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Precambrian Research 183 (2010) 725–737Contents lists available at ScienceDirectPrecambrian Researchjournal homepage: www.elsevier.com/locate/precamres≥3700 Ma pre-metamorphic dolomite formed by microbial mediation in the Isuasupracrustal belt (W. Greenland): Simple evidence for early life?Allen P. Nutman a,b,∗ , Clark R.L. Friend c , Vickie C. Bennett d , David Wright e , Marc D. Norman da School of Earth and Environmental Sciences, University of Wollongong, Wollongong, NSW 2522, Australiab Beijing SHRIMP Centre, Institute of Geology, Chinese Academy of Geological Sciences, 26 Baiwanzhuang Road, Beijing 100037, Chinac 45 Stanway Road, Headington, Oxford OX3 8HU, UKd Research School of Earth Sciences, Australian National University, Canberra ACT 0200, Australiae Department of Geology, University of Leicester, University Road, Leicester LE1 7RH, UKarticleinfoabstractArticle history:Received 19 August 2009Received in revised form 22 July 2010Accepted 13 August 2010Keywords:Isua supracrustal beltEarly lifeDolomiteREE + Y chemistryChemical sedimentary rocksPillow lava intersticesChemical (meta)sedimentary rocks in the amphibolite facies ≥3700 Ma Isua supracrustal belt (W. Greenland)are mostly strongly deformed, so there is only a small chance of the survival of features such asstromatolites or microfossils that would be direct proof of a ≥3700 Ma biosphere. Therefore the searchfor evidence of ≥3700 Ma life in Isua rocks has focused on chemical signatures, particularly C-isotopes.The new approach presented here is based on whole rock chemistry rather than isotopic signatures. Isuachemical sedimentary rocks have Ca–Mg–Fe bulk compositions that coincide with ferroan dolomite –siderite/Fe-oxide mixtures. Most have low Al 2 O 3 , TiO 2 contents (

726 A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737Fig. 1. (a) Map of the Isua supracrustal belt (southern West Greenland) showing location of discussed samples. (b) Detailed map of a small part of the eastern part of theIsua supracrustal belt, showing the geological relationships between altered intermediate (basaltic-andesite?) pillowed volcanic rocks, chemical sedimentary rocks and anoverlying thrust sheet of boninitic amphibolites. All rocks and the thrust were folded prior to intrusion of Ameralik dykes at ca. 3510 Ma.improbable (Appel et al., 2003), and the search for early life has concentratedon how to interpret 13 C-depleted signatures of graphitefrom these rocks. Some workers have concluded that these signaturesare evidence that life was already established by 3700 Ma(Schidlowski et al., 1979; Mozjsis et al., 1996; Rosing, 1999). Unresolvedissues over the interpretation of these signatures includewhether the protoliths of the studied samples are really sedimentaryand whether the graphite formed by inorganic decarbonationreactions instead of reduction of biogenic precursors (Perry andAhmed, 1977; van Zuilen et al., 2002). From one unit of Isuaquartzo-feldspathic rocks of sedimentary origin, Rosing (1999) documentedfine-grained 13 C-depleted graphite which he interpretedto be derived from microbial carbon. However, Schoenburg et al.(2002) raised the possibility that this graphite could be derivedfrom carbonaceous chondrites, particularly as the same rock unitmight contain a tungsten isotopic anomaly of non-terrestrial origin.Thus the C-isotope evidence for Eoarchean life is tantalizing,but not definitive. From another perspective, microbial mediationmay have been necessary for deposition of banded iron formations(BIF) (Cloud, 1973; Garrels et al., 1973) such as those at Isua, perhapsproviding evidence for early life.Based on petrographic, mineralogical and whole rock geochemicalevidence for the existence of low-temperature (premetamorphism)dolomite in Isua sedimentary rocks and pillowinterstices, this paper proposes a simple new line of evidencefor life: that this earliest dolomite formed by ≥3700 Ma bacterialmediation. This is based on the observation that low-temperaturedolomite formation in modern sedimentary and volcanic rocks hasonly been replicated in the laboratory through microbial mediation(Vasconcelos et al., 1995; Roberts et al., 2004; Wright and Wacey,2005).2. Setting of Eoarchaean chemical (meta)sedimentaryrocks in Greenland2.1. Itsaq Gneiss Complex regional geologyThe ca. 3000 km 2 Itsaq Gneiss Complex in the Nuuk region ofW. Greenland is dominated by tonalitic rocks, with inclusions ofamphibolites of mostly island arc tholeiite and boninitic chemicalaffinity, showing that it formed from the products of magmatic arcsat convergent plate boundaries (Nutman et al., 2007, 2009; Dilekand Polat, 2008). These rocks were affected by ductile deformationunder amphibolite to granulite facies conditions in the Eoarchaean(3650–3550 Ma) and again under amphibolite facies conditionsin the Neoarchaean (Nutman et al., 1996). This means that onlya few fortuitous areas escaped most Neoarchaean deformation,allowing rare insights into Eoarchaean geological processes. Thelargest of these low Neoarchaean strain areas is in the north ofthe complex around the ≥3690 Ma Isua supracrustal belt (Fig. 1a;e.g. Moorbath et al., 1973; Allaart, 1976; Nutman et al., 1996,2009, 2002; Rosing et al., 1996; Nutman and Friend, 2009). However,the Isua supracrustal belt contains an Eoarchaean suturezone between a ∼3800 Ma terrane to the south and a ∼3700 Materrane to the north (Fig. 1a; Nutman et al., 1997, 2009, 2002;Crowley, 2003). Thus its volcanic and sedimentary rocks are nowmostly strongly foliated and/or lineated amphibolite facies tectonites(Nutman et al., 1996; Myers, 2001), despite having escapedmuch of the superimposed Neoarchaean deformation. Thus, onlyrarely are sedimentary layering and volcanic structures preservedin the belt (albeit always still deformed under amphibolite faciesmetamorphism; Fig. 2a). It is the Isua rocks that are the focus of thispaper.

A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737 727Isua supracrustal belt carbonates has recently been confirmed bytheir seawater-like REE + Y (SN) (rare earth element and yttriumpost-Archaean-average-shale-normalized (SN)) trace element patterns(Bolhar et al., 2004; Friend et al., 2007). Preservation in somecarbonate rocks of seawater-like REE + Y (SN) patterns strongly suggeststhat their chemistry has little or no modification throughhigh temperature metasomatic processes, as can be observed insome marbles (e.g. Hecht et al., 1999). This does not precludethe possibility that other carbonates entirely of high temperature(600–500 ◦ C) 500 ◦ C) metasomatic origin can also be present (suchas within ultramafic rocks as proposed by Rose et al., 1996; seeFig. 2b).2.3. Isua supracrustal belt carbonate-bearing pillow lavaintersticesFig. 2. (a) Low strain lithon (right) in generally strongly deformed BIF (left). Sedimentarylayering is still preserved in the lithon, although the rock has been deformedand recrystallised under amphibolite facies conditions. Tip of pen for scale in topright of image. (b) Early boundinaged carbonate vein (v) cutting Isua metasomatisedultramafic schists (um) dominated by orthoamphibole. This demonstrates carbonateveining, as seen in orogenic belts of all ages. Pen for scale in the upper middlepart of the image.2.2. Isua supracrustal belt chemical sedimentary rocks – BIFs andcarbonatesIsua magnetite-banded and quartz-rich rocks have been widelyaccepted as having chemical sedimentary protoliths of BIF andchert respectively (Moorbath et al., 1973; Allaart, 1976; Dymekand Klien, 1988; Bolhar et al., 2004; Friend et al., 2007). However,the origin of the carbonate-bearing rocks that are widespreadin some Isua lithological units has been debated. Some workersregard at least some of these rocks as sedimentary in origin (Allaart,1976; Nutman et al., 1984; Dymek and Klien, 1988) whereas othershave proposed that the carbonates are entirely metasomatic(Rose et al., 1996; Rosing et al., 1996). The former interpretationacknowledges the presence of secondary carbonate formation ormobilisation during post-depositional metasomatism. Clear evidenceof this is carbonate veins that cut Isua supracrustal beltrocks (Fig. 2b). Such carbonate mobility is common in highlydeformed and metamorphosed rocks of all ages. However, an originallysedimentary or diagenetic source for at least some of theA focus of this paper is a locality of low strain with preservedpillow structures in metavolcanic rocks first described bySolvang (1999) together with adjacent more deformed carbonaterichchemical metasedimentary rocks (Figs. 1 and 3b). The pillowedvolcanic rocks are basic-intermediate in composition (e.g. analysisG05/22, Table 1). The shapes of the best-preserved pillows (Fig. 3a)show that the volcanic rocks face towards the more deformedchert + calc-silicate + carbonate and BIF rocks that wrap aroundtheir northwestern side (Fig. 3b). The pillows become increasinglybrecciated and traversed by carbonate-bearing veins as the contactis approached, but there is no definitive indication that the contactbetween them is tectonic. Therefore we interpret these chemicalsedimentary rocks to have been deposited on top of the volcanicrocks, and subsequently heterogeneously deformed together(Friend et al., 2007). Evidence supporting this is that the metasedimentaryrocks contain rare detrital/volcanic zircons with ages ofonly 3690 Ma (Nutman et al., 2009; repository Table 1), whereaspillow lava sample G05/22 (GPS 65 ◦ 10.765 ′ N, 49 ◦ 48.211 ′ W –WGS84 datum) from the underlying basic to intermediate metavolcanicrocks yielded a single, prismatic, igneous zircon dated at3709 ± 9Ma(repository Table 1). Therefore the volcanic rocks mustbe marginally older than the overlying metasedimentary rocks(Fig. 4). These volcanic rocks and capping chemical sedimentaryrocks are preserved in an isoclinal fold nose, and are separatedfrom structurally overlying amphibolites of boninitic chemicalaffinity (Polat et al., 2002) by an early isoclinally folded mylonite(Fig. 1b). The interior of the pillow interstices are siliceous and havecarbonate-rich rinds (regrettably not sampled). The interior of thepillows commonly show carbonate alteration, and intact vesiclesare filled by quartz or calcite.3. Chemistry of Isua supracrustal belt chemical(meta)sedimentary rocks and a pillow interstice3.1. Filtering of dataChemical sedimentary rocks can become polluted by terriginousadditions, namely windblown dust, subaqueous detrital input orvolcanic ash (Bolhar et al., 2004 and references therein). The REEabundance of these pollutants is much higher than in chemical sediments(e.g. Bolhar et al., 2004). Therefore only small terriginousadditions will mask the low abundance REE signature related toformation of chemical sedimentary rocks. In the whole rock chemistrythis pollution is readily detected by increases in Al 2 O 3 andTiO 2 . Hence here we have applied an arbitrary filter of

Table 1Whole rock geochemistry.Sample no JHA170728 IF-G# G05/17 G07/22 G07/08 G93/28* G04/54* G04/55* G04/85* G91/75* G05/22 G05/23 G91/25* G91/26R*Sampletype andprotolithdolomiterock metadolostonechemicalsedimentBIFchemicalsedimentdolomitic-BIFchemicalsedimentdolomitic-BIFchemicalsedimentdolomitic-BIFchemicalsedimentdol. calc-sil.chemicalsedimentdol. calc-sil.chemicalsedimentdol. calc-sil.chemicalsedimentdolomitic-BIFchemicalsedimentchertalteredpillow lavaandesite?pillowintersticechemicalsedimentdolomitic-BIFchemicalsedimentdolomitic-BIFUnit and Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Isua amph-f Akilia gran-f Akilia gran-fgradeLatitude 65 ◦ 05.4 ′ N 65 ◦ 09.869 ′ N 65 ◦ 08.059 ′ N 65 ◦ 08.432 ′ N 65 ◦ 05.62 ′ N 65 ◦ 10.767 ′ N 65 ◦ 10.767 ′ N 65 ◦ 10.277 ′ N 65 ◦ 12.14 ′ N 65 ◦ 10.765 ′ N 65 ◦ 10.765 ′ N 63 ◦ 55.73 ′ N 63 ◦ 55.73 ′ NLongitude 50 ◦ 06.5 ′ W 49 ◦ 48.936 ′ W 50 ◦ 11.464 ′ W 50 ◦ 09.302 ′ W 50 ◦ 09.73 ′ W 49 ◦ 48.227 ′ W 49 ◦ 48.227 ′ W 49 ◦ 48.968 ′ W 49 ◦ 47.40 ′ W 49 ◦ 48.211 ′ W 49 ◦ 48.211 ′ W 51 ◦ 41.08 ′ W 51 ◦ 41.08 ′ WSiO 2 (wt%) 3.19 41.20 66.26 71.28 89.73 71.86 54.04 86.60 53.78 97.07 63.04 85.37 82.97 66.80TiO 2 0.02 0.01 0.01 0.01

A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737 729Fig. 4. 238 U/ 206 Pb versus 207 Pb/ 206 Pb concordia diagram for rare small zircons fromchemical metasedimentary rocks G04/54, G04/85 and IEh50ch/a (Nutman et al.,2009), plus three age determinations on a single prismatic zircon recovered fromvolcanic pillow G05/22. Errors are depicted at the 2 level. See online repositoryTable 1 for U–Pb analytical data of sample G05/22.3.2. REE + Y signaturesFig. 3. (a) Weakly deformed pillow lavas (p) with quartz + carbonate + tremoliteactinoliteinterstices (int) at GPS 65 ◦ 10.765 ′ N, 49 ◦ 48.211 ′ W. Note stronglyweathered carbonate-filled margins and cracks in the pillows. Pen for scale onleft hand side. (b) Layered quartz + carbonate + tremolite-actinolite rocks. SampleG04/54 and G05/55 came from where the person is standing. These rocks displayseawater-like REE + Y (SN) trace element signatures (Friend et al., 2007), and thereforeare derived from chemical sedimentary rocks. However, due to high strainin them, the alternation between quartz and carbonate-rich layers should not beregarded as well-preserved sedimentary layering. (c) Transmitted light photomicrographof pillow interstice sample G05/23. Qtz, quartz; tr, tremolite-actinolite;cc, calcite; cp, chalcopyrite. Field of view ∼2.5 mm across. Phases were identified byelectron microprobe energy dispersive analyses using the JEOL JXA-8800R instrumentat the Chinese Academy of Geological Sciences. Cc + tr would have formed byprogressive reaction between qtz + dol in the pre-metamorphic rock. In this casesilica was present in excess, leaving a dolomite-free assemblage.Dolomite-rich rock JHA170728 (Fig. 5a, Table rock JHA170728(Fig. 5a, Table 1) shows an identical PAAS – normalised (REE + Y (SN) )pattern (and in this case elemental abundances) to IF-G, the geochemicalstandard BIF from Isua (Bolhar et al., 2004). This clearlyshows that the common origin of such Isua carbonate rocks andsilica – Fe rocks of accepted sedimentary origin and regarded asa seawater proxy. Regardless of silica content and proportions ofwhole rock Fe to Ca–Mg components, all the Isua samples with lowAl 2 O 3 and TiO 2 show REE + Y (SN) seawater-like patterns with clearpositive Y, Eu, La anomalies and generally discernable Ce anomalies(Fig. 5b and data in Bolhar et al., 2004; Friend et al., 2007; see thelatter for the analytical method of the new data presented here).This REE + Y (SN) pattern is produced by precipitation directlyfrom seawater, and also from low-temperature groundwaters (e.g.Bolhar et al., 2004; Johannesson et al., 2006 and references therein).On the other hand, hydrothermal fluids and carbonate–silica rocksdeposited from such fluids have different REE + Y (SN) signatures(e.g. Bolhar et al., 2005 and references therein). This is demonstratedhere by Precambrian hydrothermal siderite such as sample177886 from the Pilbara (van Kranendonk et al., 2003; analysis4inFig. 5c), plus metasomatic dolomites from the Göpfersgrüntalc deposit (Hecht et al., 1999) and modern oceanic hydrothermalbrines (Douville et al., 1999; Fig. 6). In the Göpfersgründeposit Hecht et al. (1999) analysed progressive generations ofdolomites, from regional dolomitic marble to early hydrothermalvein dolomite and later hydrothermal vein dolomite (analyses 1–3respectively in Fig. 5c). The regional dolomitic marble still displaysa weak Y anomaly and a downward-bowed PAAS-normalisedsignature for the light rare earth elements. In the successivegenerations of hydrothermal dolomite, the positive Y anomalyand downward-bowed PAAS-normalised signature for the lightrare earth elements disappear (Fig. 5c). The Pilbara hydrothermalsiderite of van Kranendonk et al. (2003) also displays neither a positiveY anomaly nor a seawater-like PAAS-normalised signature forthe light rare earth elements (analysis 4 in Fig. 5c). Thus an impor-

730 A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737Fig. 6. REE=Y (SN) plot for Isua pillow lava G05/22, pillow interstice G05/23 and Isuacarbonate-rich sedimentary rocks G04/54 (Friend et al., 2007), Pilbara Warrawoonadolomite-rich pillow interstice is from Yamamoto et al. (2004), Mid-Atlantic Ridgehydrothermal fluid is from Bau and Dulski (1999). Also shown for reference aremetasomatic dolomites of Hecht et al. (1999).Fig. 5. PAAS (Post-Archean-Australian-Shale; McClennan, 1989) normalized REE + Ypatterns (REE = Y (SN) ). (a) IF-G Isua BIF (Bolhar et al., 2004) and metadolomiteJHA170728 (this paper), modern Mg-calcite microbialite from Heron Island,Great Barrier Reef (Webb and Kamber, 2000) and a Warrawoona ∼3.45 Ga lowtemperaturedolomitized stromatolite (van Kranendonk et al., 2003). (b) Isuadolomitic calc-silicate rocks and BIF with a dolomitic component. (c) Metasomaticdolomites and host dolomitic marble (Hecht et al., 1999) and Pilbara ∼3.45 Gahydrothermal siderite deposited from hydrothermal waters (van Kranendonk et al.,2003).tant feature is that hydrothermal fluids and carbonates precipitatedfrom them show no Y anomaly (or it is extremely muted – Fig. 5c).Conversely this is a consistent feature of seawater, some groundwatersand the sedimentary proxies deposited from them. Analyticallythe Y anomaly is particularly useful, given the higher abundanceof Y and the neighboring reference Dy and Ho, compared to theLREE that in seawater proxies occur at an order of magnitude lowerabundances.Thus for the Isua samples we consider here, their REE + Y (SN)signature indicates pre-metamorphic low-temperature depositionrather than from hydrothermal fluids. Deposition could haveoccurred in a range of surficial (directly from seawater) to nearsurfacesettings (from groundwater during diagenesis). Thereforewe contend from the REE + Y data that the Isua supracrustal beltcontains a diverse suite of chemical sediments whose chemistryis controlled by deposition from seawater or during near-surfacediagenetic processes. Their composition ranged from almost puredolomite (JHA170728), to clean chert (G91/75, Friend et al., 2007)to magnetite-rich BIF (IF-G, Bolhar et al., 2004 and G04/85, Friendet al., 2007), and are united in their origin by showing identicalREE+Y (SN) patterns (Fig. 5a and b).3.3. Major element variation of Isua chemical sedimentary rocksThe low Al 2 O 3 and TiO 2 samples including data from the literatureand this study (Table 1), show a range in SiO 2 content from 3%to 97% (e.g. the dolomite-rich sample JHA170728 presented here,and metachert G91/75 in Friend et al., 2007), and when cast intomolecular proportions, mostly form a linear array from Fe to 0.5Ca,0.5Mg in a Fe–Mg–Ca ternary plot (Fig. 7). Even within the samelithological unit (such as the locality discussed in most detail inthe east of the belt where weathered pillow basalts are capped bychemical sedimentary rocks – Fig. 1b), there is a spread in bulkcompositions from close to ferroan dolomite (sample G04/54 – GPS

A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737 731Fig. 7. Molar Ca, Mg, Fe proportions of Isua chemical sedimentary rocks and pillowinterstice G05/23 with low Al 2O 3 and TiO 2. Analyses of samples are in Table 1, Dymekand Klien (1988) and Bolhar et al. (2004). Pilbara interpillow dolomite-rich sampleis from Yamamoto et al. (2004).65 ◦ 10.760 ′ N, 49 ◦ 48.153 ′ W and sample IS625-13B of Dymek andKlien, 1988; Table 1), to other more iron rich compositions (sampleG04/55), to BIF (e.g. sample G04/85 – 65 ◦ 10.310 ′ N, 49 ◦ 49.224 ′ W;Fig. 1b; Table 1).This range of Ca–Mg–Fe compositions is found in other unitsof definite and proposed chemical sedimentary rocks throughoutthe Isua supracrustal belt (Figs. 1 and 7a, Table 1). The ends of thisCa,Mg–Fe array are represented by the sample JHA170728 of massivedolomite rock interlayered with metabasaltic amphibolitesand the geochemical standard BIF sample IF-G (Fig. 7). We note thatalthough JHA170728 contains dolomite, this will have been completelyrecrystallised during superimposed metamorphism. For thepurpose of description, we have divided this compositional arrayinto four groups of rocks; dolomites (JHA170728); dolomitic calcsilicaterocks (e.g. G04/54 and -55), BIF with a dolomitic component(e.g. G07/22, G93/28) and BIF (e.g. IF-G and G04/85; Fig. 7). TheREE+Y (SN) pattern of all of these groups are seawater-like (Fig. 5,Bolhar et al., 2004; Friend et al., 2007), that indicates this Ca,Mg–Fearray is a low-temperature, pre-metamorphic, non-metasomaticfeature.We interpret these compositions to have originally containedboth iron rich phases (iron hydroxides and/or siderite) + dolomite,related either to a diagenetic phenomenon or to high strain telescopingof chemical sedimentary rock layers of different Fe/Mgratio into the scale of a single hand specimen. Similar compositionalranges are found in younger, better-preserved (lowermetamorphic grade and undeformed) Precambrian chemical sedimentarysequences such as the Transvaal Supergroup of southernAfrica, where there is both sedimentary interlayering of BIF, chertsand carbonate rocks, and diagenetic dolomite growth (Beukes,1987).3.4. Chert + dolomite protolith of an Isua pillow intersticeSample G05/23 (GPS 65 ◦ 10.765 ′ N, 49 ◦ 48.211 ′ W) represents theinterior of a pillow interstice. It is devoid of a foliation, andin decreasing modal abundance consists of quartz + tremoliteactinolite+ calcite with accessory chalcopyrite in textural equilibrium(Fig. 3c). It was taken from metabasaltic rocks underlyingthe chemical sedimentary unit represented by samples G04/54 andG04/55. It has low Al 2 O 3 (0.28 wt%) and TiO 2 (

732 A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737perature metamorphism, metasomatism and ductile deformation.It demonstrates that dolomite precipitation was occurring in differentlow-temperature settings, whose modern analogs involve theaction of anaerobic bacteria. There is also strong independent evidencefor life in the Warrawoona sequences including preservedmicrofossils and undisturbed stable isotope signatures (reviewedby Schopf, 2006).4.2. Significance of REE + Y (SN) trace element patterns in IsuadolomitesThe seawater-like REE + Y pattern recurs in chemical sedimentaryrocks throughout the geological record from Isua in theEoarchaean (this paper and Bolhar et al., 2004; Friend et al., 2007)up to the Holocene, where it is present in low-temperature settingssuch as microbialite carbonate cements (Webb and Kamber,2000) or groundwaters (Johannesson et al., 2006). Therefore, wecontend that the Isua supracrustal belt contains a diverse suite ofchemical sediments whose trace element chemistry and the presenceof dolomite in the protoliths was controlled by seawater ornear-surface diagenetic processes.4.3. Low-temperature dolomite formation mechanisms4.3.1. BackgroundDolomite, broadly defined here to as including ferroan varietieswith some substitution of Mg 2+ by Fe 2+ , is a commonlow-temperature sedimentary and diagenetic mineral in carbonatesediments from sabkha to deep ocean environments (see Wrightand Wacey, 2005 and references therein) and occurs in intersticeswithin basalts (Burns et al., 2000; Yamamoto et al., 2004). Despiteover two centuries of intense research, numerous low-temperaturephysico-chemical experiments, including one running for >30 yearsat 1000 times oversaturation, have failed to precipitate dolomite(Machel and Mountjoy, 1986; Hardie, 1987; Wright, 1997; Purseret al., 1994; Land, 1998; Wright, 2000; Wright and Wacey, 2004).This is all the more surprising given that thermodynamic considerationsdictate that in the marine environment, not onlyshould dolomite precipitate spontaneously, but any solid calciumcarbonate should also be immediately dolomitized (Lippmann,1973). However, many observed chemical trends in saline solutionsdo not follow thermodynamic predictions (e.g. Kelleher andRedfern, 2002). Although seawater is supersaturated with respectto dolomite, no spontaneous precipitation of the mineral has everbeen observed, and it is extremely doubtful that dolomite couldhave spontaneously precipitated from normal seawater in the past(e.g. Wright, 2000; Wright and Altermann, 2000). Precipitation ofdolomite from a fluid or by replacement of a CaCO 3 precursor hasonly been produced experimentally at or near hydrothermal conditions(e.g. Gaines, 1980; Lumsden et al., 1995; Tribble et al., 1995;Land, 1998), yet vast amounts of sedimentary (low-temperature)dolomite are present in the geological record, especially in thePrecambrian.4.3.2. Solution chemistry of dolomiteClearly, the precipitation of dolomite from supersaturated seawatersolution under earth-surface conditions is inhibited in someway, and research has shown that this can be attributed to molecularkinetics, which affect the behaviour of the component ions ofdolomite in solution. These kinetic barriers have been identified as:(1) The high hydration energy of the Mg 2+ ion (Lippmann, 1973;Dasent, 1982; Slaughter and Hill, 1991). Cations in aqueoussolutions form hydration shells, which enhance solubility. Themagnesium cation is strongly attached to the oxygen atoms ofsix water molecules that act as electric dipoles whose removalrequires energy.(2) The extremely low concentration and activity of CO 3 2− (e.g.Garrels and Thompson, 1962; Lippmann, 1973; Slaughter andHill, 1991).(3) The presence of even very low concentrations of sulphate,which causes ion complexing including the formation ofstrongly bonded neutral MgSO 4 0 and CaSO 4 0 ion pairs(Usdowski, 1967; Baker and Kastner, 1981; Kastner, 1984;Morrow and Ricketts, 1988; Slaughter and Hill, 1991; Wacey,2003; Wright and Wacey, 2004, 2005; Wright and Oren, 2005).Both (1) and (3) significantly increase the solubility of the Mg 2+ion, which significantly in experiments is always the last to precipitateout from evaporated brines (e.g. Borchert and Muir, 1964,personal observation). In saline solutions, ion pairs form due toshort-range interactions of adjacent ions, attracted by coulombicforces. This complexing reduces the ions’ activities below theirmolarities, making precipitation of carbonate minerals unlikely.For example more than 90% of total CO 3 2− in seawater is complexedwith hydrated metal cations, chiefly magnesium (Garrelsand Christ, 1965).4.3.3. Sulphate and sulphate reducing bacteriaThe requirement to overcome these inhibitors for inorganicprecipitation led some researchers to investigate settings wheremodern sedimentary dolomite is forming, specifically to understandhow the kinetic barriers might be overcome in the naturalenvironment. A common link was subsequently observed betweenmicrobial mediation, organic degradation, raised carbonate alkalinity,removal or absence of sulphate and dolomite formation in bothmodern and ancient sediments (e.g. Gieskes et al., 1982; Compton,1988; Vasconcelos et al., 1995; Vasconcelos and McKenzie, 1997;Wright, 1997, 1999, 2000; Warthmann et al., 2000; van Lith et al.,2003; Roberts et al., 2004; Wright and Wacey, 2004, 2005; Wrightand Oren, 2005; Wacey et al., 2007; Kenward et al., 2009; Oliveri etal., 2010).Sulphate is an abundant component of seawater, and formsneutral ion pairs with metal cations, enhancing their solubility;moreover the proportion of ion pairs increases with ionic concentration.The work of e.g. Walter (1986) has shown that sulphatecan inhibit calcite precipitation, and this is entirely consistent withdolomite inhibition – the same kinetics are involved, with sulphateforming neutral ion pairs with Ca 2+ , but the hydration shell surroundingthe Ca 2+ ion is less tightly bound than that surroundingthe Mg 2+ ion (1926 kJ per mole at infinite dilution).Sulphate reducing bacteria (SRB) can remove all kineticinhibitors to dolomite formation through dissociation of MgSO 40ion pairs as a part of their metabolic activity, the formation ofcarbonate ions through buffering of bicarbonate during organicdiagenesis (Slaughter and Hill, 1991; Wright, 1999), and the disruptionof hydration shells at electrically charged surfaces of degradedorganic material (Kafkafi et al., 1967; Slaughter and Hill, 1991;Balarew et al., 2001). Charged cell walls attract both Mg 2+ andCa 2+ cations, with subsequent adsorption of carbonate ions andthe formation of fine-grained carbonates, and with extracellularpolysaccharides serve as nucleation centres for dolomite precipitation(e.g. van Lith et al., 2003).Hardie (1987) used the presence of sulphate in certain lakewaters to argue against the effectiveness of sulphate as an inhibitorto dolomite formation. However, the presence of sulphate is necessaryfor SRB to metabolise, and dolomite is formed not in aerobiclake waters but in anoxic conditions associated with the zone ofmicrobial sulphate reduction. The experiments of Usdowski (1967),Baker and Kastner (1981) and Morrow and Ricketts (1988), providecompelling evidence for sulphate inhibition of dolomite precipi-

A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737 733tation. Sulphate is of course present in seawater, but in moderndeep-sea organic-rich sediments, where continual diffusion of seawaterSO 4 2− feeds and controls populations of SRB, dolomite formsin sulphate-free anoxic interstitial waters immediately beneath thezone of bacterial sulphate reduction (Gieskes et al., 1982; Compton,1988). Where sulphate reduction was absent in deep-sea samples,no dolomite was reported (Gieskes et al., 1982). These observationsprovide the necessary evidence that the removal of sulphateby SRB from a solution in which it was originally present can leadto dolomite formation.Observations and data reported in Wright (1999, 2000), Wacey(2003), Wacey et al. (2007) and in experiments by Wright andWacey (2004, 2005) also demonstrate that primary dolomite formswhere bacterial sulphate reduction has removed sulphate in thepresence of organic matter. For example, high counts of SRBrecorded from Coorong distal ephemeral lake sediments (up to of3.74 × 10 6 /ml in dolomitic lakes) and enrichment of 34 S in residuallakewaters, indicate flourishing microbial populations and bacterialfractionation of sulphate during sulphate reduction. Sulphateconcentrations were extremely high (>20,000 mg/l) in the lakewatersamples, but declined dramatically and progressively withdepth through the sulphate reduction zone in dolomitic sedimentpore-water samples. By the end of the evaporative cycle, sulphatewas entirely removed. These observations, data and experimentsindicate that it is bacterial sulphate reduction and the consequentbiochemical interactions that drive dolomite formation.As long ago as 1928, Georgii Nadson, a Soviet biologist, recognizedan association between microbial mediation and carbonate(including dolomite) formation, associated with bacterial sulphatereduction. Other early studies in which bacteria were implicated indolomite formation are those by Neher (1959), Neher and Rohrer(1958) and Mansfield (1980). Oppenheimer and Master (1964)reported dolomite formation in association with ‘algal mats’ whileDavies and Ferguson (1975) established a ‘causal relationship’between dolomite and decaying organic matter in experiments.Methanogens and dissimilatory iron reducing bacteria have alsobeen strongly implicated in dolomite formation (e.g. Roberts et al.,2004).4.3.4. Abiological mechanismsThe organogenic approach dating back to Nadson was supplantedby a series of theoretical physico-chemical models thatbecame popular but remained unsupported by experimental evidence(for reviews, see Hardie, 1987; Tucker and Wright, 1990;Wright, 2000). When interpreting dolomite formation in the contextof these ‘conventional’ models, consideration of fundamentalchemical constraints has often been avoided or neglected. Mostof these models involve dolomitization of limestone by circulatingbrines or fluids operating in specific hydrologic systems (e.g.Adams and Rhodes, 1960; Badiozamani, 1973; Folk and Land, 1975;Machel and Mountjoy, 1987).One model commonly referred to as the ‘magnesium pump’ wasinvoked to explain massive dolomitization of carbonate platformsby the circulation or tidal pumping of normal seawater through carbonateplatforms (e.g. Carballo et al., 1987; Land, 1991). However,seawater is unable to act as a dolomitizing solution because of thehigh enthalpy of hydration of the magnesium ion, the low activityof the carbonate ion and the presence of sulphate – all effectiveinhibitors to dolomite formation. Lippmann (1973) argued that fordolomitization of calcite to even begin in normal seawater, all Ca 2+must first be removed by precipitation, and a significant increase incarbonate alkalinity would be required for it to proceed – changesin chemical composition that are impossible to maintain in normalseawater.A similar case can be made against the mixing zone or ‘Dorag’model (Badiozamani, 1973), in which mixed marine and meteoricwaters produce a solution both undersaturated in calcite and supersaturatedin dolomite. Subsequent high Mg:Ca ratios are unlikelyto lead to dolomitization because there is no mechanism to breachthe high hydration barrier of the magnesium ion, nor increase carbonateion activities. This model also predicts that dolomitizationwould be concomitant with the dissolution of calcite, but as Hardie(1987) reports, this does not occur in modern mixing zones. Smithet al. (2003) provides an example of dolomite formation by SRB ina modern mixing zone.Overriding kinetic problems are also inherent in models invokingan evaporative mechanism, including seepage reflux (Adamsand Rhodes, 1960) and evaporative pumping (Hsu and Siegenthaler,1969). Lippmann (1973) showed that evaporation would not favourdolomitization despite consequent high Mg:Ca ratios, because ofthe accompanying decrease in the activity of the CO 3 2− free ion. Thealready low activity of the CO 3 2− ion would be further suppresseddue to its complexing to form neutral ion pairs, giving MgCO 3 0 ,while Mg 2+ would also complex with sulphate, as MgSO 4 0 .Experimental evaporation of brines in beakers has not produceddolomite in the sequence of precipitated minerals (e.g. Borchertand Muir, 1964, personal observation) clearly indicating that simpleevaporation does not produce dolomite. Proposals that dolomite isan evaporative mineral (e.g. McKenzie et al., 1980), are thus shownto be untenable. Where dolomite occurs beneath the Abu Dhabisabkha, it does not cross-cut facies, as might be expected fromseepage reflux or evaporative pumping, but is located within horizontalbeds associated with cyanobacterial mats (McKenzie et al.,1980; Wright, 2000) – suggesting that microbial mediation may beinvolved. More recent work by Bontognali et al. (2010) attributesthe Abu Dhabi dolomite formation to a microbially mediated origin.4.3.5. Documentation of anaerobic microbial mediationIn contrast, anaerobic microbial mediation has been observedto cause the precipitation of dolomite in modern sedimentsand groundwater-basalt systems, with several different metabolicpathways and waters of different composition (Moore et al., 2004;Roberts et al., 2004; Douglas, 2005; Wright and Oren, 2005 and referencestherein; Wright and Wacey, 2005). These cases of dolomiteformation have been replicated under controlled conditions inthe laboratory (e.g. Vasconcelos et al., 1995; Roberts et al., 2004;Wright and Wacey, 2004, 2005), making a compelling case thatmicrobial mediation is essential for low-temperature precipitationof dolomite in saline solutions (Douglas, 2005; Wright and Oren,2005; Wacey et al., 2007).The idea that dolomite needs thousands or even millions ofyears to form is manifestly false, since recent experiments haveconclusively demonstrated that dolomite can be formed in daysthrough microbial mediation (Wright and Wacey, 2004, 2005). Onthe antiquity of microbial sulphate reduction, Shen et al. (2001)and Shen and Buick (2004) have used sulphur isotope data to arguethat sulphate reducing microbes had evolved by the early Archaean.The large spread of 34 S values of microscopic pyrites aligned alonggrowth faces of former gypsum crystals in the ∼3.47-Ga North Polebarite deposit of northwestern Australia provide the oldest evidenceof microbial sulphate reduction and the earliest indicationof a specific microbial metabolism. Microbial ecosystems dominatedthe depositional environments of Archaean, Proterozoic andsome Phanerozoic carbonate platforms, in which sulphate reduction,organic diagenesis and dolomite formation were widespread(e.g. Wright and Altermann, 2000; Shen et al., 2005; Altermann etal., 2006; Gandin and Wright, 2007). Such ecosystems thus actednot only as the ‘engine house’ of early carbonate production, butalso provide a ‘process analogue’ for dolomite formation.In conclusion, although seawater is supersaturated with respectto dolomite, the mineral is prevented from precipitating undernormal earth-surface conditions by kinetic inhibitors. Numer-

A.P. Nutman et al. / Precambrian Research 183 (2010) 725–737 735interpretation of the process of low-temperature dolomite growthin some Isua rocks is correct, then by 3700 Ma, life was alreadywell established in a range of ecological niches, including in asub-surface volcanic pillow interstice environment. Such an environmentwould afford life protection from any surface-sterilizingmassive impacts at that early stage of Earth’s history (summarizedby Koeberl, 2006), and from whence surficial environments couldbe readily colonized again.5. Conclusions(1) Dolomite was a pre-metamorphic mineral in some Isua chemicalsedimentary rocks and pillow interstices.(2) The Isua dolomite (e.g. sample JHA170728) has a REE + Y (SN)seawater-like signature identical to international BIF standardIF-G from Isua, showing it formed in a pre-metamorphic lowtemperatureenvironment, during deposition or diagenesis.(3) Recent low-temperature dolomite precipitation in sedimentarysystems and intra-basalt ground waters apparently requiresmicrobial mediation, and this has been replicated under controlledlaboratory conditions. On the other hand, years ofexperimentation has failed to precipitate low-temperaturedolomite by abiotic means. Given our results and theseobservations, we propose that geochemical evidence for premetamorphicIsua low-temperature dolomite is simple, directevidence for microbial life by 3700 Ma.(4) If some Akilia association BIF with a dolomitic component andseawater-like REE + Y (SN) signatures really are ≥3850 millionyears old, then the dolomite component in these rocks is evidenceof life at the start of Earth’s known sedimentary record.AcknowledgementsThis research was funded by Australian Research Council grantDP0342794, 2006–2009 support from the Chinese Academy ofGeological Sciences and currently an operating grant from theUniversity of Wollongong. The Geological Survey of Denmark andGreenland is thanked for permission to publish this paper.Appendix A. Supplementary dataSupplementary data associated with this article can be found, inthe online version, at doi:10.1016/j.precamres.2010.08.006.ReferencesAdams, J.E., Rhodes, M.L., 1960. Dolomitization by seepage refluxion. 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