developed. This may have occurred asa result of sulfide released from sediments or from enhancedsulfate reduction in the water column such as that observedby Crowe et al. (2008). Support for this has been provided by Poulton et al. (2004), who provide evidence from iron speci-ation that points to the development of sulfidic waters in Pro-terozoic shelf environments. Lower sulfate reduction ratesare observed today in abyssal sediments (~0.121 mol/cm2 / yr) than in shelf or slope sediments (~40270 and ~ 1574 mol/cm2 /yr, respectively; Canfield et al. 2005). There is no

reason to suspect that a similar relative relationship betweensulfate reduction rates did not occur in the Proterozoic. Thismay have limited the supply of sulfide from Proterozoicabyssal sediments to abyssal waters, and it may have produceda situation in which the competition between sulfide anddeep sources of iron may not have favored euxinia, allowingfor formation of iron oxides without appreciable iron sulfidein a number of Proterozoic VMS deposits (Slack et al., 2007).It is not clear whether intermediate waters were sulfidic, suchas those that develop in modern settings like that of LakeMatano (Crowe et al., 2008), but this is a possibility that somefavor. We suggest that the Proterozoic oceans may have had a

524 FARQUHAR ET AL.

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TABLE 2. Observations Relevant to Proterozoic Ocean Chemistry

Type of evidence Age Implications for oceans Setting Reference

Evidence for oxygenated oceanic pools

Presence of hematite in the jasper and iron formation 1740 Ma, Oxygenated water Deep water 1, 2Jerome deposit Since 1.85 Ga (>1,000 m)

Cerium abundances in Proterozoic jasper from deep-water 1740 Ma Low oxygen, Whole ocean 1exhalites suggest limited formation of Mn and Fe-, Mn-oxides limited formationof Fe,- Mn-oxide

Mo isotope evidence suggests that the global ocean chemistry 1.8 Ga 1.0 Ga Low oxygen/ Whole ocean 3, 4did not preserve the same signal as seen in the oceans today Limited formationthat results from formation of Mn and Fe-, Mn-oxides of Fe,- Mn-oxide

Occurrence of non-sulfide zinc deposits such as the Franklin Late Low sulfide, possibly Below wave base 5and the Sterling Hill deposits of New Jersey Mesoproterozoic low total sulfur.

Evidence for sulfidic oceanic poolsLack of a barite cap on the Sullivan CD Pb-Zn deposit ~1430 Ma Low sulfate Below wave base 6

Evidence from Mo abundances and Mo-TOC suggest a ~1.8 Ga 551 Ma. Sulfidic waters or Whole ocean 7significant sulfidic sink for Mo other Sulfidic sinks

Presence of pigment biomarkers from two classes of ~1.64 Ga Sulfidic waters Photic zone/ 8, 9, 10phototrophic sulfide oxidizing bacteria (purple and green possible restrictedsulfur bacteria) in McArthur basin sediments marine basin

Change in iron speciation in the Gunflint Iron Formation and ~1.85 Ga Transition from ferrous Shelf below 11the overlying Rove Fm. to sulfidic waters wave base

Models for formation of CD Pb-Zn deposits in the McArthur ~1.6 Ga Sulfidic bottom waters Below wave 6, 12basin that include mixing of deep (oxidized brines) with base, possiblesulfidic deep ocean-derived shallow fluids restricted basin

Long-term, positive 34S of Mesoproterozoic sulfate and sulfide 1.8 Ga 1.0 Ga Missing sulfidic sink Whole ocean Figures 1,(average value) from sedimentary rocks and ore deposits implies sulfidic waters 810 andsuggests a missing negative34S pool since the ultimate and sediments in see 12, 13,sources of sulfur to the surface sulfur cycle (input of juvenile deep water 14, 15sulfur) have34S values close to zero

(1) Slack et al. (2007), (2) Slack and Cannon (2009), (3) Anbar and Knoll (2002), (4) Arnold et al. (2004), (5) Holland (2005), (6) Cooke et al. (2000), (7)Scott et al. (2008), (8) Brocks et al. (2004), (9) Brocks et al. (2005), (10) Brocks and Schaeffer (2008), (11) Poulton et al. (2004), (12) Lyons et al. (2004), (13)Logan et al. (1995), (14) Canfield (2004), (15) Johnston et al. (2007)


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complex structure with respect to sulfidic, oxygenated, andferruginous waters. Whereas surface waters were oxygenated, water over the shelves, the intermediate-water, and perhaps anumber of deep basins may have developed euxinic condi-tions, and other deep-water settings may have been anoxic,nonsulfidic, or even ferruginous.

Some features of the record preserved by ore depositsand sedimentary geochemistry suggest an even more com-plex picture. These include the following: (1) a suggestionby Johnston et al. (2005a) that the oxidation pathways forsulfide in the Proterozoic oceans underwent a change in the 33S of sulfate that started at ~1.4 Ga, (2) the appearanceand disappearance of nonsulfide zinc deposits (Holland,2005), (3) a drop in the frequency of CD Pb-Zn deposits inthis interval (Lyons et al., 2004), (4) the absence of signifi-cant evidence for the presence of euxinic conditions be-tween ~1.5 Ga and 750 Ma (Condie et al., 2001; Meyer andKump, 2008; Scott et al., 2008), and (5) the suggestion by Canfield et al. (2008), on the basis of iron speciation, thatanoxic but ferruginous water replaced sulfidic waters in theNeoproterozoic because of enhanced iron delivery or lim-ited sulfur. These changes suggest something happened with the sulfur chemistry of the oceans in the late Protero-zoic. One explanation is that they reflect the limitation of sulfur on a global scale brought about by a shrinking exo-genic sulfur pool, a suggestion made by Canfield (2004) onthe grounds that sulfide deposited in abyssal sediments may have been lost from the system to subduction. Another ex-planation by Holland (2005) attributes these changes tohigher fluxes of sinking organic matter, exceeding the ven-tilation (oxygenation) of the deep ocean that left Fe2+ in ex-cess. It is not clear if they can be linked to a single overriding

cause, but it would appear that further studies of ore de-posits and sedimentary sulfur in combination with biogeo-chemical models are needed to provide an answer to why these changes have been suggested.

SynthesisEvidence from ore deposits and the sedimentary rock

record yield several important observations on the connec-tions between the sulfur cycle and the secular evolution of ocean chemistry. The most significant of these are summa-rized below and are used to develop an evolutionary modelfor the principal features of the oceans (Fig. 12).

The presence of Superior-type banded iron formations(Cloud, 1973) is evidence for significant solubility of ferrousiron (Fe2+) in the oceans that limited oxygen and free sulfide.On the basis of trace metal and sulfur isotope data, it has beensuggested that local, or spatially constrained oxygenated(Wille et al., 2006; Anbar et al., 2007; Kaufman et al., 2007),or sulfidic pools (Reinhard et al., 2009; Poulton et al., 2009)developed in the Neoarchean ocean that reflected the ap-pearance of oxygenic photosynthesizers. These suggestionsindicate that the oceans of early Earth evolved from a Pale-oarchean and Mesoarchean state with ferruginous deep wa-ters and relatively small amounts of cycling to a Neoarcheanstate with ferruginous deep waters and surface waters withproduction of oxygen and local competition between iron,oxygen, and sulfur. Nevertheless, oceanic sulfate levels likely remained below 200 M during the entire Archean (Habichtet al., 2002), and there is no conclusive evidence for spikes insulfate concentration. This is consistent with the lack of sig-nificant variability for34S and the preservation of MIF-S inthe sedimentary record.


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Sulfidic

Ferruginous F e r r u g

i n o u s

Oxic

Oxic

D e e p

O c e a n

C h e m

i s t r y

A t m o s p

h e r e

SurfaceOcean

Oxic Anoxiclow high

Ferruginous

O c e a n

i c

s u l f a t e

c o n c e n

t r a

t i o n

Time (Ga)

3.8 2.8 2.42 1.8 1.4 0.8 0.55 0.3 Today

?

gypsum before halitehalite before gypsum

bioturbation

OAEs

nonsulfidicanoxic

FIG. 12. Inferred evolution of atmosphere, surface ocean, and deep ocean and oceanic sulfate concentration. Note thatthe representation does not imply depth or geographic distribution of pools in the deep ocean.


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Whereas some earlier studies of ore deposits have arguedfor high levels of sulfate in the Paleoarchean (e.g., Vearn-combe et al., 1995), this may not be correct (Huston andLogan, 2004). It is possible, however, that the Neoarcheanmay have been characterized by a drop in seawater sulfateconcentrations because of the growth of sinks associated with microbial sulfate reduction. This is consistent with arise in the 34S of the oceanic sulfate pool and possibly ex-plains the lack of barite in Neoarchean VMS deposits. It alsomay have contributed to development of euxinia in somebasins where sulfate reduction rates were enhanced, and if the concentrations of sulfate exceeded the local supply of ferrous iron.

The rise in oxygen levels at ~2.42 Ga was accompanied by a rise in seawater sulfate concentrations, the shutting downof MIF-S processes, an increase in the variability of34S,and a relationship between 34S, MIF-S, and 13C. ThisGOE was followed by a protracted change in the chemistry of iron in the oceans, indicated by 56Fe and Mo data(Rouxel et al., 2005; Johnson et al., 2008b; Scott et al.,2008). Other observations include a change in the sequenceof precipitation of gypsum and halite in evaporites(Grotzinger and Kasting, 1993; Pope and Grotzinger, 2003)that suggest sulfate and calcium concentrations exceeded asaturation threshold of ~23 mM 2 at about 1.8 Ga. Quartznodules with palmate upright shapes and chevron termina-tions consistent with a selenite gypsum precursor havebeen reported in 2.1 Ga evaporites of the Lucknow For-mation, South Africa (Schrder et al., 2008, p. 110), indicat-ing that gypsum precipitation before halite occurred at aneven earlier date. There are other reports of higher sulfatelevels at 2.4 to 1.8 Ga, including evidence for gypsum pre-cipitation (e.g., Mehlezhik et al., 2005). This time alsomarked the appearance of the most ancient CD Pb-Zn de-

posits, MVT deposits and sediment-hosted copper deposits, which all have been related to rising sulfate levels in near-surface environments.

A change in ocean chemistry occurred at ~1.9 to 1.8 Ga, when sulfidic and/or anoxic nonsulfidic, nonferruginousocean pools were established. The balance between sulfidicand nonsulfidic anoxic pools in the 1.8 Ga to 800 Ma interval was a balanced system that responded by changes in the pro-portion of sulfidic to nonsulfidic anoxic pools.

Stabilization of sulfidic pools was favored in shelf settings where sulfate delivery rates and sulfate reduction rates werehigh, and also in intermediate waters below the photic zone where delivery of organic donors was high (Crowe et al.,2008). Development of nonsulfidic low oxygen pools oc-

curred where ventilation of the deep water exceeded the oxy-gen sinks from settling organic matter and also from sourcesof sulfide and ferrous iron. Evidence from CD Pb-Zn de-posits in continental rifts and sag basins and from VMS de-posits, as well as from sedimentary iron speciation, supports atransition to an ocean with sulfidic as well as nonsulfidic andanoxic (Slack et al., 2007; Slack and Cannon, 2009) interme-diate and deep water at this time.

The second half of the Neoproterozoic, and possibly theentire Neoproterozoic, was dominated by ferruginous oceanbottom water in shelf settings (Canfield et al., 2008). The per-sistence of some sulfidic pools is indicated by Mo data (Scott

et al., 2008), and the existence of shallow-water euxinia is in-dicated by a biomarker (Olcott et al., 2005). The reasons forthe re-establishment of iron-rich conditions are debated, butinclude hypotheses related to the size of the exogenic sulfurpool (Canfield, 2004; Canfield et al., 2008) or the change inthe supply of hydrothermal ferrous iron.

Starting in the latest Neoproterozoic (~600 Ma), a secondoxidation of the Earths surface is indicated by the coloniza-tion of the oceans by animal life (e.g., Logan et al., 1995). Ithas also been suggested this change corresponded to a rise of sulfate above Proterozoic levels that is attributable to theonset of bioturbation (Canfield and Farquhar, 2009). Evi-dence for this rise is indirectly tied to a change in the range of variation for34S. Direct evidence for episodes

connections between sulfur cycle evolution, sulfur isotopes, sediments,and base metal sulfide...

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