stephen j. mojzsis - mojzsis research group, cu...

Precambrian Ophiolites and Related RocksEdited by Martin J. van Kranendonk, R. Hugh Smithies and Vickie C. BennettDevelopments in Precambrian Geology, Vol. 15 (K.C. Condie, Series Editor) 1© 2007 Elsevier B.V. All rights reserved.DOI: 10.1016/S0166-2635(07)15075-1

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Chapter 7.5

SULPHUR ON THE EARLY EARTH

STEPHEN J. MOJZSIS

Department of Geological Sciences, Center for Astrobiology, University of Colorado,2200 Colorado Avenue, Boulder, CO 80309-0399, USA

7.5-1. INTRODUCTION

On the young Earth, the surface zone of the planet was shaped by physical and chem-ical interactions between the crust, hydrosphere, atmosphere and the emergent biosphere,on the geochemical cycles of the biologically important elements (H, C, N, O, P, S, Feand others). An explicit record of the early co-evolution of the atmosphere and biospherewith the geosphere has been notoriously difficult to investigate. By its very nature theatmosphere is ephemeral, and the result has been that a direct record of its chemical evo-lution remained frustratingly out of reach. It is now reasonably well understood that thegeochemical behaviour of the multiple sulphur isotopes are the best tool for investigationsof the long term evolution of the surface system, and that they provide a uniquely powerfulproxy for long-term changes in atmospheric chemistry and other key processes (Thiemens,1999, 2006). The geochemistry of sulphur is complex. There is a long history of study onthe topic so that a single review on the generalities of sulphur biogeochemistry on the earlyEarth cannot hope to cover all progress in this rapidly changing field (please see Ohmotoand Goldhaber (1997), Canfield and Raiswell (1999), Canfield (2001), Farquhar and Wing(2003) and Seal (2006) for comprehensive reviews). Instead, the goal here will be to out-line several key aspects of sulphur that are of particular interest to studies of Earth’s oldestrocks (limited to before ca. 3.5 Ga), and to make the case that a comprehensive programof retrospective multiple sulphur isotope analyses is required on samples for which prior34S/32S data have been compiled (Strauss, 2003).

When one attempts to review processes operative on the early Earth, global environ-mental aspects unique to the young planet need to be acknowledged. In the first billionyears, higher crustal heat flow and enhanced mantle outgassing, widespread (and rapid?)crustal recycling and considerable changes in continental volume, and the emergence oflife and its subsequent take-over of several important aspects of chemistry of the surfacezone, were all likely governing factors that influenced the various sub-cycles of the globalsulphur cycle schematically represented in Fig. 7.5-1. Furthermore, exogenous processesunique to the young planet were also important, even if they are more difficult to accountfor in physical models. These processes included a higher frequency of asteroid and comet

dpg15 v.2007/05/23 Prn:4/06/2007; 13:56 F:dpg15035.tex; VTEX/JOL p. 1aid: 15075 pii: S0166-2635(07)15075-1 docsubty: REV

http://dx.doi.org/10.1016/S0166-2635(07)15075-1

2 Chapter 7.5: Sulphur on the Early Earth

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Fig. 7.5-1. Schematic representation of global geochemical cycles of sulphur and the averagesulphur isotope compositions of important reservoirs in the Earth system. The interface betweenthe geosphere and hydrosphere is dominated by exchange of sulphur between seawater and rocksby fluids, weathering, dissolution and re-precipitation. Sulphide will precipitate with an ade-quate supply of dissolved iron (and other metals) and can be abiogenic or biogenic. Sulphurmay also be sequestered in evaporates, brine pools or crustal fluids as sulphate, sulphide or ele-mental sulphur. Chemical weathering is the principal mechanism that returns sulphur from rocks andsediments on land to the hydrosphere. Hydrothermal activity and subduction transports sulphur backto the geosphere. Submarine dissolution of basalt and other rocks also returns S to the hydrosphere.


7.5-1. Introduction 3

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Fig. 7.5-1. (Continued.) Volcanism injects sulphur gases to the hydrosphere and atmosphere. Vol-canic gases in the atmosphere can become transformed by light-driven reactions and returned tosediments from aerosol deposition. Subduction will bring all components back to the mantle. Ex-traterrestrial input is minimal, but could be important in a few cases of large impacts.

impacts, an enhanced rate of extraterrestrial dust delivery compared to present, as well asreduced solar luminosity (but higher ultraviolet flux) from the young Sun.

However, there exists a further and yet more fundamental limitation to what can beknown about the earliest Earth. A geologic record of the interactions outlined in Fig. 7.5-1has apparently been lost for the first ca. 750 My due to the effective recycling of all crustexcept for remnant ca. 4.4–3.8 Ga zircons contained in younger sediments (see Cavosieet al., this volume). Sedimentary rocks �3.7 Ga in age, which could potentially preservea direct sample of the ancient sulphur cycle, occur only as rare and deformed enclavesin high-grade Archaean metamorphic granitoid-gneiss terranes. Nevertheless, given recentadvances in instrumental techniques, experimental and observational geomicrobiology andmolecular biochemistry, chemical oceanography, and access to an increasingly large sam-ple base of ancient rocks, our knowledge of early Earth environments has accumulated tothe point where it is now perhaps worthwhile to revisit what we know of the interactionsbetween the geochemical cycles of sulphur at that time. Specifically:(1) Sulphur comes in four stable isotopes of well-defined relative abundance. What is the

origin of sulphur to the Earth? How do changes occur to the abundance ratios of thesulphur isotopes? What is the significance of mass-independent isotope fractionationas opposed to ‘normal’ mass-dependent effects in the early Earth record?

(2) The infall of meteoritic and cometary debris to the planets was greater in the first billionyears than at present. What contributions could extraterrestrial sources have made tothe global geochemical sulphur cycle in the first billion years? Could a trace of thisextraterrestrial material remain in the oldest sedimentary rocks?

(3) Subsequent to primary outgassing, the atmosphere underwent chemical changes. Whatclues can sulphur provide to the composition and chemical evolution of the atmosphereprior to about 3.5 Ga?

(4) Sulphur compounds are central to many fundamental biochemical processes. Whatrole could sulphur have played in the establishment of the biosphere? Could sulphurisotope fractionations serve as a robust isotopic biomarker in highly metamorphosedrocks?

(5) All rocks older than about 3.5 Ga have been strongly transformed by metamorphism.How well is the sulphur isotope record preserved in the oldest rocks?

(6) Banded iron-formations (BIFs) are the most common sedimentary rock preserved fromthe early Archaean, but their derivation remains enigmatic. What can sulphur isotopesinform us of the origins of the BIFs?



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7.5-2. ORIGIN OF SULPHUR TO THE EARTH

The elemental inventory of the planets was determined by the composition of the neb-ular gas and dust from which the solar system formed. In stars greater than 8 times solarmass, and at temperatures near 1.5 × 109 K, oxygen fusion begins as a result of the slow(s-process) fusion of neutrons (n), protons (p) and 4He nuclei (α) to a variety of seed nuclei(Cameron, 1979; Woosley and Weaver, 1986). In this scenario, the various sulphur isotopescan be produced from the following important oxygen-burning nuclear reactions:

16O(16O, γ

) → 32S; 32S(4He, γ

) → 36S; 32S(n, γ ) → 33S; 33S(n, γ ) → 34S

Here, it can be seen how the major oxygen-oxygen reaction 16O(16O, γ ) → 32S yieldsthe abundant 32S isotope and energy. Neutrons evolved from side reactions such as16O(16O, n) → 31S can result in further s-processes and the formation of the minor isotopes(Seeger et al., 1965). As a result of these processes, there are four stable isotopes of sulphurwhich have the following relative abundances as defined by the meteorite sulphur standardCañon Diablo troilite (Hultson and Thode, 1965; see Ding et al., 2001 and Beaudoin et al.,1994):

32S = 95.02%; 33S = 0.75%; 34S = 4.21%; 36S = 0.02%

The high relative abundance of 32S can therefore be understood from the fact that it isdirectly synthesized from O-burning reactions. Sulphur is an exceptionally stable element.It has an even atomic number (Z = 16) and the nuclide 32S has both a mass number (A)

and a neutron number (N) that are multiples of 4 (the mass of the He nucleus), so thatA = Z + N = 16 + 16 + 32. Since the minor S isotopes are more closely tied to ratesand types of various s-process reactions which also include synthesis from 4He-fusion withthe stable isotopes of Si (28Si, 29Si and 30Si), the relative abundances of 33S, 34S and 36Sare much lower. This inventory of the different sulphur isotopes was imparted to the solarnebula at some time prior to the accretion of the planets.

7.5-2.1. Primordial Sulphur

The formation time of Earth and other planets, and the elemental inventory of Earth, hasbeen deduced in part on the basis of comparison with primitive meteorites (type 1 carbona-ceous chondrites; CI) and the solar photosphere. The average compositions of CI mete-orites and the Sun have been compiled by numerous workers (e.g., Anders and Grevesse,1989), who showed that CI:solar photosphere displays a nearly 1:1 correspondence in sul-phur content normalized to Si, so that bulk meteoritic sulphur is probably a reasonablygood approximation for bulk planetary sulphur isotopic compositions. Primitive carbona-ceous chondrites contain approximately 5.4 wt% sulphur (Dreibus et al., 1995) and S is thetenth most abundant element in the solar system (Anders and Ebihara, 1982). Sulphur ranksas eighth in abundance for the whole Earth (silicates + core; Morgan and Anders, 1980)and the Bulk Silicate Earth concentration is about 250 ppm (McDonough and Sun, 1995).


7.5-2. Origin of Sulphur to the Earth 5

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A summary of the various concentrations of sulphur and sulphur-containing compounds inthe terrestrial reservoirs in the atmosphere, hydrosphere, biosphere and geosphere is pro-vided in Table 7.5-1. The sulphur isotopic composition of the bulk Earth is assumed to beclose to that of the major S-phases of iron meteorites. Because the record of sulphur isgeochemical, a brief review of the stable isotope geochemistry of sulphur is warranted.

7.5-2.2. Sulphur Isotope Fractionations

For low atomic number elements like sulphur, the mass differences between the various iso-topes are large enough for many physical, chemical, and biological processes or reactionsto change the relative proportions of the isotopes (reviewed in Hoefs (1997)). In normalphysical-chemical interactions on Earth, two isotope effects – equilibrium and kinetics –result in isotope fractionation. Due to these fractionation processes, products often developunique isotopic compositions (normally expressed as ratios of heavy to light isotopes) thatmay be indicative of their source, or of the processes that formed them (Faure, 1986, andreferences therein). Reaction rates depend on the ratios of the masses of the isotopes andtheir zero-point energies. As a general rule, bonds between the lighter isotopes are severalcalories less than bonds between heavy isotopes. The lighter isotopes tend to react morereadily and become concentrated in the products, and the residual reactants tend to becomeenriched in the heavy isotopes. The magnitude of the fractionation strongly depends on thereaction pathway utilized and the relative energies of the bonds broken and formed by thereaction. Slower reactions (for instance those proceeding at lower temperatures) tend toshow larger isotopic fractionation effects than faster ones.

Equilibrium isotope-exchange reactions involve the redistribution of isotopes amongvarious chemical species or compounds. During equilibrium reactions, the heavier iso-topes preferentially accumulate in the compound with the higher energy state. Kineticisotope fractionations occur when systems are out of isotopic equilibrium, such that for-ward and back reaction rates are not identical. When reactions are unidirectional, reactionproducts become physically isolated from the reactants. Biological processes are kineticisotope reactions that tend to be unidirectional. Both diffusion and metabolic discrimina-tion against the heavier isotopic species occurs because of the differences in zero-pointenergies of the different isotopes. This can result in significant fractionations between thesubstrate (isotopically heavier) and the biologically mediated product (isotopically lighter).Many reactions can take place either under purely equilibrium conditions or be affected byan additional kinetic isotope fractionation, and metabolic (oxido-reduction) reactions willpreferentially select molecules with lighter isotopes (Schidlowski et al., 1983).

To follow the various isotopic fractionation pathways, partitioning of stable isotopes be-tween two substances A and B is expressed by use of the isotopic fractionation factor (α):

α(A − B) = RA/RB

where R is the ratio of the heavy to light isotope (e.g., 34S/32S or 33S/32S). Values for α

tend to be close to 1.00 so that differences in isotopic compositions are usually expressed inparts-per-thousand, or ‘per mil’ (�). Kinetic fractionation factors are typically described


6C

hapter7.5:

Sulphuron

theE

arlyE

arth

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Table 7.5-1. Average concentration of sulphur in various reservoirs on Earth (values are in 10−6 g/g unless otherwise indicated)

Reservoir Reduced SH2S

Elemental S

S2 (S0, S8)

Oxidized SSO2 [SO]

Sulphate

SO42−

OtherOCS

TotalS

Atmospherea 3–10 × 10−5 ∼1.5 × 10−5 2–9 × 10−5b – 5 × 10−4 <1 × 10−3

Biosphere – – – – – 1000c

HydrosphereRiversd – – – 11.5 – –Oceans – – – 2784e – ∼2800f

Shale – – – – – 2400g

Granite – – – – – 260h

Mafic rocks andUltramafics

– – – – – ∼1000–4000i

Volcanic gasesj 0.057–3.21 3 × 10−4–1.89 0.006–47.7 – 2 × 10−5–0.08[0.03–0.06]

Comet particles 95480k

IDPs 41664l

Notes:apresent troposphere (Warneck et al., 1988); bas CS2; cdry weight of tissue (Berner and Berner, 1996 – citing Zinke, 1977); Schlesinger, 1997); dStumm and

Morgan (1996); eMillero and Sohn (1992); fLi (1991); gTurekian and Wedepohl (1961); hMason and Moore (1982); iWedepohl and Muramatsu (1979);jSymonds et al. (1994) values expressed in mol%; kKissel et al. (1986); lSchramm et al. (1988).

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in terms of various enrichment or discrimination factors and these are defined in variousways by different researchers (see O’Neil, 1986).

7.5-2.3. Mass-Dependent Fractionation of Sulphur Isotopes

The numerous paths of sulphur between the different reservoirs shown in Fig. 7.5-1 of-ten involve changes in oxidation states that are modulated through both abiotic and bioticequilibrium and kinetic reaction processes, and the different degrees of S-isotope fraction-ations depend on these modes of isotopic exchange. In order to track these fractionationsin nature, sulphur isotope compositions of substances are expressed with the conventionalδ notation as:

δ34SCDT(�) = (R34

sample/R34CDT − 1

) × 1000

δ33SCDT(�) = (R33

sample/R33CDT − 1

) × 1000,

where R34 = 34S/32S, R33 = 33S/32S and the reference standard is the pyrrhotite end-member phase troilite (FeS) from the Cañon Diablo meteorite (CDT; see Beaudoin et al.,1994). Similarly, this expression holds for δ36S, however the value has not usually beenreported because the trace abundance of 36S makes the measurement difficult with nor-mal mass-spectrometric techniques (e.g., Hulston and Thode, 1965; Hoering, 1989; Gaoand Thiemens, 1989). As reviewed elsewhere (Mojzsis et al., 2003; Farquhar and Wing,2003), most conventional sulphur isotope studies omitted the minor isotope 33S with theview that there would be no new information conveyed by the measurement of 33S/32S(cf. Hoering, 1989; Lundberg et al., 1994). This is because normal physical, chemical,or biological processes fractionate isotopes with predictable mass-dependent behaviour(Urey, 1947; Bigeleisen and Mayer, 1947). In other words, on sulphur three-isotope plotsof [(33S/32S)/(34S/32S)] or [(36S/32S)/(34S/32S)], the relative mass differences of the iso-topomers are such that variations in δ33S and δ33S ought to be correlated with those in δ34S.Mass-dependent fractionation laws arise from the kinetic and equilibrium reactions thatfractionate sulphur in approximate proportion to their relative mass difference. Becausethe mass difference between 33S and 32S is about half of that between 34S and 32S, anymass-dependent isotope effects in δ33S would be about half those in δ34S. Mass-dependentfractionation (MDF) processes lead to isotopic compositions that can be described by theequation:

(33S/32S)/(33S/32S

)CDT = [(34S/32S

)/(34S/32S)

CDT

]λ.

In this treatment, linear regression in (x, y) with y = ln(33S/32S) and x = ln[(34S/32S)/

(34S/32S)CDT] has slope λ. Theoretical analyses of equilibrium MDF laws define slopesof λ ≈ 0.515 (Bigeleisen and Mayer, 1947; Hulston and Thode, 1965). York regressionsthrough large datasets from sulphur standards measurements yield values λ = 0.515 ±0.005 (Farquhar et al., 2000a, 2000b, 2000c; Mojzsis et al., 2003; Farquhar and Wing,2003; Ono et al., 2003; Hu et al., 2003; Papineau et al., 2005, 2006; Papineau and Mojzsis,2006; Cates and Mojzsis, 2006; Whitehouse et al., 2005; Ono et al., 2006, 2007; Papineau



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et al., 2007). Empirical determinations of 33S/32S are obtained by fixing λ at 0.515 with(34S/32S)V-CDT = (22.6436)−1 (Ding et al., 2001), so that regressions through standardsdata with fixed λ result in intercepts that equal ln[(33S/32S)V-CDT]. This yields (33S/32S)CDTvalues = (126.3434)−1, and with the thousands of standards measurements reported thusfar, mass-dependent values have been narrowly defined about [(34S/32S)/(33S/32S)]CDT =0 ± 0.1�(2σ).

Different modes of mass-dependent fractionations can arise from different kinetic andequilibrium effects (Bigeleisen and Wolfsberg, 1958). What is little-discussed but germaneto sulphur isotope geochemistry is that the scatter often seen in [(34S/32S)/(33S/32S)] datacan deviate slightly from orthodox mass-dependency because different slopes in λ mayresult from these different mass-dependent modes (e.g., Young et al., 2002) captured inthe compositions of natural standards. In light of this observation, Papineau et al. (2005)argued that MDF trajectories on a three-isotope plot should be projected as a band thatencompasses the different mass-dependent laws, rather than as a simple ‘terrestrial (mass-)fractionation line’. Based on this analysis, the width of the York-MDF band is defined asthe standard deviations in y about the fixed-slope York regression through all standards in(x, y)-space with a slope λ = 0.515 ± 0.005 when transformed to δ34SCDT vs δ33SCDTspace. This scrutiny of the various modes of mass-dependent sulphur isotope fractionationsin the three-isotope system yields external errors in the range of ±0.20�. Other, moredetailed studies that explore the quadruple sulphur isotope system at high resolution withlaser fluorination methods is reviewed in Ono et al. (2003, 2006, 2007).

Admittedly, there remains considerable room for improvement in these measurementsand it is expected that when synthetic sulphur standards become available, the externalerrors will decrease and finer details of sulphur isotope fractionations will be revealed.Theoretical treatments of the MDF boundaries within the range of possible MDF typesshow that they plot as shallow curves, and the York MDF-bandwidth (approximately±0.1�, 2σ in the theoretically best possible case) increases slightly with absolute valuesof (δ33S/δ34S)CDT (Fig. 7.5-2). However, some unusual photolytic atmospheric reactionsthat occur in the gas-phase, and some nuclear effects in space, produce sulphur that signif-icantly deviates from mass-dependent fractionation laws in what has come to be known as‘mass-independent fractionation’.

7.5-2.4. Mass-Independent Fractionation of Sulphur Isotopes

In order to better grasp the significance of mass-independent sulphur isotope effects cap-tured in the geologic record of the early Earth, it is worthwhile to briefly summarise theinitial discovery of this phenomenon in the oxygen isotopes. Mass-independent fraction-ation (MIF) behaviour of stable isotopes was found in the oxygen isotopic compositions(16O, 17O, 18O) of carbonaceous meteorites (Clayton et al., 1973). In extensive work since,Clayton and others have shown that on a three-isotope oxygen plot1, anhydrous miner-als in carbonaceous meteorites lie along a line of slope of +1 (Fig. 7.5-3) instead of a

1(δ17O vs δ18O; where δ1xO = ((1xO/16O)sample/(1xO/16O)V-SMOW − 1) × 1000�, x is either 7 or 8 and

V-SMOW is the Vienna standard mean ocean water).



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Fig. 7.5-2. Three-isotope plot (δ33S vs δ34S) for various sulphide standards that are used to definethe York-MDF bandwidth. For clarity, the error bars for the individual measurements have beenomitted (from Papineau et al., 2005).

slope of approximately +0.5 expected to define normal mass-dependent isotope effects onEarth. The initial interpretation that nucleosynthetic processes produced mass-independentisotopic signatures was re-visited by Thiemens and Heidenreich (1983) who showed ex-perimentally that the origin of MIF oxygen isotopes was likely photochemical (Fig. 7.5-4).Photolysis of diatomic oxygen under ultraviolet radiation produces MIF oxygen isotopesin the residual molecular oxygen (O2) and in the product ozone (O3). Fig. 7.5-4 definesa theoretical mass fractionation (band), which is the regime wherein all O-isotope MDFprocesses plot according to the equation:

�17O = 1000 ×{(

δ17O

1000+ 1

)−

(δ18O

1000+ 1

)λ}.

In this case, �17O is the permil deviation from mass-dependent fractionation and λ repre-sents the slope, which expresses the mass-dependent fractionation relationship betweenδ17O and δ18O (approximately δ17O = 0.52 × δ18O). Mass-dependent fractionationprocesses result in isotopic compositions that have a �17O of 0�, while mass-independentprocesses that arise from ultraviolet photolysis will result in �17O �= 0�, or significantdeviation from the mass-dependent fractionation band. In Fig. 7.5-4, product ozone (O3)from the photolysis of molecular oxygen (O2) has a positive �17O, which means that it issignificantly more enriched in 17O than it would if the fractionation process was purely



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Fig. 7.5-3. Three-isotope plot (δ17O vs δ18O) of anhydrous minerals in carbonaceous meteoritesshowing MIF oxygen isotopes (from Clayton et al., 1973).

mass-dependent. The isotopic composition of residual O2 also shows MIF in isotopicmass-balance and has a negative �17O, which is a depletion of 17O relative to isotopic com-positions characteristic of mass-dependent processes. The slope of the data in Fig. 7.5-4 isclose to +1, while the slope for MDF processes is about +0.5. These results convincinglydemonstrated that UV photochemical reactions can lead to products and residual reac-tants that exhibit mass-independent fractionation. Note that the range of δ18O and δ17Oin Fig. 7.5-4 for the photochemical experiments is different than that observed in car-bonaceous meteorites (Fig. 7.5-3). This could mean that photochemical reactions are moreefficient in a controlled experiment and/or that isotopic dilution occurred from mixingmass-independently fractionated components within dominantly mass-dependently frac-tionated oxygen reservoirs. Further progress in our understanding is hampered by the factthat the actual mechanism(s) for mass-independent fractionation have yet to be fully elu-cidated with quantum-chemical theory (Thiemens, 1999; Mauersberger et al., 1999; Gaoand Marcus, 2001; Deines, 2003).



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Fig. 7.5-4. Three-isotope plot (δ17O vs δ18O) for oxygen isotopes in product O3 and residual O2after irradiation of O2 with ultraviolet light from electrical discharge (from Thiemens, 1999).

Sulphur is isoelectronic with oxygen2 and has also been found to display mass-independent isotope fractionation. Multiple sulphur isotope data are reported using theconventional notation as:

�33S = 1000 ×{(

δ33S

1000+ 1

)−

(δ34S

1000+ 1

)λ}

�36S = 1000 ×{(

δ34S

1000+ 1

)−

(δ36S

1000+ 1

)λ},

where �33S and �36S are the permil deviations from mass-dependent isotopic fraction-ation processes, δ3xS = ((3xS/32S)sample/(

3xS/32S)CDT − 1) × 1000� where x = 3, 4or 6, and λ represents the mass-dependent fractionation relationship between δ33S andδ34S (δ33S = 0.52 × δ34S) or between δ36S and δ34S (δ36S = 1.9 × δ34S). Like oxygen

2Only mass-dependent fractionations have been reported for Selenium (Johnson and Bullen, 2004). Telluriumdoes not appear to display MIF, although searches have been conducted (Fehr et al., 2005).



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Fig. 7.5-5. Three isotope plots (δ33S vs δ34S) showing (a) experimental results with an ArF excimerlaser on SO2 gas and (b) Precambrian sedimentary sulphide and sulphate multiple sulphur isotopecompositions (from Farquhar et al., 2001).

�17O, sulphur �33S and �36S values ∼0� result from MDF processes such as kinetic,equilibrium and oxido-reduction reactions. Laboratory experiments have shown that mass-independently fractionated sulphur isotopes can arise in gas-phase reactions with startingmixtures of SO2, H2S and CS2 (Zmolek et al., 1999; Farquhar et al., 2000a, 2000b, 2001).Results from photolysis experiments with an ArF excimer UV laser at 193 nm in a SO2

gas are shown in Fig. 7.5-5(a). These experiments documented MIF sulphur isotopes inproduct S0, SO4

2− and residual SO2. The �33S values of these S-phases are positive forproduct elemental sulphur so that they plot above the mass-dependent fractionation band,and negative for product sulphate (i.e., they plot below the MDF band).

The MIF behaviour of three sulphur isotopes (32S, 33S and 34S) in photochemical ex-periments resembles multiple sulphur isotopes from Precambrian sedimentary sulphidesand sulphates as seen in Figs. 7.5-5(b) and 7.5-6. Similar to the results for oxygen iso-topes, the main difference between the plots is that the range of δ33S and δ34S values islarger in the photochemical experiments (tens of permil) and much smaller (a few permil)for natural samples that exhibit mass-independent behaviour such as found in pre-2.45 Gasedimentary sulphides and sulphates. As for the oxygen case discussed above, this may



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Fig. 7.5-5. (Continued.)

arise from more efficient photochemical reactions in controlled experiments and/or fromisotopic dilution occurring in natural systems.

Barring further detailed experimental verification from a wider range of UV wave-lengths (Lyons, 2006), data from ancient sulphur are consistent with the view that shortwavelength solar radiation (i.e., λ < 220 nm) penetrated deeply in the early terrestrialatmosphere in the absence of an effective ozone screen to lead to significant MIF in thegeologic record (Farquhar et al., 2001). It is interesting to note that the higher extreme UV(EUV) output observed for young solar-type stars (Wood et al., 2002) requires that in theearly Earth’s atmosphere the incident solar EUV flux fell from ∼30× to approximately 4×present between 4.4 to 3.5 Ga (Zahnle, 2006). The evolution in the incident UV flux hasyet to be accounted for in our models for the generation of mass-independent fractionationof sulphur isotopes in the ancient atmosphere. Mass-independently fractionated signaturesin early Precambrian sedimentary sulphides and sulphates could be coupled to results ofexperiments that explore the magnitude of MIF under specific atmospheric regimes andUV intensities at different wavelengths. These data could then be used to trace the chemi-cal evolution of the atmosphere in relation to the long-term interaction of the surface zonewith the evolving young Sun.



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Fig. 7.5-6. Plot of �33S vs δ34S for sulphides and sulphates from pre-2.45 Ga samples (triangles)with data from samples younger than ca. 2 Ga (from Farquhar and Wing, 2003). The field of isotopecompositions for barite and hydrothermal sulphide (volcanogenic massive sulphide ‘VMS’ deposits)is shown. Data for pre-2.45 Ga samples tend to plot with an aspect ratio similar to the 193 nmphotolysis array of SO2 and SO (discussed in Farquhar et al., 2001; Pavlov and Kasting, 2002).

It is thought that the early Earth experienced significant mass input from extraterrestrialsources from the sweep-up of remnant debris from solar system formation. Therefore, itmakes sense to explore whether extraterrestrial sources of sulphur could have been impor-tant to the earliest sulphur cycle.

7.5-3. EXTRATERRESTRIAL INPUTS TO THE SULPHUR CYCLE ON THE EARLYEARTH

Dust delivery has always been important to planets, and interplanetary dust was likelya major contributor to the surface inventory of biogenic elements at the formation andearly evolution of terrestrial planets (Maurette et al., 2000; Pavlov et al., 1999; Martyand Yokochi, 2006). Micrometeorites and interplanetary dust particles dominate the con-temporary extraterrestrial matter flux to Earth on the order of 103–104 tons/yr (Love and


7.5-3. Extraterrestrial Inputs to the Sulphur Cycle on the Early Earth 15

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Brownlee, 1993), or about 2–3 orders of magnitude above the calculated present-day me-teorite flux (Bland, 2005).

In the last decade evidence has emerged that many (or most) stars are born with cir-cumstellar disks with abundant dust and rocky debris (Habing et al., 2001; Carpenter etal., 2005). The solar system’s original disk became severely depleted by planetesimal ac-cretion and by planet formation within about the first 100 My (Rieke et al., 2005) haveshown that some debris disk systems observed around young (<1 Gyr old) stars by theSpitzer telescope look brighter than others of similar age and may represent dust gener-ated from evaporation of comets and/or planetary-scale bombardment events (Kenyon andBromley, 2005) well after the primary phase of planetary accretion should have ended (Be-ichman et al. 2005; Rieke et al., 2005; Meyer et al., 2006). If cometary dust productionrates from comet Hale-Bopp (Bocklée-Morvan et al., 2004) are any measure of the com-position of the early disk dust for our solar system, then the combined sulphur speciesmeasured for coma dust (OCS, CS2, H2CS) amounts to a mere ∼0.65% of the comet’s wa-ter content. At first glance, it appears that late implantation of extraterrestrial sulphur fromdust would have been a relatively minor player in the chemistry of early planetary surfacesand atmospheres. Since some meteorite components show mass-independently fraction-ated sulphur isotopes, it is germane to explore whether these inputs were at all significantto the post-primary accretion sulphur budget of the early Earth.

7.5-3.1. Extraterrestrial Dust

Interplanetary dust particles compositionally resemble the CM type carbonaceous chon-drites (Kurat et al., 1994). Murchison (CM) meteorite contains 3.3% S (Wasson andKallemeyn, 1988), and has 34S/32S and 33S/32S values comparable to terrestrial crust andmantle (Gao and Thiemens, 1993a) with average �33S = 0.013 ± 0.054�. These valuesare well within the terrestrial mass-dependent field of the sulphur isotopes. Other primi-tive meteorites such as Orgeuil (CI), considered by some to be a representative of Jupiterfamily comet compositions (Gounelle et al., 2006), contains upwards of 5% total sulphur.Orgeuil has 34S/32S and 33S/32S values slightly enriched in the heavier isotopes comparedto Murchison, and preserves an average �33S value of −0.005� (Gao and Thiemens,1993a). The majority of meteoritic materials that have been analysed so far for the minorsulphur isotopes indicate that inhomogeneities in S isotope compositions possibly relatedto nucleosynthetic, spallation reactions on Fe, or gas-phase (ion-molecule) photolytic re-actions in space, were mixed-in for the most part in the early solar system (Hulston andThode, 1965; Gao and Thiemens, 1993a, 1993b). Values of �33S not equal to 0� havebeen found in a class of igneous meteorites called the ureilites (Farquhar et al., 2000b), inorganic residues from carbonaceous chondrites (Cooper et al., 1997), in the acid-resistant‘phase Q’ of the Allende meteorite (Rees and Thode, 1977), in various phase separates ofchondrite (Rai and Thiemens, 2007) as well as achondrite (Rai et al., 2005) meteorites.Greenwood et al., 2000a) raised the possibility that the increased rates of extraterrestrialmatter input to the planets of the early solar system from later bombardments could haveimplanted heterogeneous distributions of the sulphur isotopes to the early surface.



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7.5-3.2. Sulphur and the Late Heavy Bombardment

We have no direct measure of the influx of extraterrestrial dust to Earth before about 3.85Ga during the so-called ‘late lunar cataclysm’ or ‘late heavy bombardment’ (Tera et al.,1974; Ryder, 1990). Hartmann et al. (2000) calculated that the total mass of impactors tothe Moon at ca. 3.9 Ga was approximately 6 × 1021 g, and that the amount of mass ac-cumulated by the Moon in the 4.4–4.0 Ga timeframe was roughly equivalent to the massinflux that formed the major lunar basins during the late heavy bombardment (4.1–3.85 Ga;reviewed in Mojzsis and Ryder (2001)). Therefore, the total amount of dust accumulatedto the Moon during the Hadean could have been on the order of 1.2 × 1022 g. Severalstudies have attempted to acquire information on the rate of accumulation and intensity ofimpacts at the tail end of the bombardment epoch from the study of Earth rocks. Koeberland Sharpton (1988) searched for shock features in quartz from ∼3.7–3.8 Ga rocks fromIsua (West Greenland). Later, Koeberl et al. (1998, 2000) looked for impact shocked zir-cons from the same Isua rocks, and of the hundreds of crystals studied, none showed anyevidence of optically resolvable shock deformations (Koeberl, 2003). Mojzsis and Ryder(2001) and Anbar et al. (2001) could likewise find no unequivocal chemical evidence fromplatinum-group elements for the late heavy bombardment in sedimentary rocks with ages atleast to ∼3.83 Ga (Nutman et al., 1997; Manning et al., 2006). On the other hand, Schoen-berg et al. (2002) reported tungsten isotope anomalies in ca. 3.7–3.8 Ga paragneisses fromWest Greenland and Labrador (Canada) which they interpreted to be remnants of the latebombardment in younger materials. However, follow-up studies with platinoid elements,Cr and other isotopic systems (e.g., Frei and Rosing, 2005) could find no corroboratingevidence for cosmic materials in the Isua sediments (Koeberl, 2006). To further compoundthese difficulties, the various studies cited above explored rocks with formation times (ca.3.7–3.8 Ga) that likely did not overlap with the main period of late heavy bombardment tothe inner solar system (Kring and Cohen, 2002).

Since evidence for the bombardment epoch from terrestrial rocks remains elusive, allestimates of the role of extraterrestrial matter to the early Earth must be treated with cau-tion. Marty and Yokochi (2006) estimated the total range of all extraterrestrial material(including dust and other meteoritic debris) collected on Earth’s surface from 4.4–3.8 Gato have been in the range of 2.4–7.2 × 1023 g. With the average sulphur content of CMand CI meteorites provided above as a guide, this would amount to an accumulation ofabout 0.8–3.6 × 1022 g of extraterrestrial sulphur to the planet integrated over the first 600My Although significant when compared to the cumulative mass delivered to the Moon asprovided in Hartmann et al. (2000), this amount is still �1% of the total sulphur inven-tory for all sediments + seawater presently in the surface zone of the Earth (Holser et al.,1988). In terms of MIF sulphur, Farquhar et al. (2002) argued that bulk extracts of manydifferent meteorite classes indicate that �33S values are homogeneous and range between−0.01 and +0.04�. Earth’s mantle did not inherit S-isotope heterogeneities from its ownformation. It appears that terrestrial sulphur was dominantly indigenous to the Earth sinceprimary accretion ceased prior to ∼4.4 Ga and that extraterrestrial matter was, overall, arelatively minor contributor to the global sulphur cycle to the early Earth. This is important,


7.5-4. Sulphur and the Hadean Earth Surface State 17

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because mass-independently fractionated sulphur isotopes documented in the oldest Earthrocks and minerals can with reasonable confidence be assigned an origin from terrestrialatmospheric reactions.

7.5-4. SULPHUR AND THE HADEAN EARTH SURFACE STATE

No terrestrial rocks have yet been found that yield ages older than ca. 4.03 Ga (Bowringand Williams, 1999). What we do know of the geochemical evolution of Hadean Earthcomes primarily from the study of pre-4.0 Gyr old Hadean detrital zircons shed in youngersediments captured in the Mount Narryer (Froude et al., 1983) and the Jack Hills (Comp-ston and Pidgeon, 1986) sedimentary belts in Western Australia (see Cavosie et al., thisvolume).

7.5-4.1. Surface Environments in the Hadean

A subset of Hadean zircons, which are up to 4.38 Gyr old (Harrison et al., 2005, 2006)are enriched in heavy oxygen (Wilde et al., 2001; Mojzsis et al., 2001; Peck et al., 2001;Cavosie et al., 2005, 2006; Trail et al., 2007). This observation, coupled with characteristictrace element patterns (Maas and McCulloch, 1991; Maas et al., 1992; Peck et al., 2001;Crowley et al., 2005) has provided mutually consistent evidence that an evolved rock cyclewas present within ∼150 My of the formation of the Moon (cf. Whitehouse and Kamber,2002; Nemchin et al., 2006; Coogan and Hinton, 2006). Detailed studies of Hadean zirconsfrom 176Hf/177Hf systematics show that large volumes of continental crust were presentthroughout the Hadean (Amelin et al., 1999; Harrison et al., 2005, 2006; cf. Valley, 2006).Zircon thermometry based on Ti contents of individual pre-4.0 Ga Jack Hills grains alsoshows that formation temperatures cluster narrowly around 680 ◦C (Watson and Harrison,2005). The simplest explanation for this result is that low-temperature wet minimum-meltconditions during granite genesis prevailed in Hadean melts that gave rise to the oldestzircons (Watson and Harrison, 2006). Studies of pre-4 Ga zircons now provide a ratherfirm foundation to arguments that water and a sustained atmosphere was present within thefirst 100 My of the formation of the solar system. In the absence of an actual rock recordfrom the first ∼500 My, what data for the earliest surface state might possibly be gleanedfrom sulphur isotopes?

7.5-4.2. Sulphide in Hadean Zircon

Some Hadean zircons contain inclusions of diverse minerals, but few studies have providedcomprehensive results bearing on identified inclusions in these samples (Maas et al., 1992;Cavosie et al., 2004; Crowley et al., 2005). In fact, zircons with noticeable inclusions areoften avoided during the selection process as an ion probe spot analyses that overlap in-clusions may affect data (e.g., Amelin et al., 1999). Individual Hadean zircons have beenshown to contain polyphase domains of albite, Ca-Al silicate, muscovite/quartz/Fe-oxide,



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Fig. 7.5-7. Optical micrographs of Hadean zircon sample JH992_CF01_88 from Mojzsis et al.(2001). The location of a Ni-rich pyrite inclusion is shown by the arrow. The multicollector ion mi-croprobe analysis area is indicated by the dashed oval (transmitted light micrograph at top, reflectedlight below).



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abundant quartz, K-feldspar, and phosphates such as monazite, apatite and xenotime (Maaset al., 1992; Cavosie et al., 2004; Crowley et al., 2005). In these studies, it is important todistinguish between primary melt inclusions and secondary/exsolution products of zircon.Mineral inclusions encapsulated inside zircons are only of relevance to host rock charac-teristics if it is possible to show that they represent patches of melt now solid at normalsurface temperatures and specifically of magmatic origin. Such inclusions are often ovoidand chemically incompatible with zircon. For example, in sample JH992_88 from Mojzsiset al. (2001), a small (<10 µm diameter) Ni-rich pyrite inclusion was identified near thecentre of a 96% concordant 4.105 Ga zircon with igneous [Th/U]Zr values = 0.52 (Mojzsisand Harrison, 2002). This zircon has no visible alteration or cracks leading to or from theinclusion (Fig. 7.5-7). Zircon JH992_88 has δ18OVSMOW values from two separate analy-sis spots of +5.9 ± 0.7� and +6.3 ± 0.4� (reported in Mojzsis et al., 2001), which areslightly enriched in 18O compared to average crust and mantle oxygen (+5.5�; Eiler etal., 1997), but still within the field of I-type magmas (White and Chappell, 1977). Thesulphide contained ∼2 wt% Ni as determined by wavelength dispersive spectroscopy byelectron microprobe, which along with oxygen values, is consistent with a mafic- (or ultra-mafic?) origin for zircon and its inclusion (e.g., Fleet, 1987). Three separate measurementsof the multiple sulphur isotope composition of this sulphide were obtained following thetechniques described in Mojzsis et al. (2003). The δ34S values yield a narrow range of−8.1 to −10.1� (1σ uncertainties were ±0.17�); these values are lower in 34S/32S thanaverage MORB and mantle (Chaussidon et al., 1987) but comparable to those reportedfor hydrothermally altered oceanic crust sulphides (Alt, 1995). Two of the three �33S val-ues measured for the inclusion are within 1σ error of the York-MDF band (+0.5 ± 0.5�and −0.9 ± 1.2�), and the third has a slightly 33S-depleted value of −1.3 ± 0.7� (Ta-ble 7.5-2). The absence of resolvable �33S in 2/3 of the measurements may in part be aconsequence of the low S count rates for this minute sulphide inclusion. It is obvious thatfurther analyses of this kind are needed on larger sulphide inclusions in Hadean zirconwhen they can be found. Other sulphide inclusion studies in resistant mineral hosts such

Table 7.5-2. Sulphur isotope composition of sulphide inclusion in JH992_CF01_88 by in situ ionmicroprobe multicollection

spot 34S/32S ±1σ 33S/32S ±1σ

CF01_88@1 4.4640E–02 7.89E–06 7.9399E–03 5.53E–06CF01_88@2 4.4549E–02 7.60E–06 7.9464E–03 3.94E–06CF01_88@3 4.4603E–02 7.09E–06 7.9399E–03 9.28E–06

spot δ34SVCDT ±1σ δ33S ±1σ �33S ±1σ

CF01_88@1 −8.11 0.18 −5.53 0.70 −1.35 0.71CF01_88@2 −10.12 0.17 −4.71 0.50 0.50 0.52CF01_88@3 −8.92 0.16 −5.53 1.17 −0.93 1.18

Notes: zircon 207Pb/206Pb age (1σ) = 4105±11 Ma; [Th/U]Zr = 0.52; δ18OZr = +5.9±0.7� and +6.3±0.4�at two separate spots as reported in Mojzsis et al. (2001). Average 32S count rates were: 3.62E7 to 4.23E7 cps.



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as in diamonds (Chaussidon et al., 1987) have been used to draw conclusions about thenature of sulphur in the mantle as well as crust of the Archaean. For instance, Farquhar etal. (2002) reported the presence of resolvable �33S anomalies in Archaean eclogite-typediamonds from the Orapa kimberlite (Kaapvaal Craton, Botswana) and inferred from theseresults that atmospheric MIF sulphur was transferred to the mantle. If mass-independentlyfractionated sulphur isotopes are preserved in Hadean zircons, as hinted at by the singleanalysis that provided a slightly negative �33S composition, it may be possible in the fu-ture with a greatly expanded data set to place some general constraints on the couplingbetween Hadean mantle, crust and atmosphere.

It is interesting to note that in the conceptual model of Farquhar et al. (2002) for theearly Archaean sulphur cycle, sulphur dioxide expelled from volcanoes and transformedby ultraviolet radiation can be deposited to surface reservoirs as aerosols and ultimatelyrecycled into the mantle. In their schema, sulphate in oceanic crust accumulates uniformlynegative �33S values (Farquhar and Wing, 2003) which can become reduced to sulfide(biologically or thermo-chemically), subducted, and recycled into arc magmas (Fig. 7.5-8).

Fig. 7.5-8. A schematic model for sulphur cycles on the early Earth (from Mojzsis et al., 2003).The figure represents a passive margin environment, for a model of the possible contribution ofmass-independent sulphur to subduction products on the early Earth see Farquhar et al. (2002).



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Fig. 7.5-9. Changing CO2 levels in the Earth’s atmosphere over time expressed as both partial pres-sure of CO2 in bars (left), and against present atmospheric level (PAL; right). The grey shaded regionis the range of concentrations required to compensate for the fainter young Sun (Kasting, 1993).Other data points shown are reviewed in (Rollinson, 2007). The multiple sulphur isotopes provide astrong upper constraint on the limit of CO2 partial pressure to below 0.5 bar as far back as 3.83 Ga,and perhaps before 4.0 Ga (Cates and Mojzsis, 2007b). The same may be said for partial pressuresof methane, which had to be below the threshold necessary for haze formation (Pavlov et al., 2001a,2001b).

A key feature of these arguments is that the atmosphere must have been transparent toultraviolet and that an ozone UV shield was not present for MIF sulphur isotopes to be im-parted to crustal rocks (Farquhar and Wing, 2003). This is because ultraviolet absorption bythe Rayleigh scattering tail of the CO2 absorption spectrum at partial pressures that exceed∼0.5 bar would have blocked UV sufficiently (Shemansky, 1972) to stop the reactionsresponsible for mass-independent sulphur isotope fractionations (Farquhar et al., 2001).Under these conditions, and with most solar models which show that solar luminosity was∼70% of present day values, other greenhouse gases were required to keep the HadeanEarth from freezing over (Kasting, 1993). Abundant methane is the prime candidate for agreenhouse gas that supplemented water vapor and CO2 throughout the Hadean-Archaeanto have kept the Earth above the freezing point of water (Sagan and Mullen, 1972; Kasting,2005). Sagan and Chyba (1997) and Pavlov et al. (2001a, 2001b) suggested that photolysisof methane in the mesosphere by wavelengths shorter than 145 nm could have generatedan organic haze that served as an UV shield early in Earth’s history to stabilise ammonia. Ifhowever short wavelength UV sulphur dioxide photolysis accounts for the Archaean mass-



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independently fractionated sulphur isotopes, data are inconsistent with the presence of ahigh-altitude UV shield, whether from ozone, dense CO2 or an organic haze (Fig. 7.5-9).

It may also be that the atmosphere was mostly transparent to UV for much of the Hadeanand the planet was cool (Zahnle, 2006), although the lack of evidence for glaciations olderthan ∼2.9 Ga (Young et al., 1998) apparently weakens this hypothesis.

7.5-5. SULPHUR AND THE EARLY ARCHAEAN ATMOSPHERE

Quantitative constraints on the composition of the Archaean-Proterozoic atmospherehave been attempted for many years based on investigations of redox-sensitive mineralsin the geologic record used as input parameters for model calculations. The prevailingview has been that free oxygen (O2) levels were substantially lower than present values,with estimates of pO2 � 10−2 present atmospheric level (PAL) before about 2.3 Ga (e.g.,Cloud, 1968, 1972; Walker, 1990; Holland, 1984, 1994; Kasting, 1993; Rye et al., 1997;Rye and Holland, 1998; Farquhar et al., 2000a; Pavlov and Kasting, 2002). Alternatively,some have argued for an O2-rich surface zone, or at least the condition that the atmospherewas not anoxic, for much or all of the Archaean (Dimroth and Kimberley, 1976; Towe,1990; Ohmoto, 1997; Watanabe et al., 1997). Regardless of these different arguments foran oxic Archaean, it has not been specified whether these models would also apply to theHadean Earth. Several reviews of the geologic evidence used to support the contentionthat the Archaean was anoxic by Holland (1984, 1994) noted detrital uraninite and pyritegrains in pre-2.3 Ga sediments as well as other geochemical data (e.g., Fe(II)/Fe(III) inpalaeosols) as indicative of global anoxia at time of deposition (Rye and Holland, 1998;Rasmussen and Buick, 1999).

As outlined in Mojzsis et al. (2003), the extensive production of mass-independentlyfractionated S isotopes formed on Earth in gas-phase reactions in UV transparent anoxicatmospheres, and preserved in terrestrial rocks and minerals in the absence of enough MDFoceanic sulphate to dilute mass-independently fractionated sulphur products to the surfacezone, potentially provides three fundamental constraints on the nature of the early Earth’satmosphere:(1) Separate mass-dependent arrays of slope m ≈ 0.5–0.52 within MIF δ33S vs δ34S data

from ancient sulphur minerals show that effective oxidation of sulphur in the surfacezone and homogenization of reduced and oxidised sulphur reservoirs by biological cy-cling was either absent or suppressed. In a model study, Pavlov and Kasting (2002)argued that in addition to the palaeosol and detrital mineral data cited above, pO2

levels had to be below ∼10−6 PAL throughout the Archaean. A conclusion of thiswork, and separate experimental studies of microbial sulphate reduction capabilitiesof organisms at low seawater SO4

2− concentrations (Habicht et al., 2002), is that sul-phate was severely limited in the oceans for microbial sulphate reduction before therise of atmospheric oxygen in the Proterozoic (Papineau et al., 2005, 2007; Papineauand Mojzsis, 2006). This observation is important because in standard 34S/32S isotopic


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Fig. 7.5-10. Plot of �33S vs δ34S for closed cell photolysis experiments of SO2 gas mixtures irra-diated at 193 nm (from Farquhar and Wing, 2003; data in Farquhar et al., 2001).

analyses of ca. 3.5 Ga barite deposits from the North Pole (Warrawoona Group, Pil-bara Craton, Western Australia), Shen et al. (2001) concluded that a 20� range in δ34Svalues was evidence for effective microbial sulphate reduction by 3.5 Ga. As noted pre-viously, photolysis reactions that involve SO2 +SO and other gases (e.g., H2S, CS2) inanoxic atmospheres are known to produce large MIF effects (Fig. 7.5-10) with a rangein 34S/32S of product sulphur that can mimic the entire range of biological fractiona-tions of the sulphur isotopes. This observation underscores the now critical importanceof multiple isotope measurements in all subsequent sulphur studies of ancient rocks,and warrants a comprehensive re-visit of pre-2.3 Gyr old terranes previously sampledfor sulphur (reviewed in Strauss (1997, 2003)).

(2) Experiments to delimit the range of conditions necessary for UV light-driven pho-tolysis reactions with sulphur gases deep in the atmosphere, and to investigate themagnitude of isotopic fractionations in the mass-independently fractionated products(reviewed in Thiemens (1999) and Farquhar et al. (2001)), have succeeded in repro-ducing the large range of MIF observed in Archaean sulphides. A consensus view hasemerged that column abundances of O2 and O3 before ∼2.3 Ga were very low if, forexample, 190 nm UV and other wavelengths could penetrate deep into the atmosphere



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to result in widespread MIF. Farquhar et al. (2000a, 2001) showed that this result canbe used to place an upper limit on pO2 in the Archaean of <10−2 PAL and an upperlimit on pCO2 of ∼0.5 bar (1.5 × 103 PAL). Methane may have been present, but didnot produce aerosols to attenuate UV (Fig. 7.5-9).

(3) In Fig. 7.5-6 the principal laboratory reaction that has been documented to be mostconsistent with the data thus far obtained for MIF S in the geologic record involvesSO2 photolysis at 193 nm UV (Farquhar and Wing, 2003). Pavlov and Kasting, 2002used one-dimensional photochemical models of atmospheres at various O2 concen-trations and found that the transfer of MIF S to the surface by aerosol deposition isonly possible in anoxic conditions of the early Earth (e.g., Kasting et al., 1989, 1992).The proposed mechanism is that MIF in sulphur aerosols are preserved from homoge-nization in oceanic sulphate reservoirs only in atmospheres where pO2 � 10−6 PAL.Pavlov and Kasting (2002) concluded that it is not possible to preserve MIF δ33S/δ34Sin high-O2 atmospheres; in these models, the demise of MIF S in the geologic recordis reached even before the pO2 = 10−6 PAL threshold was overtaken in the Paleopro-terozoic (Farquhar et al., 2000a; Mojzsis et al., 2003; Ono et al., 2003; Bekker et al.,2004), sometime between 2.35 and 2.42 Ga during glaciations (Papineau et al., 2006,2007).

As ever larger data sets become generated for multiple S isotopes in early Archaean rocks,results provide an increasingly robust indication that a chemically reactive, UV transparent,anoxic atmosphere existed on Earth prior to about 2.35 Ga. This anoxic regime stretchedas far back as ∼3.83 Ga, and perhaps well into the Hadean. The diverse �33S values forsulphides reported from Archaean and early Proterozoic rocks means that up to severalpercent of the marine sulphur cycle was affected by atmospheric deposition of MIF Saerosols following residence in the atmosphere.

A consequence of planetary-scale deposition of sulphur aerosols (Fig. 7.5-8) is that theycould have been tapped as a food source for early microbial communities that metaboliseelemental sulphur and sulphate.

7.5-6. SULPHUR AND EARLY LIFE

Sulphur has played a central role in biological systems since the origin of life and hasfollowed the genomic and metabolic diversification of organisms since their emergencesome 4 b.y. ago (DeDuve, 2005). A landmark in studies of the evolutionary relationshipsof organisms was established with the comparative sequence analysis of genomes used toinfer phylogenetic relationships (Woese and Fox, 1977; Woese, 2000). It is now widelyheld that the evolutionary history of life is recapitulated by its molecular biochemical in-heritance (Pace, 1997, 2001), and that all life can be classified into three Domains withinthe small subunit ribosomal RNA Phylogenetic Tree: Bacteria, Archaea and Eukarya. Inthis context, a major objective to analyse the sulphur isotope compositions of Archaeansediments has been to identify and trace the long-term evolution of microbial metabolisms(e.g., Monster et al., 1979; Schidlowski et al., 1983), and lately, to be able to relate the


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divergence of different sulphur metabolisms to events in the geologic and paleontologicalrecord (Canfield, 2004). This rationale is based on several key observations: (i) some ofthe most deeply branching lineages in the universal phylogenetic tree that reside close tothe ‘Root’ population that gave rise to all extant life, can metabolize sulphur compounds;(ii) significant sulphur isotopic fractionations can result from enzymatic processing of S-containing molecules; and (iii) sulphides preserved in the geological record have a largerange of isotopic compositions and data show that this range has increased in time (re-viewed in Canfield and Raiswell (1999)).

7.5-6.1. Sulphur Chemistry and the Origin of Life

Reduced sulphur as H2S, and its organic derivatives in the thiols (R′–SH) and the thioesters(CO–S), form from the dehydration condensation of carboxylic acid (R′–COOH) and thiolin the simple reaction:

These compounds were likely participants in the most basic proto-metabolic processes inthe origin of life. Furthermore, no discussion of sulphur biochemistry is complete withouttaking into account the rich chemistry of metal-sulphur compounds. In all contemporaryorganisms, transition metal sulphide clusters play a central catalytic role in biological en-ergy conversion systems (Beinert, 2000) and have been important since the establishmentof the earliest metabolic cycles (Rees and Howard, 2003). It is the connection betweenthe essential role of thioesters and the predominance of mineral sulphides in hydrother-mal and volcanic vents that has led a number of workers to speculate how life couldhave emerged in such environments (e.g., Russell and Hall, 1997; Russell and Martin,2004). Hydrothermal cells form wherever there is hot rock in contact with water, and suchenvironments may have been widespread in the Hadean crust. It is thus highly likely –because of the ubiquity of hydrothermal environments away from a surface zone affectedby impacts and intense UV – that the deep biosphere was amongst the earliest on Eartheven if it was not necessarily the first (Russell and Arndt, 2005). As a basis for an iron-sulphide theory for the origin of life, Wächtershäuser (1988, 1992) proposed a ‘fast origin’by an autotrophic proto-metabolism of low-molecular weight constituents (Fig. 7.5-11)in hydrothermal environments. Along these lines, Cody et al. (2000) presented experi-mental support for the formation of pyruvic acid (CH3–CO–COOH) from formic acid(HCOOH) in the presence of organo-metallic sulphur compounds such as nonylmercap-tane (Fe2(CH3S)2–(CO)6), and iron sulphide (FeS2) at high temperatures and pressuresconsistent with shallow oceanic crust near volcanic centres (250 ◦C and 200 MPa). Mix-ture of these reaction products with lower temperature fluids that emanated further awayfrom vent centres, such as is found at the crust-ocean interface, would have been moreconducive to the survivability of other important biomolecules such as amino acids and



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Fig. 7.5-11. Basic proto-biotic reactions that could have taken place in the ‘Iron-Sulphur World’(modified after Wächtershäuser (2000)).

nucleosides (Orgel, 1998). In a separate analogy with the H2S/CO system, the formationof FeS2 in the simple reaction:

FeS(pyrrhotite) + H2S → FeS2(pyrite) + H2

produces molecular hydrogen. The FeS/H2S couple is a strong reductant that could haveplayed an important supporting role in prebiotic organic synthesis under crustal conditionscommon to the Hadean Earth.

Therefore, a self-consistent picture has emerged that hydrothermal zones in shallowHadean oceanic crust were, if not the birthplace of life on Earth as advocated by Russelland Arndt (2006) and Russell et al. (2005), at least were the crucible wherein early biolog-ical forms found a ready source of fluids rich in reduced carbon and more complex organicmolecules (Shock and Schulte, 1998). In this scenario, electron acceptors such as S0 eitherfrom recycled sulphur in oceanic crust, from sulphidic hydrothermal fluids, or transferredfrom the water column to the crust plumbing system (Fig. 7.5-1) to be catalysed by car-bonylated iron-sulphur clusters (Cody et al., 2000), could have provided both a kick-startto, and a rich food source for, the Hadean biosphere.

A primordial repose for life in deep hydrothermal vent systems that metabolized re-duced sulphur compounds is further supported by molecular phylogenetic studies. Bio-chemical evidence shows that some of the earliest biomes were located in deep hydrother-



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Fig. 7.5-12. Phylogenetic tree based on ssu 16s rRNA comparisons which shows Bacterial and Ar-chaeal lineages of extant organisms capable of elemental sulphur reduction (framed characters),sulphate reduction (underlined lineages) and sulphide oxidation (boxed lineages). This figure is mod-ified from Pace (1997) with information from Klein et al. (2001), Canfield and Raiswell (1999),Stetter and Gaag (1983).

mal environments. A crucial caveat to these studies is that before we actually know howlife on Earth began there is no definite requirement that life started in a mild-, medium- orhot-hydrothermal vent field, even if it is a compelling and convenient place for the originof life (Nisbet and Sleep, 2001). Sulphate and other oxidised S species figure prominentlyas electron acceptors for those deeply branching microbes that perform microbial sulphatereduction, yet were probably of limited use in the absence of significant sulphate beforethe Proterozoic. Instead, the ancient metabolic cycles of life that arose out of the metal-Schemistry described above in all probability metabolized volcanogenic sulphur gases suchas H2S, SO2 and sulphur compounds including elemental S.

7.5-6.2. Sulphur Metabolic Styles of Early Life

Many hyperthermophilic, deep-branching organisms in the domains Bacteria (Thermoto-gales, Proteobacter) and in various Archaea (sub-Domain Crenarcheota and Pyrococcus,Thermoplasma in sub-Domain Euryarcheota) are known to reduce elemental sulphur to



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H2S with both H2 and organic compounds as an electron donor (Fig. 7.5-12). This path-way could have evolved from proto-metabolic metal-sulphur cycles, and presently involvesa short electron transport chain controlled by the activity of key enzymes including sulphurreductase, polysulphide reductase and hydrogenase (Laska et al., 2003). The antiquityof this metabolism is further evident in that both deep-branching bacterial and archaeallineages share this metabolic style (Stetter and Gaag, 1983; Stetter, 1996; Canfield andRaiswell, 1999), an observation that strongly suggests that elemental sulphur reduction(ESR) might be an inherited trait from the last common ancestor of all life that resided ina high-temperature environment (Corliss, 1990).

Supposing that ESR is truly ancient, it still remains unknown whether the last commonancestral population at the base of the phylogenetic tree (labelled ‘Root’ in Fig. 7.5-12)included the genes for this particular metabolic style. Elemental sulphur reduction couldvery well have evolved later in one of the domains and subsequent lateral gene transferoccurred between domains. Indeed, the lateral-transfer argument holds for all the discus-sions on the early rise of metabolisms interpreted from extant phylogenies if they cannotbe specifically tied to molecular or isotopic biomarkers in the geologic record. It is antici-pated that future phylogenetic comparisons with amino acid sequences of specific enzymeswill better constrain the branching sequence of lineages. If a large number of lineages thatperform a specific metabolic pathway are known from both the bacterial and archaeal do-mains, evidence would seem to favour a deeply ancient origin. A critical clue may be thatno known Eukaryal organism has this metabolic style. Anaerobic respiration with elemen-tal sulphur is performed by several bacteria and archaea, but has only been investigated ina few organisms in detail (Hedderich et al., 1998). As was noted by Canfield and Raiswell(1999), organisms that do ESR can readily obtain elemental sulphur from rapid cooling ofvolcanogenic sulphur gases where:

SO2 + 2H2S ↔ 3S0 + 2H2O.

This reaction shifts the equilibrium in favor of S0 (Grinenko and Thode, 1970). For exam-ple, with S0 present, microbes can perform the reaction:

4S0 + 4H2O → 3H2S + SO42− + 2H+

and in the presence of iron hydroxides, which were abundant in the oceans before 1.9 Ga(see Holland, 1984; Bjerrum and Canfield, 2002; Thamdrup et al., 1993):

H2S + 4H+ + 2Fe(OH)3 → 2Fe2+ + S0 + 6H2O

As implied by Fig. 7.5-11, the rich chemistry at the interface of metal-sulphur and bio-chemistry is not restricted to iron, even if that particular branch of inorganic biochemistryhas always been dominated by Fe-S (Daniel and Danson, 1995). Many metals which hap-pen to be abundant in most mafic and ultramafic rocks readily interact with sulphur, suchas Zn, Mo, Co and Ni. Average MORB contains 79 ppm Zn, 0.46 ppm Mo, 47 ppm Coand 150 ppm Ni (Hertogen et al., 1980; Newsom et al., 1986; Hofmann, 1988). Tran-sition metal-sulphur chemistry figures prominently in the biochemistry of many deeply-branching microbes (reviewed in Konhauser, 2007). Most of the various catalytic roles of



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metal ion + S clusters other than Fe await experimental verification from analysis of con-trolled (culture) experiments and such experiments represent an important new avenue forresearch into the question of a biological origin of sulphur isotope fractionations in thegeologic record.

Microbial sulphate reduction (MSR) is an anaerobic metabolic pathway in which SO42−

is used as an electron acceptor to be reduced to H2S with electron donors such as H2 and/ora variety of organic molecules to produce hydrogen sulphide in the general form:

SO42− + 5H2 → H2S + 4H2O

Sulphate reduction in bacteria and archaea can proceed via dissimilatory or assimilatorypathways that either excrete H2S or insert it in organic compounds. A few key enzymesand compounds in MSR are illustrated in Fig. 7.5-13. The timing of evolution of thesegenes is not accurately known, but qualitative information can be obtained from molecularphylogeny of sulphate reducers.

7.5-6.3. Biological Sulphur Isotopic Fractionations

Sulphate-reducers preferentially metabolize 32SO42− rather than 34SO4

2− to produce iso-topically light H2S with a range of δ34S values between −4 and −46� under non-limiting concentrations of SO4

2− (Canfield and Raiswell, 1999, and references therein).Microbially-produced H2S can react with dissolved Fe2+ or other metal cations to precip-itate sulphides in the simple reaction:

2H2S + Fe2+ → FeS2 + 2H2

Diagenetic sulphides formed from H2S and FeS (or other metal-sulfide) reactions carry thecombined mass-balanced sulphur isotopic composition without significant isotopic frac-tionation (Butler et al., 2004; Böttcher et al., 1998). These isotope ratios can then becomeincorporated and preserved in the geological record of authigenic sulphides. The sulphurisotopic composition of sedimentary sulphides can therefore be used to pinpoint the iso-topic fractionation between dissolved SO4

2−, and product H2S as an isotopic biosignatureto trace the early evolution of microbial sulphur metabolisms.

It has been experimentally verified that microbial sulphate reduction under limitingSO4

2− concentrations (i.e., lower than 200 µM) has insignificant effect on the δ34S valueof product H2S (Habicht et al., 2002). If the concentration of SO4

2− in the environment islow, microbial sulphate reduction produces H2S (and thence sulphide) with approximatelythe same δ34S value as the available δ34Ssulphate. The δ34S of sulphides from ancient sed-iments can also be used to trace the concentration of dissolved SO4

2− in seawater at thetime of sedimentation. Sedimentary sulphides with negative δ34S values (i.e., <−10�)are consistent with microbial sulphate reduction in high seawater SO4

2− concentrations.Variability in sulphur isotopic fractionations during MSR can also depend strongly on thesources of organic compounds consumed during growth (Kleikemper et al., 2004) andon the different lineages of sulphate reducers (Detmers et al., 2001). Some organisms



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Fig. 7.5-14. Isotopic fractionation in the modern sulfur cycle showing the values for average sulfurfractionations (± standard deviations) associated either with reduction or oxidation. Each dispropor-tionation reaction couples the production of both H2S and SO4

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can also disproportionate thiosulphate (S2O32−), SO3

2− and/or S0, and the disproportion-ation of sulphur-containing compounds produces different molar amounts of both H2Sand SO4

2− with isotopic fractionations characterized by enrichments of 34S in SO42−

and depletions of 34S in H2S (Böttcher, 2001; Habicht et al., 2002). Natural microbialcommunities are known to produce sulphide significantly more depleted in 34S than isexpected from microbial sulphate reduction alone. The disproportionation of intermedi-ate sulphur-compounds metabolized by sulphate reducers can recycle sulphur through yetgreater degrees of isotopic fractionation and this has been used to explain the observedlarge range of δ34S in modern sediments (Habicht and Canfield, 2001; Canfield and Tham-drup, 1994). A schematic representation of the different sulphur isotopic fractionationsis shown in Fig. 7.5-14. Sulphur oxidation does not significantly fractionate δ34S valuesbut returns oxidized sulphur compounds to the environment, which in the context of a hi-erarchical microbial community can then be recycled as electron acceptors for microbialsulphate reduction or disproportionation reactions.

7.5-6.4. Non-Authigenic Sulphur in Sediments

Because sulphur is so widely distributed in the Earth and occurs in oxidized forms(sulphate), neutral state (elemental S0) and reduced forms (sulphides), it readily travelsthroughout the rock cycle. Criteria used to distinguish sedimentary from hydrothermal sul-phides include the crystal habit, sulphur isotopic composition, and sometimes an unusualabundance of trace metals. Detrital sulphides, which can sometimes be identified fromtheir rounded appearance and occurrence with other characteristic detrital minerals, areolder than the depositional age of the sediment. Sulphide formation can postdate the origin



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of host sedimentary rocks if they formed from the crystallization of remobilized sulphurfrom post-depositional circulating fluids. Hydrothermal fluids may assimilate sulphur fromhost sedimentary rocks during water-rock interactions and therefore carry the δ34S valueof the host rock. Pyrite in sediments may have the isotopic signature of microbial sulphatereduction, and subsequent hydrothermal circulation can transport Cu that will react withpyrite to form chalcopyrite with the same δ34S as the precursor pyrite (Ohmoto and Gold-haber, 1997; Papineau and Mojzsis, 2006). High concentrations of trace metals such as Co,Zn and Ni in sulphides may be used as an indicator of diagenetic conditions or interactionwith magmatic or other fluids. The δ34S value of sulphur in igneous rocks from the uppermantle is in the range of −0.7 ± 5.0� (Chaussidon and Lorand, 1990) and more generallyis in the range of −3.0 to +3.0� (Ohmoto, 1986).

Sulphur isotopes of Archaean magmas (Hattori and Cameron, 1986) have mass-dependently fractionated �33S values between −0.3 and +0.3�, unless the magma orig-inated primarily from recycled Archaean sediments with a nonzero �33S and the MIFsulphur isotopes were isolated from dilution by mantle sulphur (e.g., Farquhar et al., 2002;Mojzsis et al., 2003). Metamorphic effects only slightly fractionate sulphur isotopes, andthese minor effects are mass-dependent. Metamorphism can potentially modify the �33Svalue of sedimentary sulphides but it is thought that this can only be achieved by inter-action with sulphur-containing metamorphic fluids such that a pre-existing MIF signatureis either diluted or remobilized by the fluid. At 900–950 ◦C, the self-diffusion of S in sul-phides (pyrrhotite) is slow, (log D = 10−16 m2 s−1; Condit et al., 1974), which will tend topreserve mass-independent sulphur isotope signatures even in metamorphic sulphides, butan important caveat to this is that it must otherwise occur in the absence of massive S-richfluid infiltration and replacement (Papineau and Mojzsis, 2006).

Because of these competing factors, it must be noted that near-zero �33S values forArchaean sedimentary rocks may not necessarily imply high levels of atmospheric O2at time of deposition. Mass-independently fractionated sulphur isotopes can be dilutedby mass-dependently fractionated sulphur, especially if sulphate production is enhancedin local Archaean environments (Huston et al., 2001a). This point probably explains theobservation that some sulphides with nonzero �33S values from Archaean sediments co-exist in samples that also have mass-dependently fractionates sulphur isotopes (Mojzsis etal., 2003; Hu et al, 2003; Farquhar et al., 2000a; Whitehouse et al., 2005). The presence of amass-independently fractionated sulphur isotopic signature is consistent with sedimentarysulphur that cycled through the atmosphere. However, to propose that mass-independentsulphur isotopes were absent at time of deposition requires other lines of evidence such asδ34S, sulphide composition, morphology and petrogenesis to firmly assign a sedimentarysource of sulphur. In sum, without more data to support a sedimentary source, the near-zero �33S values of these samples cannot be used to constrain atmospheric O2 at time ofdeposition.

The crystal habit of the analysed sulphide crystals from metasedimentary rocks canvary from anhedral to euhedral and may be related to degree of recrystallization relatedto metamorphism. All Archaean sedimentary rocks experienced some degree of metamor-phism, which may have modified the original sulphide morphologies. However, analysis



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Fig. 7.5-15. Sulphur multiple isotope data (expressed as �33S) for sedimentary sulphides and sul-phates as a function of age (replotted after Farquhar et al. (2000); see also Seal (2006)). A large rangein �33S values for samples older than 2.4 Ga most likely reflects a change in atmospheric oxygenconcentration past the pO2 = 10−6 PAL threshold (Pavlov and Kasting, 2002). Filled symbols aredata from sulphides, open symbols are sulphates.

of visible light and electron micrographs of sulphides thus far analysed in situ have failedto show any obvious correlations between sulphur isotopic composition and grain shapeand/or size (e.g., Papineau and Mojzsis, 2006). Papineau et al. (2005, 2007) have arguedthat the habit of a sulphide grain is a weak criterion to assess the sedimentary origin of thesulphur, unless specific features can be identified, such as sulphide nodules. Metamorphicrecrystallization is probably responsible for most of the large euhedral crystals observed insamples of various ages. Visible light, as well as electron micrograph images have so farfailed to show any evidence for sulphide overgrowths on pre-existing sulphide cores in thehundreds of analysed sulphides from various Archaean localities worldwide.

7.5-6.5. Tracking the Antiquity of Sulphur Metabolisms

A consequence of the model for anomalous �33S sulphur aerosol formation is that pyrite,forming via elemental sulphur reduction by organisms to H2S, would retain positive �33Ssignatures, while pyrite from sulphate reduction would be expected to preserve negative�33S values. The latter case may be seen in published �33S data for sedimentary rocks



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Fig. 7.5-16. Three-isotope plot of sulphur isotope data from sulphides analysed in Isua supracrustalbelt metasediments showing the MDF bands (grey lines) reported in Papineau and Mojzsis (2006).The separate MIF arrays in the data shows that effective homogenization of the S reservoirs bymixture with sulphate was suppressed in the anoxic Eoarchean oceans. Mojzsis et al. (2003) notedthat the paucity of negative �33S in early Archaean rocks argues against widespread MSR at thattime and could provide crucial clues about BIF origins.

�2.5 Ga where pyrites preserve both negative and positive mass-dependent fractionations(Fig. 7.5-15). In the majority of Archaean samples for which multiple sulphur isotope dataare available, most sulphides have strongly positive �33S values. Mojzsis et al. (2003)and Papineau and Mojzsis (2006) interpreted these data to be inconsistent with sulphatereduction and more in line with elemental sulphur reduction. The presence of such positive�33S values provides a strong case that the source of sulphur in these systems was cycledthrough the atmosphere and introduced to the oceans in the early Archaean, and separatelinear mass-dependent arrays in δ33S/δ34S space show that the sulphur was subsequentlyisolated there from re-homogenisation by volcanic and hydrothermal sources (Fig. 7.5-16).

Elemental sulphur (S0, S8) aerosols derived from mass-independent isotope fraction-ations from atmospheric reactions and implanted into the oceans would have provided aready source of sulphur to a plethora of environments at the global scale from the Hadeanonwards. In the absence of oxidative weathering of surface pyrite or other ready sources ofsulphate, anoxic waters contain no more than trace sulphate over HS− + H2S (Thorsten-son, 1970). Although far from representing a consensus view, the current paradigm holdsthat the Archaean oceans were generally poor in sulphate relative to present values (Hol-


7.5-7. Sulphur Isotopes in Early Archaean Rocks 35

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land, 1984) at least before the Paleoproterozoic oxygenation of the surface environment(Papineau et al., 2007).

7.5-7. SULPHUR ISOTOPES IN EARLY ARCHAEAN ROCKS

Sulphides are often preserved in ancient marine sediments but sedimentary sulphateminerals have yet to be reported in rocks older than about 3.5 Ga (Huston and Logan,2004). The oldest evidence of sedimentary sulphate occurs as ∼3.4–3.5 Ga barite fromWestern Australia, South Africa and India (reviewed in Van Kranendonk (2006)). It isbecause of the paucity of early Archaean multiple sulphur isotope data that the sulphurisotopic composition of seawater sulphate during the 3.5–3.2 Ga interval is not well es-tablished, and not established at all for before 3.5 Ga. However, it is not unreasonable tosuppose that seawater δ34Ssulphate at the time of deposition of Isua sedimentary protolithsat ca. 3.8 Ga (see below) was not significantly different from the 3.4–3.5 Ga seawaterδ34Ssulphate of between +2.7 and +8.7� (Strauss, 2003).

As opposed to the conventional 34S/32S measurements, an increasingly large body ofmultiple sulphur isotope data has been collected in the last few years for several of thekey metasedimentary units in Western Australia (ca. 3.5 Ga) and West Greenland (ca. 3.8Ga). These studies have brought to light several important features of the early Archaeansurface state.

7.5-7.1. Sulphides and Sulphates from the (ca. 3.49 Ga) Dresser Formation, WarrawoonaGroup, Western Australia

Microscopic sulphide (pyrite) grains in ca. 3.49 Ga barite beds from the Dresser Formation,Warrawoona Group at North Pole (Western Australia) have δ34S values up to 24� lighterthan co-existing sulphate (barite) from the same unit. These data were interpreted by Shenet al. (2001) as evidence of microbial sulphate reduction in the Paleoarchaean, an interpre-tation that could also be viewed consistent with phylogenetic arguments that suggest anearly evolution of microbial sulphate reduction (Fig. 7.5-12). However, to assign an actualtime of relative phylogenetic divergence episodes on the RNA phylogenetic tree (Pace,2001) should be attempted with extreme caution because changes per nucleotide/time havenot yet been established with confidence. Smaller ranges of δ34S values in pyrite fromMesoarchaean marine metasedimentary rocks of between −2.5 and +8.5� have been putforth as evidence for sulphate-rich Archaean oceans with concentrations >1 mM that ex-isted under 10−13 PAL O2 (Ohmoto and Felder, 1987; Ohmoto, 1997, 1999; Ohmotoet al., 1993). Such concentrations are about 5 times higher than what is expected fromexperiments that show similarly small sulphur isotopic fractionations when microbial sul-phate reduction occurs at seawater sulphate concentrations of less than 0.2 mM (Habichtet al., 2002). If the Archaean atmosphere was oxygen-rich and stabilised sufficient sul-phate in seawater to form the large barite deposits, it requires distinctive interpretations ofgeologic data and of thermodynamic arguments for oxidative weathering of siderite and



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Fig. 7.5-17. Evolution of 34S/32S sulphur isotope ratios in the geological record. The two upperlines represent the band of δ34Ssulfate values and the diamonds are compiled published δ34Ssulfidedata for sedimentary rocks. Before ∼1.7 Ga, constraints on δ34Ssulfate are scarce and the lower lineis a 55� fractionation from contemporaneous seawater sulphate (from Shen et al., 2001). Data forpre-3.5 Ga sediments are incomplete for this graph.

detrital heavy minerals (Ohmoto, 1999; cf. Rasmussen et al., 1999), a different explana-tion of BIF genesis, and alternative interpretations of δ13Corg, δ13Ccarb and δ34Ssulphidedata for Archaean sedimentary rocks (Ohmoto, 1997). The proposed evolution for sea-water δ34Ssulphate is illustrated as a continuous band in Fig. 7.5-17. The observation thatthe range of δ34Ssulphide values has increased over time – specifically during the Pale-oproterozoic and the Neoproterozoic – is instead most easily explained by significantaccumulations of atmospheric O2 after about 2.5 Ga which helped stabilise oceanic sul-phate at higher concentrations (Holland, 1984). The generally small ranges of δ34S valuesin Archaean metasedimentary rocks appear to indicate that either sulphate-reducing mi-crobes were frustrated by the generally low concentration of seawater sulphate and limitedto specific microenvironments where sulphate was present, or they had not yet evolved.

The data of Shen et al. (2001) were collected in cherty barite-containing rocks fromthe Dresser Formation. The rocks experienced low degrees of metamorphism and slightdeformation (Van Kranendonk, 2006), which argues against extensive post-diagenetic al-teration of the sulphur isotopes. Groves et al. (1981) interpreted the depositional settingfor these units as sedimentary deposits with precipitated gypsum (later replaced by barite)



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from brine ponds separated from the sea by a sand berm. Buick and Dunlop (1990) pos-tulated the origin of the precipitated sulphate from originally low sulphate sea water thathad seeped into briny lagoons and evaporated to concentrate sulphate. These authors alsosuggested that sulphate may have been locally supplemented by the phototrophic oxidationof volcanogenic sulphide.

In a series of papers, Van Kranendonk et al. (2001), Van Kranendonk and Pirajno (2004)and Van Kranendonk (2006) presented an alternative explanation for the origin of thesulphate with new high-resolution mapping and geochemical analyses of the North Polerocks. This work convincingly showed that the extensive bedded chert + barite units ofthe Dresser Formation formed during discrete episodes of volcanogenic hydrothermal cir-culation during exhalative cooling of a felsic magma chamber. A similar conclusion wasreached for the origin of sulphates in the 3.2–3.6 Ga Panorama volcanic-hosted sulphidedistrict, also in Western Australia (Huston et al., 2001a). Under conditions analogous tocontemporary S-rich volcanism at Mt. Pinatubo (Philippines), a silicic magma chamberwith associated caldera collapse and seawater incursions will evolve highly oxidized (Scail-let et al., 1998) volatile-laden saline fluids (Newton and Manning, 2005). Such conditionsstabilise SO2 and lead to the expulsion of oxidised fluids, which in the hydrolysis reaction(after Hattori and Cameron, 1986):

4H20 + 4SO2 ↔ H2S + 3H+ + 3HSO−4

leads to the highly localised enrichment of sulphate, without the intervention of oxygenicphotosynthesizers. In the case of the Dresser Formation barites, locally recharged subma-rine brine-pools with reactive Ba2+ leached from basalt would have formed barite (cf. VanKranendonk, 2006). Sulphate would have been rapidly scavenged by Ba2+; the solubilityof barite is low (�H ◦

r = 6.35 kcal mol−1, log K = −9.97; Stumm and Morgan (1996)),and Ba2+ leached from basalt (average Archaean basalt contains 569 ppm Ba; Condie(1993)) was precipitated as barite syndepositionally with pyrite.

Multiple S-isotope measurements of sulphate-sulphide pairs provide a crucial clue tothe origin of the sulphur in the Dresser rocks. Barites from Dresser sample GSWA 169711(A-I) reported in Farquhar et al. (2000a) have average �33S values = −0.98 ± 0.02� andaverage δ34SVCDT = +4.9� that form a well-defined (r2 = 0.977) linear array of MDFslope λ = 0.51 ± 0.04 (Fig. 7.5-18(a)). It is thus apparent from these data that the sulphurresponsible for this barite resided for some time in the Archaean atmosphere before it wasdeposited in the water and stabilised as sulphate. Core rock samples of black chert withfinely disseminated pyrite, also from the Dresser Formation (core sample GSWA 169729),have average �33S values = +3.67 ± 0.03� and average δ34SVCDT = +2.70 ± 0.01�(Table 7.5-3) that form a linear array (r2 = 0.893) of non-MDF slope λ = 0.837 ± 0.16(Fig. 7.5-18(b)). When these data are plotted in �33S vs δ34S space (Fig. 7.5-19), it isapparent that they are consistent with samples that experienced SO2 (or SO) photolysis atshort UV (e.g., 193 nm) wavelengths (compare with Fig. 7.5-6).

The differences in pyrite and barite multiple sulphur isotope compositions from theDresser rocks show that deposition of MIF sulphur as pyrite and barite was swift onceaerosols reached the water column. Because neither hydrothermal nor biological cycling



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Fig. 7.5-18. A. Plot of multiple S-isotope measurements of barites from Dresser sample GSWA169711 (A-I) reported in Farquhar et al. (2000a) shows a well-defined (r2 = 0.977) linear array ofMDF slope λ = 0.51 ± 0.04. B. Core samples of pyrite from black chert from the Dresser Forma-tion sample GSWA 169729 form a less well-defined linear array (r2 = 0.893) of non-MDF slopeλ = 0.837 ± 0.16.

of sulphate had sufficient time to homogenise the �33S values, it may be the case thatneither process was important at time of deposition.


7.5-7.Sulphur

Isotopesin

Early

Archaean

Rocks

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11

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33

44

55

66

77

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Table 7.5-3. Sulphur isotope composition of sulphides in Dresser Mine core sample 169729-1B by in situ ion microprobe multicollection

spot 34S/32S ±1σ 33S/32S ±1σ δ34SVCDT ±1σ δ33S ±1σ �33S ±1σ

G1_s1 4.5135E–02 6.32E–07 8.0006E–03 1.67E–07 2.89 0.01 5.22 0.02 3.73 0.03G1_s2 4.5114E–02 5.63E–07 7.9969E–03 2.06E–07 2.44 0.01 4.76 0.03 3.50 0.03G1_s3 4.5110E–02 3.74E–07 7.9962E–03 1.98E–07 2.35 0.01 4.67 0.02 3.45 0.03G2_s1 4.5113E–02 6.50E–07 7.9986E–03 1.88E–07 2.41 0.01 4.97 0.02 3.72 0.03G2_s2 4.5134E–02 2.24E–07 8.0011E–03 2.53E–07 2.89 0.00 5.29 0.03 3.79 0.03G2_s3 4.5119E–02 5.82E–07 7.9993E–03 1.63E–07 2.53 0.01 5.07 0.02 3.75 0.02G3_s1 4.5134E–02 7.65E–07 8.0003E–03 1.63E–07 2.88 0.02 5.19 0.02 3.69 0.03G3_s2 4.5132E–02 3.42E–07 8.0005E–03 1.81E–07 2.84 0.01 5.21 0.02 3.74 0.02G3_s3 4.5128E–02 6.38E–07 7.9990E–03 1.33E–07 2.75 0.01 5.02 0.02 3.59 0.02G4-s1 4.5140E–02 6.76E–07 8.0014E–03 2.10E–07 3.00 0.01 5.32 0.03 3.76 0.03

Notes: Nomenclature is grain number + ion microprobe spot number. Error demagnification and corrections for instrumental mass bias were applied to the data.Details for these measurement conditions are reported in Papineau et al. (2005).

dpg15

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docsubty:

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Fig. 7.5-19. Plot of �33S vs δ34S for sulphides and sulphates from the 3.46 Ga Dresser Fm. (sul-phate data from Farquhar et al. (2000)). The field of isotope compositions for barite and hydrothermalsulphide (volcanogenic massive sulphide ‘VMS’ deposits) is also labelled. The data for the sul-phide-sulphate pair tend to plot with an aspect ratio similar to the 193 nm photolysis array of SO2and SO (see Fig. 7.5-6).

7.5-7.2. Sulphides from the (3.8–3.7 Ga) Isua Supracrustal Belt, West Greenland

Amphibolites of dominantly basaltic composition from the ca. 3.77 Ga Isua supracrustalbelt (ISB), 150 km northeast of Nuuk, West Greenland, have δ34S values between −1.0 and+2.6� (Monster et al, 1979), consistent with a magmatic origin and an Archaean mantlewith near-zero δ34S values (Hattori and Cameron, 1986). Rocks of sedimentary protolith,such as banded iron-formations (BIF) and garnet-biotite schists of probable ferruginousclay-rich sedimentary origin are recognizable throughout the ISB (reviewed in Nutmanet al. (1984a)). Because these rocks are of marine sedimentary origin (Dymek and Klein,1988) they are important candidate materials to host information relevant to the sulphurgeochemistry of the Archaean oceans. In line with these observations, it was proposedsome time ago that sulphur isotopes in sulphides from the metamorphic equivalents of



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Fig. 7.5-20. Comparison of published �34S values for various sulfide phases in Isua supracrustalbelt metasedimentary rocks (from Papineau and Mojzsis, 2006).

pelagic sedimentary rocks could preserve isotopic evidence of the early biosphere (Monsteret al., 1979; Schidlowski et al., 1983).

Bulk sulphides from Isua banded iron-formations have δ34S values in the range of −1.0to +2.0�, comparable to values measured from associated garnet-amphibolite rocks ofbasaltic composition noted above, which volumetrically dominate the lithotypes of the ISB.Similar δ34S values for pyrite in other BIF units from Isua range between −1.2 to +2.5�,and were interpreted by Strauss (2003) to reflect magmatic and hydrothermal processesin the sulphur cycle, with no apparent biological fractionation. The total range of δ34Svalues measured for metasedimentary rocks of the ISB extends over ∼9� and is cen-tred around +1� (Fig. 7.5-20). This δ34S range is similar to other sulphur isotope rangesthat characterize some Mesoarchaean marine sediments (Ohmoto et al., 1993; Ohmotoand Felder, 1987). A possible explanation is that the small variations in δ34S of sulphidesfrom Isua metasediments reflect mass-dependent fractionation processes during metamor-phism, which depend on temperature, sulphide phase, cooling time, etc. Alternatively, sucha range may be consistent with biological fractionation at the time of deposition, but therange in δ34S is small and as outlined previously, microbial sulphate reduction was eithersuppressed or absent at Isua time. More data are required to test the hypothesis that el-emental sulphur reduction or some other early metabolic cycling of sulphur was present



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when the Isua sediments formed, especially if the range of δ34S for co-existing sulphurphases can be expanded further.

Published sulphur isotope analyses of anhedral pyrites that occur as interstitial blebsin biotite from various quartz-biotite schists (‘metapelites’ of clay-rich origin, as opposedto banded iron-formations) from the ISB showed positive �33S values between +1.06and +1.43�, and a range of δ34S between −2.82 and +5.44� (Mojzsis et al., 2003;Papineau and Mojzsis, 2006). Banded iron-formation samples have also been investigatedfor their multiple sulphur isotopes from the northeast sector of the ISB. Many of these rockscontain bands of large (<200 µm) pyrite grains within grunerite+magnetite and quartzbands. Published data for subhedral to anhedral sulphide for the ISB can be used to definemaximum ranges of δ34S data for sulphides (n = 161) as between −3.1 and +5.8�,and for reported �33S data (n = 99) between −0.87 and +3.41�. A compendium ofpublished and recent δ34S data from metasedimentary rocks from the ISB is presented inFig. 7.5-20.

From the standpoint of environmental chemistry, the multiple sulphur isotope data forthe Isua supracrustal belt provides evidence for an anoxic atmosphere at 3.8 Ga and lowconcentrations of dissolved seawater sulphate and atmospheric oxygen. Under these con-straints, it is interesting to consider whether the early Archaean oceans were severely lim-ited in many oxidants necessary for various microbial respiration pathways. The dominantbiome could have been carbon-fixing autotrophs as suggested by 13C-depleted graphiteparticles reported from Isua (Ueno et al., 2002; Rosing, 1999; Mojzsis et al., 1996).This conclusion would mean that the low abundance of oxidants in early Earth environ-ments was not conducive to an extensive biosphere nor to metabolically diverse microbialcommunities before the Paleoproterozoic transition from a globally anoxic world to thebeginnings of a oxygenated planet (Papineau and Mojzsis, 2006; Papineau et al., 2007).

7.5-7.3. Sulphides from the �3.83 Ga Akilia Association, Akilia Island, West Greenland

As outlined in Manning et al. (2006, and references therein), the supracrustal rocks ofthe Akilia association at the type locality of Akilia, an island ∼30 km S of Nuuk, WestGreenland, have undergone a complicated metamorphic history that led to recrystalliza-tion under granulite facies conditions (Griffin et al., 1980). The complexity of these rocks,compounded with a long history of ‘reconnaissance studies’ from this terrane, has beenthe principal reason for current debates in the interpretation of protoliths for the Akiliaenclaves. On the south-western peninsula of Akilia are two units of ferruginous quartz-pyroxene rocks sandwiched between amphibolites (Manning et al., 2006). The mineralcomposition of the units is dominated by quartz, clinopyroxene, magnetite, hornblende,sulphides and minor accessory phases including zircon and graphite (Nutman et al., 1997).The age of the quartz-pyroxene rock has been the subject of debate. One interpretationholds that the units are greater than the ca. 3.83 Ga zircon U-Pb ages for tonalitic or-thogneisses, some of which were shown to intersect the supracrustal enclave (Manninget al., 2006). Published U-Th-Pb depth-profiles from Akilia orthogneiss zircons by ionmicroprobe (Mojzsis and Harrison, 2002a) and Pb-Pb ages on apatite separates in the



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quartz-pyroxene rock (Sano et al., 1999b; cf. Mojzsis et al., 1999) indicate a complicatedmetamorphic history with regional major events at ∼3.73 Ga, ∼3.65 Ga, ∼2.65 Ga and∼1.50 Ga (reviewed in Cates and Mojzsis, 2006). The protolith of the quartz-pyroxenerocks was originally interpreted as banded-iron formation (BIF) by Nutman et al. (1997)and was used by them to earmark these particular samples for biosignatures analyses ofC-isotopes. Detailed geochemical results published in the last few years provide supportingevidence for a sedimentary for these rocks. These data include major, minor, trace elementand rare earth distributions (Mojzsis and Harrison, 2002b; Friend et al., 2002c; Nishizawaet al., 2005; Manning et al., 2006), δ56Fe and δ57Fe values3 (Dauphas et al., 2004), δ18Ovalues (Manning et al., 2006) and �33S values (Mojzsis et al., 2003). In an alternativeinterpretation, REE patterns and trace elements have been used to argue instead for a meta-somatized ultramafic igneous protolith for the Akilia quartz-pyroxene rocks (Fedo andWhitehouse, 2002; Bolhar et al., 2004; cf. Johannesson et al., 2006). Also, Whitehouse etal. (2005) were unable to corroborate previously reported mass-independent sulphur iso-topes values for these units (Mojzsis et al., 2003). Lastly, Manning et al. (2006) showed thatonly an origin as chemical sediment with minor detrital input satisfies all field, petrologicand geochemical tests, and because the Akilia quartz-pyroxene rocks represent a primarylithology that was deposited in the original volcanosedimentary succession, the rocks havethe same age as the enclave as a whole or �3825 Ma.

Whole-rock sulphur isotopic analyses of Akila quartz-pyroxene sample GR9707 re-ported in Farquhar et al., 2000) revealed �33S = +0.12�. Subsequent ion microprobemeasurements of sulphides in Akilia sample GR9707 reported in Mojzsis et al. (2003) re-vealed a large positive �33S anomaly at +1.47 ± 0.14� (2σ). The data also showed apopulation of positive �33S values that ranged between +0.84 to +0.30�. A relativelylarge range (∼12�) of δ34S was also observed (−6.7 to +5.2�) from this sample. Thesedata are not that different from reported values for various ISB metasedimentary rocks. Ina separate study on a different Akilia sample, Whitehouse et al. (2005) measured the multi-ple sulphur isotope composition of quartz-pyroxene sample G91-26 (Nutman et al., 1997)and found no clear evidence for mass-independently fractionated sulphur. Their �33S dataranged from −0.55 to +0.35� and they describe a range in δ34S of +0.02 to +2.45�,substantially smaller than that reported in Mojzsis et al. (2003). Based on their analysis,Whitehouse et al., 2005) concluded that multiple sulphur isotopes could not be used toinfer protolith for highly metamorphosed rocks.

It is likely that the complexity of the Akilia quartz-pyroxene rocks revealed by the vari-ous isotopic and trace element data cited above is manifest in heterogeneities in the sulphurisotopes at the outcrop scale. Further detailed work is needed on these rocks, but fortu-nately, Fe-rich chemical metasedimentary rocks with abundant sulphides turn out to berelatively common in Akilia association enclaves beyond the ‘type-locality’ on Akilia Is-land. Indeed, rocks of the ‘Akilia type’ in mineralogy, bulk composition and antiquity canfound throughout scattered throughout the southern limits of the Itsaq Gneiss Complex of

3The δ5xFe(�) = ((5xFe/54Fe)sample/(5xFe/54Fe)IRMM-014 − 1) × 1000 and IRMM-014 is the Fe-isotope

standard 014 of the Institute of Reference Materials and Measurements.



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West Greenland as well as in other early Archaean terranes of the North American Craton(Cates and Mojzsis, 2006, 2007a).

7.5-7.4. Sulphides from the �3.77 Ga Akilia Association Enclaves on InnersuartuutIsland, West Greenland

Finely-laminated quartz + magnetite ± garnet banded iron-formations from the main is-land of Innersuartuut (approximately 30 km south of Akilia at N63◦50.227′, W51◦41.173′)were sampled for sulphur isotope analysis and reported in Cates and Mojzsis (2006).This rock preserves quartz-magnetite bands that apparently escaped the pervasive strain/recrystallization/silica-mobility recorded by Akilia association units on Akilia Island. Inthis regard, it most resembles samples 119233 from Ugpik, and 119217 from an unspeci-fied field locality on Innersuartuut, as described in McGregor and Mason (1977).

Multiple sulphur isotope geochemistry of this rock, which apparently experienced thesame regional metamorphisms as other Akilia association samples, provides a useful andindependent (from Akilia Island) benchmark for the preservation of mass-independentsulphur in granulite-grade supracrustal enclaves (Fig. 7.5-21). Sulphides are present aspyrrhotite blebs in the quartz-magnetite matrix with fine intergrowths of SiO2. Theanalysed pyrrhotites preserve consistently large �33S values between +1.31 ± 0.19� and+1.49 ± 0.19� (2σ external errors) and record a small range in δ34S of ∼1.5�. Otherferruginous quartz-pyroxene rocks also reported in Cates and Mojzsis (2006) from the In-nersuartuut localities are similar in both mineralogy and gross geochemical composition tothe samples G91-26 of Mojzsis et al. (1996) and GR9707 of Mojzsis et al. (2003). Like theAkilia Island rocks, these samples contain abundant pyrrhotite and occasional chalcopyritewith subhedral habit in a dominantly quartz + hedenbergite matrix. All analysed pyrrhotitein the Innersuartuut samples preserve MIF S isotopes with large �33S values between+1.90 ± 0.25� and +3.15 ± 0.25� (2σ external errors) and range in δ34S by ∼2.5�(Cates and Mojzsis, 2006). Although they record a smaller total range in δ34S comparedto sample GR9707 (∼12�) from Akilia Island (Mojzsis et al., 2003), the Innersuartuutsamples preserve exceptionally large �33S values akin to those reported for ISB bandediron-formation sample 248474 (Baublys et al., 2004; ca. +3.3�) that was reanalysed byWhitehouse et al., 2005). The Innersuartuut rocks are also notable in that they preservea unit of quartz-biotite-garnet schist what contains abundant pyrrhotite (∼1 mode%). Allanalysed sulphides from this particular unit, which Cates and Mojzsis (2006) interpreted tobe a paragneiss of probable pelitic origin, contain MIS sulphur isotopes with well-resolved

Fig. 7.5-21. Compilation plot of �33S vs δ34S for sulphides and sulphates from pre-3.7 Ga metased-iments in the pre-3.77 Ga Isua supracrustal belt and the pre-3.83 Ga Akilia association rocks fromAkilia (island) and Innersuartuut (West Greenland). Data are from Farquhar et al. (2000), Mojzsiset al. (2003), Hu et al. (2003), Whitehouse et al. (2005), Cates and Mojzsis (2006), Papineau andMojzsis (2006), and unpublished data. Compare the different trends in multiple isotope end-membercompositions to the 193-nm results in Fig. 7.5-6.



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�33S values, between +0.67±0.21� and +0.98±0.12� (2σ external errors), and rangein δ34S by ∼3�.

These results for mass-independent sulphur isotopes in amphibolite- to granulite-gradesedimentary rocks from various early Archaean localities in West Greenland underscorethe observation that �33S values in a rock containing accessory sulphur minerals can bepreserved. The simplest explanation for the small spread of δ34S values in these data forthe oldest known rocks of marine sedimentary origin is by re-precipitation of sulphidesduring metamorphic processes. Dilution by invasive fluids (H2S etc.) which carried exoticsulphur does not appear to be an important process in rocks that contain no more than ∼1mode percent sulphur (Ohmoto, 1986). Scatter in �33S values is minimal for each theseunits, which suggests that no significant foreign component of sulphur was added to thesystem since formation as aqueous sediments in the early Archaean ocean.

7.5-8. SUMMARY: BANDED IRON-FORMATIONS, SULPHUR AND LIFE ON THEEARLY EARTH

The Archaean banded iron-formations are regarded solely as a chemical sediments thathave been enriched in metals from volcanic/hydrothermal sources (reviewed in Polat andFrei (2005)) and probably represent part of a relatively deep (>100 m) marine sedimentaryfacies derived from either the photo-oxidation of upwelled Fe(II)-rich seawater to ferro-ferri-oxyhydroxides (Braterman et al., 1983) or biological precipitation and/or oxidationof insoluble iron oxides (e.g., Konhauser et al., 2002). In either of these models, the Fe-precipitates subsequently underwent diagenetic transformation to magnetite. Models forbiological modulation of iron-formation deposition via anoxygenic photosynthetic Fe(II)oxidation, and the nature of microbially-mediated iron-oxide deposition and organic matterrecycling (Kappler et al., 2005), are still in their early stages of development. Cloud (1973)postulated that ferrous iron [Fe(II)aq] oxidation by molecular oxygen after cyanobacteriaevolved on Earth was responsible for the banded iron-formations. However, it has beenargued that the anoxygenic photoautotrophic bacteria are the most ancient type of photo-synthetic organisms (Xiong et al., 2000), and these can readily catalyse Fe(II) oxidationunder anoxic conditions (Widdel et al., 1993). The Fe isotope compositions of bandediron-formations are consistent with the possibility that anoxygenic photoautotrophs suchas purple bacteria played a role in their deposition. Calculations based on experimentallydetermined Fe(II) oxidation rates by anaerobic photoautotrophic bacteria under low-lightregimes representative of ocean depths of a few hundred meters suggest that, even in thepresence of cyanobacteria, anoxygenic phototrophs living beneath a wind-mixed surfacelayer provide the most likely explanation for BIF deposition in a stratified ancient oceanand the absence of Fe in Precambrian surface waters (Kappler et al., 2005). FollowingGross (1980), the Archaean BIFs such as those at Isua and in the Akilia association tendto resemble the ‘Algoma-type’ which are thought to have been deposited in small basinsaround island arcs (Jacobsen and Pimentel-Close, 1988). Enrichment of seawater in fer-rous iron and the transformation to abundant Fe3O4 in marine sediments is considered a


7.5-8. Summary: Banded Iron-Formations, Sulphur and Life on the Early Earth 47

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signature style of sedimentation under anoxic conditions on the early Earth, when Fe(II)concentrations were high in seawater (Holland, 1984; Bjerrum and Canfield, 2002).

Banded iron-formations are strongly enriched in sulphur relative to average Archaeancrust and hydrothermal fluids; the average S composition of Isua iron-formation is takento be close to that of the international banded iron-formation standard IF-G (Govindaraju,1994) at 700 ppm. BIF sulphur was introduced to the water column from both atmosphericand magmatic exhalative sources at time of deposition (Huston and Logan, 2004), andthe major sulphide phases present in the Isua and Akilia rocks are pyrite and pyrrhotite(with minor chalcopyrite and cubanite). As noted by Farquhar and Wing (2003, and refer-ences therein), an important consequence of the model for MIF-carrying aerosol formationin the Archaean atmosphere is that sulphide formation via elemental sulphur reductionby microbial life – perhaps deposited as greigite: Fe3S4, or mackinawite: Fe1−xS (e.g.,Mann et al., 1990) which was later transformed to pyrite: FeS2 and metamorphosed topyrrhotite: ∼Fe7S8 – would be expected to retain positive �33S signatures (Mojzsis et al.,2003) with > 10� spread in δ34S (Canfield and Raiswell, 1999). Sulphide formed frommicrobial sulphate reduction would be expected to preserve negative �33S values with apotentially large spread in δ34S values. Sulphur aerosols provided a ready source of sul-phur to a plethora of microbial environments at the global scale on the early Earth, at issueof course is whether this oxidative chemistry was actually exploited by life, as perhapsFe(II)aq was by early microbial communities for an electron donor (Johnson et al., 2003;Dauphas et al., 2004). Could then the low abundance of oxidants in early Archaean envi-ronments globally have severely limited the early biosphere before the gradual build-up ofoxygen (Papineau and Mojzsis, 2006)? If the Archaean biosphere did in fact modulate BIFdeposition, it still remains to be seen whether it left traces of itself in the sulphur isotopes.

Of the principal elements involved in biological activity (C, N, O, H, S, P, Fe etc.),sulphur is probably the most versatile for geochemistry. As summarised by Goldschmidt(1954), chemists have long recognised that sulphur participates in many of the most im-portant processes in inorganic geochemistry, and serves as a bridge between the inorganicworld and the chemistry contained in the central metabolic machinery of all classes of or-ganisms. Because of its chemical reactivity, large mass difference between multiple stableisotopes (32S, 33S, 34S and 36S), diverse redox chemistry through eight valence states (2−to 6+), and tendency to form both ionic and covalent bonds in a wide variety of miner-als, sulphur and its isotopes have been used to trace igneous, sedimentary, metamorphic,hydrothermal, atmospheric and biological processes since the early days of geochemistry(Thode et al., 1949). New inroads in transition metal stable isotopic systems used to tracechanges in the surface state of the Earth with time, such as Mo (reviewed in Anbar andKnoll (2002)) and Fe (reviewed in Johnson et al. (2003)), also figure prominently in itsphase-space as the various metal-sulphides. How the chemistry of the ‘inorganic bridge’relates to the distribution of inorganic components in the cell (‘metallomics’) and in re-sponse to the evolution of the surface zone as traced by the multiple sulphur isotopes,represents an exciting new horizon that is ripe for exploration. Future work will unitetransition metal isotope systematics with the chemistry of the multiple sulphur isotopespreserved in sedimentary sulphides.



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ACKNOWLEDGEMENTS

Discussions and debates over the years on the various topics presented herein withA.D. Anbar, G.L. Arnold, A. Bekker, P.S. Braterman, R. Buick, N.L. Cates, M. Chaus-sidon, R.N. Clayton, C.D. Coath, N. Dauphas, J. Farquhar, C. Fedo, R. Frei, C.R.L. Friend,J.P. Greenwood, T.M. Harrison, K. Hattori, H.D. Holland, B.S. Kamber, A.J. Kaufman,A. Kappler, J. Karhu, J.L. Kirschvink, M.J. Van Kranendonk, A. Lepland, D. Lowe,J.R. Lyons, C.E. Manning, B. Marty, K.D. McKeegan, S. Moorbath, A.P. Nutman,H. Ohmoto, S. Ono, N.R. Pace, D. Papineau, A.A. Pavlov, M. Rosing, D. Rumble,B. Runnegar, M. Schidlowski, A.K. Schmitt, J.W. Schopf, H. Strauss, M.H. Thiemens,M.J. Whitehouse, B. Wing, Y. Watanabe and M. Van Zuilen have helped to refine andfocus this work. N. Pace at the University of Colorado is thanked for access to his li-brary of Archaea and Bacteria molecular phylogenies. Constructive reviews by N.L. Cates,M.J. Van Kranendonk and two anonymous reviewers helped a great deal to improve themanuscript. Our work on the evolution of sulphur isotopes was supported by the NationalScience Foundation grants EAR9978241 and EAR0228999 to SJM, and since 2000 bythe University of Colorado Center for Astrobiology through a cooperative agreement withthe NASA Astrobiology Institute. Correspondence and request for material should be ad-dressed to [email protected].


mailto:[email protected]

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