≥3700 Ma pre-metamorphic dolomite formed by microbial mediation in the Isua supracrustal belt (W. Greenland): Simple evidence for early life?

Precambrian Research 183 (2010) 725–737

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≥3700 Ma pre-metamorphic dolomite formed by microbial mediation in the Isua supracrustal 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 d a

School of Earth and Environmental Sciences, University of Wollongong, Wollongong, NSW 2522, Australia Beijing SHRIMP Centre, Institute of Geology, Chinese Academy of Geological Sciences, 26 Baiwanzhuang Road, Beijing 100037, China 45 Stanway Road, Headington, Oxford OX3 8HU, UK d Research School of Earth Sciences, Australian National University, Canberra ACT 0200, Australia e Department of Geology, University of Leicester, University Road, Leicester LE1 7RH, UK b c

a r t i c l e

i n f o

Article history: Received 19 August 2009 Received in revised form 22 July 2010 Accepted 13 August 2010

Keywords: Isua supracrustal belt Early life Dolomite REE + Y chemistry Chemical sedimentary rocks Pillow lava interstices

a b s t r a c t Chemical (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 as stromatolites or microfossils that would be direct proof of a ≥3700 Ma biosphere. Therefore the search for 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. Isua chemical sedimentary rocks have Ca–Mg–Fe bulk compositions that coincide with ferroan dolomite – siderite/Fe-oxide mixtures. Most have low Al2 O3 , TiO2 contents (<0.5 and <0.05 wt% respectively) showing minimal contamination from terriginous materials. Identical seawater-like REE + Y shale-normalised trace element signatures with La, Ce, Eu and Y positive anomalies are found in magnetite-rich banded iron formation (BIF – such as the geochemical standard IF-G), dolomite-rich rocks and quartz–carbonate–calcsilicate rocks. Additionally from a rare, small area of low deformation in Isua, there are ∼3700 Ma pillow lava interstices consisting of quartz + tremolite + calcite derived from pre-metamorphic dolomite + silica. Thus the dolomite in the chemical sediments and the pillow interstice was part of the pre-metamorphic assemblage, and was deposited from seawater and/or low-temperature groundwater (as shown by the REE + Y chemistry). Therefore, at least some Isua carbonate rocks are sedimentary or diagenetic in origin rather than being formed by metasomatism at 600–500 ◦ C as proposed by Rose et al. (1996. American Journal of Science 296, 1004–1044). Low-temperature dolomite formation in modern sediments (sabkha to deep ocean) and its deposition from low-temperature groundwater within basalts has only been directly observed in the field and replicated in laboratory experiments through anaerobic microbial mediation. Therefore, microbial mediation appears to be essential for the formation of low-temperature dolomite. From this, we propose that the evidence for the formation of low-temperature pre-metamorphic dolomite in Isua prior to metamorphism provides a new, simple, and relatively direct line of evidence for ≥3700 Ma life. © 2010 Elsevier B.V. All rights reserved.

1. Introduction Widely accepted evidence for Paleoarchean life is found in essentially undeformed and non-metamorphosed 3500–3400 Ma sedimentary rocks of the Pilbara region, Western Australia in the form of stromatolites and probably some microfossils, in association with stable isotope signatures consistent with a biotic origin (Schopf, 2006; van Kranendonk, 2007 and references therein). In

∗ Corresponding author at: Department of Earth and Environmental Sciences, University of Wollongong, Wollongong, NSW, 2515, Australia. Tel.: +61 02 4298 1347. E-mail address: [email protected] (A.P. Nutman). 0301-9268/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2010.08.006

contrast, all pre-3500 Ma (Eoarchaean) rocks have been metamorphosed to at least amphibolite facies conditions and are mostly strongly deformed (Nutman, 2006), which makes it much harder to find convincing evidence of early life in them. The ≥3700 Ma Isua supracrustal belt in the Itsaq Gneiss Complex of southern West Greenland contains the largest and best-preserved example of ancient terrestrial sedimentary and volcanic rocks in the world (Moorbath et al., 1973; Allaart, 1976; Nutman et al., 1996; Rosing et al., 1996; Nutman and Friend, 2009). Since the realisation of the belt’s great age in the early 1970s, the Isua rocks have remained the focus of the search for evidence of earliest life. Because of the deformation and amphibolite facies metamorphism of the Isua rocks, the survival of ≥3700 Ma microfossils in them (Pflug and Jaeschke-Boyer, 1979) has been regarded as

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Fig. 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 the Isua supracrustal belt, showing the geological relationships between altered intermediate (basaltic-andesite?) pillowed volcanic rocks, chemical sedimentary rocks and an overlying 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 concentrated on how to interpret 13 C-depleted signatures of graphite from these rocks. Some workers have concluded that these signatures are evidence that life was already established by 3700 Ma (Schidlowski et al., 1979; Mozjsis et al., 1996; Rosing, 1999). Unresolved issues over the interpretation of these signatures include whether the protoliths of the studied samples are really sedimentary and whether the graphite formed by inorganic decarbonation reactions instead of reduction of biogenic precursors (Perry and Ahmed, 1977; van Zuilen et al., 2002). From one unit of Isua quartzo-feldspathic rocks of sedimentary origin, Rosing (1999) documented fine-grained 13 C-depleted graphite which he interpreted to be derived from microbial carbon. However, Schoenburg et al. (2002) raised the possibility that this graphite could be derived from carbonaceous chondrites, particularly as the same rock unit might 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 mediation may have been necessary for deposition of banded iron formations (BIF) (Cloud, 1973; Garrels et al., 1973) such as those at Isua, perhaps providing evidence for early life. Based on petrographic, mineralogical and whole rock geochemical evidence for the existence of low-temperature (premetamorphism) dolomite in Isua sedimentary rocks and pillow interstices, this paper proposes a simple new line of evidence for life: that this earliest dolomite formed by ≥3700 Ma bacterial mediation. This is based on the observation that low-temperature dolomite formation in modern sedimentary and volcanic rocks has only 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)sedimentary rocks in Greenland 2.1. Itsaq Gneiss Complex regional geology The ca. 3000 km2 Itsaq Gneiss Complex in the Nuuk region of W. Greenland is dominated by tonalitic rocks, with inclusions of amphibolites of mostly island arc tholeiite and boninitic chemical affinity, showing that it formed from the products of magmatic arcs at convergent plate boundaries (Nutman et al., 2007, 2009; Dilek and Polat, 2008). These rocks were affected by ductile deformation under amphibolite to granulite facies conditions in the Eoarchaean (3650–3550 Ma) and again under amphibolite facies conditions in the Neoarchaean (Nutman et al., 1996). This means that only a few fortuitous areas escaped most Neoarchaean deformation, allowing rare insights into Eoarchaean geological processes. The largest of these low Neoarchaean strain areas is in the north of the 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 suture zone between a ∼3800 Ma terrane to the south and a ∼3700 Ma terrane to the north (Fig. 1a; Nutman et al., 1997, 2009, 2002; Crowley, 2003). Thus its volcanic and sedimentary rocks are now mostly strongly foliated and/or lineated amphibolite facies tectonites (Nutman et al., 1996; Myers, 2001), despite having escaped much of the superimposed Neoarchaean deformation. Thus, only rarely are sedimentary layering and volcanic structures preserved in the belt (albeit always still deformed under amphibolite facies metamorphism; Fig. 2a). It is the Isua rocks that are the focus of this paper.

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Isua supracrustal belt carbonates has recently been confirmed by their seawater-like REE + Y(SN) (rare earth element and yttrium post-Archaean-average-shale-normalized (SN)) trace element patterns (Bolhar et al., 2004; Friend et al., 2007). Preservation in some carbonate rocks of seawater-like REE + Y(SN) patterns strongly suggests that their chemistry has little or no modification through high temperature metasomatic processes, as can be observed in some marbles (e.g. Hecht et al., 1999). This does not preclude the possibility that other carbonates entirely of high temperature (600–500 ◦ C) 500 ◦ C) metasomatic origin can also be present (such as within ultramafic rocks as proposed by Rose et al., 1996; see Fig. 2b).

2.3. Isua supracrustal belt carbonate-bearing pillow lava interstices

Fig. 2. (a) Low strain lithon (right) in generally strongly deformed BIF (left). Sedimentary layering is still preserved in the lithon, although the rock has been deformed and recrystallised under amphibolite facies conditions. Tip of pen for scale in top right of image. (b) Early boundinaged carbonate vein (v) cutting Isua metasomatised ultramafic schists (um) dominated by orthoamphibole. This demonstrates carbonate veining, as seen in orogenic belts of all ages. Pen for scale in the upper middle part of the image.

A focus of this paper is a locality of low strain with preserved pillow structures in metavolcanic rocks first described by Solvang (1999) together with adjacent more deformed carbonaterich chemical metasedimentary rocks (Figs. 1 and 3b). The pillowed volcanic rocks are basic-intermediate in composition (e.g. analysis G05/22, Table 1). The shapes of the best-preserved pillows (Fig. 3a) show that the volcanic rocks face towards the more deformed chert + calc-silicate + carbonate and BIF rocks that wrap around their northwestern side (Fig. 3b). The pillows become increasingly brecciated and traversed by carbonate-bearing veins as the contact is approached, but there is no definitive indication that the contact between them is tectonic. Therefore we interpret these chemical sedimentary rocks to have been deposited on top of the volcanic rocks, and subsequently heterogeneously deformed together (Friend et al., 2007). Evidence supporting this is that the metasedimentary rocks contain rare detrital/volcanic zircons with ages of only 3690 Ma (Nutman et al., 2009; repository Table 1), whereas pillow lava sample G05/22 (GPS 65◦ 10.765 N, 49◦ 48.211 W – WGS84 datum) from the underlying basic to intermediate metavolcanic rocks yielded a single, prismatic, igneous zircon dated at 3709 ± 9 Ma (repository Table 1). Therefore the volcanic rocks must be marginally older than the overlying metasedimentary rocks (Fig. 4). These volcanic rocks and capping chemical sedimentary rocks are preserved in an isoclinal fold nose, and are separated from structurally overlying amphibolites of boninitic chemical affinity (Polat et al., 2002) by an early isoclinally folded mylonite (Fig. 1b). The interior of the pillow interstices are siliceous and have carbonate-rich rinds (regrettably not sampled). The interior of the pillows commonly show carbonate alteration, and intact vesicles are filled by quartz or calcite.

2.2. Isua supracrustal belt chemical sedimentary rocks – BIFs and carbonates Isua magnetite-banded and quartz-rich rocks have been widely accepted as having chemical sedimentary protoliths of BIF and chert respectively (Moorbath et al., 1973; Allaart, 1976; Dymek and Klien, 1988; Bolhar et al., 2004; Friend et al., 2007). However, the origin of the carbonate-bearing rocks that are widespread in some Isua lithological units has been debated. Some workers regard at least some of these rocks as sedimentary in origin (Allaart, 1976; Nutman et al., 1984; Dymek and Klien, 1988) whereas others have proposed that the carbonates are entirely metasomatic (Rose et al., 1996; Rosing et al., 1996). The former interpretation acknowledges the presence of secondary carbonate formation or mobilisation during post-depositional metasomatism. Clear evidence of this is carbonate veins that cut Isua supracrustal belt rocks (Fig. 2b). Such carbonate mobility is common in highly deformed and metamorphosed rocks of all ages. However, an originally sedimentary or diagenetic source for at least some of the

3. Chemistry of Isua supracrustal belt chemical (meta)sedimentary rocks and a pillow interstice 3.1. Filtering of data Chemical sedimentary rocks can become polluted by terriginous additions, namely windblown dust, subaqueous detrital input or volcanic ash (Bolhar et al., 2004 and references therein). The REE abundance of these pollutants is much higher than in chemical sediments (e.g. Bolhar et al., 2004). Therefore only small terriginous additions will mask the low abundance REE signature related to formation of chemical sedimentary rocks. In the whole rock chemistry this pollution is readily detected by increases in Al2 O3 and TiO2 . Hence here we have applied an arbitrary filter of <0.5 wt% Al2 O3 and <0.05 wt% TiO2 to identify relatively pure chemical sediments and discard those samples whose compositions are mixtures of chemical and terriginous components.

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Table 1 Whole rock geochemistry. JHA170728 dolomite rock metadolostone

IF-G# chemical sediment BIF

Unit and grade Latitude Longitude

Isua amph-f

Isua amph-f

65 05.4 N 50◦ 06.5 W

SiO2 (wt%) TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5 LOI

3.19 0.02 0.04 0.06 3.47 0.24 22.84 23.51 <0.01 <0.01 n.d. 44.95

Trace elements (p.p.m.) Cr Ni Co Ga Rb Sr Ba Y Zr Hf Nb Ta La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Th U

◦

24 n.d. 2.6 0.2 0.7 141 14 11.6 0.8 <0.20 0.21 <0.03 2 2.84 0.46 1.97 0.67 0.56 0.81 0.11 0.89 0.21 0.58 0.56 0.09 0.07 <0.03



G05/17 chemical sediment dolomiticBIF Isua amph-f ◦

170728 replicate analysis 23 n.d. 2.8 0.2 0.7 141 14 11.7 0.9 <0.14 0.24 <0.03 2.01 2.9 0.46 1.84 0.5 0.52 0.62 0.13 0.82 0.21 0.58 0.53 0.08 0.06 0.03



◦



G93/28* chemical sediment dol. calc-sil.

G04/54* chemical sediment dol. calc-sil.

G04/55* chemical sediment dol. calc-sil.

Isua amph-f

Isua amph-f

Isua amph-f

◦



◦



◦



G04/85* chemical sediment dolomiticBIF Isua amph-f ◦



G91/75* chemical sediment chert

G05/22 altered pillow lava andesite?

G05/23 pillow interstice

Isua amph-f

Isua amph-f

Isua amph-f

◦



◦

G91/25* chemical sediment dolomiticBIF Akilia gran-f

G91/26R* chemical sediment dolomiticBIF Akilia gran-f

65 10.767 N 65 10.767 N 65 10.277 N 65 12.14 N 49◦ 48.227 W 49◦ 48.227 W 49◦ 48.968 W 49◦ 47.40 W

65 10.765 N 65◦ 10.765 N 63◦ 55.73 N 49◦ 48.211 W 49◦ 48.211 W 51◦ 41.08 W

63◦ 55.73 N 51◦ 41.08 W

41.20 0.01 0.02 55.85 n.d. 0.04 1.89 1.55 0.03 0.01 0.06

66.26 0.01 0.19 28.41 n.d. 0.20 1.48 1.64 0.06 0.02 0.02 1.74

71.28 0.01 0.16 20.91 n.d. 0.20 2.21 2.46 0.04 0.01 0.01 2.73

89.73 <0.01 0.13 8.35 n.d. 0.15 0.88 0.53 0.01 <0.01 0.01 0.09

71.86 0.01 0.31 12.42 n.d. 0.42 3.08 6.21 0.03 0.04 0.02 5.53

54.04 0.07 1.72 12.35 n.d. 0.79 7.16 13.49 0.03 0.53 0.01 9.74

86.60 0.01 0.31 6.92 n.d. 0.32 2.29 2.78 0.01 0.01 0.01 0.64

53.78 0.01 0.19 46.10 n.d. 0.06 2.63 0.23 0.07 0.01 0.05 −3.14

97.07 <0.01 0.16 2.62 n.d. 0.08 0.67 0.94 0.01 <0.01 0.01 −1.58

63.04 0.84 13.36 7.67 n.d. 0.25 2.85 3.17 0.65 5.25 0.13 2.54

85.37 <0.01 0.28 3.15 n.d. 0.23 2.86 4.48 0.03 0.04 0.01 2.47

82.97 0.01 0.42 11.60 n.d. 0.19 3.27 3.58 0.06 0.01 0.02 −2.54

66.80 0.01 0.16 27.55 n.d. 0.33 5.06 2.20 0.06 0.01 0.04 −3.65

4 23 29.0 0.07 0.04 3 1.5 9 1 0.04 0.1 0.2 2.8 4 0.4 1.8 0.4 0.39 0.74

29

13

15

16

3.6

4.2 0.7 0.3 3.4 1.9 4.0 2.7 0.06 0.04 <0.019 1.432 2.144 0.254 1.020 0.292 0.190 0.357 0.046 0.406 0.097 0.283 0.273 0.040 0.02 0.071

1.7 0.3 0.1 1.0 0.3 1.9 0.7 <0.05 0.04 <0.011 0.486 0.886 0.097 0.406 <0.109 0.092 0.166 0.033 0.230 0.050 0.153 0.149 0.027 <0.020 0.033

6.8 0.6 1.56 8.7 7.0 4.6 2.2 0.059 0.078 0.007 2.167 3.314 0.362 1.513 0.335 0.261 0.467 0.075 0.444 0.107 0.323 0.265 0.043 0.083 0.033

37 23 3.1 2.0 16.7 18.1 349 3.1 13.5 0.412 0.458 0.028 2.041 3.111 0.351 1.378 0.307 0.263 0.362 0.059 0.353 0.077 0.230 0.211 0.033 0.295 0.064

9 21 2.4 0.4 0.47 9.8 5.0 2.4 2.0 0.046 0.073 0.006 0.517 0.736 0.087 0.395 0.108 0.124 0.200 0.034 0.238 0.056 0.174 0.150 0.023 0.034 0.022

6 18 2.9 0.5 0.13 0.4 1.0 3.2 0.6 0.016 0.087 0.003 1.959 3.026 0.305 1.231 0.242 0.175 0.304 0.050 0.329 0.077 0.263 0.305 0.053 0.032 0.015

13 8 0.7 0.2 0.39 1.6 1.0 0.9 0.5 0.024 0.027 0.006 0.461 0.610 0.063 0.251 0.066 0.056 0.083 0.014 0.087 0.021 0.059 0.060 0.008 0.026 0.012

529 90 13.1 17.8 117.8 30.4 425 20.6 126 3.58 4.647 0.354 15.547 34.18 4.248 17.72 4.029 1.133 0.705 4.047 3.964 0.766 2.155 1.838 0.267 4.41 0.725

19 47 6.1 0.6 1.2 7.0 2.1 4.8 3.7 0.053 0.05 0.006 0.363 0.256 0.037 0.242 0.121 0.022 0.088 0.340 0.630 0.158 0.481 0.421 0.053 0.08 0.596

12 59 5.8 0.7 0.34 3.3 1.4 2.9 3.1 0.105 1.04 0.016 1.103 2.220 0.288 1.293 0.328 0.147 0.411 0.073 0.434 0.088 0.258 0.208 0.032 0.130 0.040

10 172 9.7 0.8 0.13 1.7 0.4 1.2 0.9 0.019 2.94 0.023 0.729 1.230 0.128 0.485 0.096 0.058 0.156 0.026 0.164 0.038 0.115 0.122 0.020 0.080 0.073

0.8 0.2 0.63 0.6 0.09 0.1 0.02

◦

G07/08 chemical sediment dolomiticBIF Isua amph-f

65 09.869 N 65 08.059 N 65 08.432 N 65 05.62 N 49◦ 48.936 W 50◦ 11.464 W 50◦ 09.302 W 50◦ 09.73 W

0.8 4.3 45 4.4 2.8 0.08 0.04 0.02 1.192 1.810 0.210 0.902 0.250 0.175 0.464 0.072 0.511 0.115 0.357 0.332 0.058 0.05 <0.013



G07/22 chemical sediment dolomiticBIF Isua amph-f



Major elements determined by XRF (n.d. = not determined), trace elements determined by Laser Ablation ICPMS of samples fused in Li-borate glass (<, below detection limit). See Friend et al. (2007) for analytical methods, apart from IF-G (Bolhar et al., 2004). Isua = sample from the Isua supracrustal belt; Akilia = sample from the Akilia association on Akilia island, south of Nuuk town; amph-f = maximum amphibolite facies; gran-f = maximum granulite facies metamorphism. Positions are WGS84 GPS readings, apart from Google EarthTM positions for 1975 sample JHA170728 (for which the location is known less accurately), G91/75 and G93/28.

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Sample no Sample type and protolith

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Fig. 4. 238 U/206 Pb versus 207 Pb/206 Pb concordia diagram for rare small zircons from chemical metasedimentary rocks G04/54, G04/85 and IEh50ch/a (Nutman et al., 2009), plus three age determinations on a single prismatic zircon recovered from volcanic pillow G05/22. Errors are depicted at the 2 level. See online repository Table 1 for U–Pb analytical data of sample G05/22.

3.2. REE + Y signatures

Fig. 3. (a) Weakly deformed pillow lavas (p) with quartz + carbonate + tremoliteactinolite interstices (int) at GPS 65◦ 10.765 N, 49◦ 48.211 W. Note strongly weathered carbonate-filled margins and cracks in the pillows. Pen for scale on left hand side. (b) Layered quartz + carbonate + tremolite-actinolite rocks. Sample G04/54 and G05/55 came from where the person is standing. These rocks display seawater-like REE + Y(SN) trace element signatures (Friend et al., 2007), and therefore are derived from chemical sedimentary rocks. However, due to high strain in them, the alternation between quartz and carbonate-rich layers should not be regarded as well-preserved sedimentary layering. (c) Transmitted light photomicrograph of pillow interstice sample G05/23. Qtz, quartz; tr, tremolite-actinolite; cc, calcite; cp, chalcopyrite. Field of view ∼2.5 mm across. Phases were identified by electron microprobe energy dispersive analyses using the JEOL JXA-8800R instrument at the Chinese Academy of Geological Sciences. Cc + tr would have formed by progressive reaction between qtz + dol in the pre-metamorphic rock. In this case silica 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 geochemical standard BIF from Isua (Bolhar et al., 2004). This clearly shows that the common origin of such Isua carbonate rocks and silica – Fe rocks of accepted sedimentary origin and regarded as a seawater proxy. Regardless of silica content and proportions of whole rock Fe to Ca–Mg components, all the Isua samples with low Al2 O3 and TiO2 show REE + Y(SN) seawater-like patterns with clear positive Y, Eu, La anomalies and generally discernable Ce anomalies (Fig. 5b and data in Bolhar et al., 2004; Friend et al., 2007; see the latter for the analytical method of the new data presented here). This REE + Y(SN) pattern is produced by precipitation directly from 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 rocks deposited from such fluids have different REE + Y(SN) signatures (e.g. Bolhar et al., 2005 and references therein). This is demonstrated here by Precambrian hydrothermal siderite such as sample 177886 from the Pilbara (van Kranendonk et al., 2003; analysis 4 in Fig. 5c), plus metasomatic dolomites from the Göpfersgrün talc deposit (Hecht et al., 1999) and modern oceanic hydrothermal brines (Douville et al., 1999; Fig. 6). In the Göpfersgrün deposit Hecht et al. (1999) analysed progressive generations of dolomites, from regional dolomitic marble to early hydrothermal vein dolomite and later hydrothermal vein dolomite (analyses 1–3 respectively in Fig. 5c). The regional dolomitic marble still displays a weak Y anomaly and a downward-bowed PAAS-normalised signature for the light rare earth elements. In the successive generations of hydrothermal dolomite, the positive Y anomaly and downward-bowed PAAS-normalised signature for the light rare earth elements disappear (Fig. 5c). The Pilbara hydrothermal siderite of van Kranendonk et al. (2003) also displays neither a positive Y anomaly nor a seawater-like PAAS-normalised signature for the light rare earth elements (analysis 4 in Fig. 5c). Thus an impor-

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Fig. 6. REE = Y(SN) plot for Isua pillow lava G05/22, pillow interstice G05/23 and Isua carbonate-rich sedimentary rocks G04/54 (Friend et al., 2007), Pilbara Warrawoona dolomite-rich pillow interstice is from Yamamoto et al. (2004), Mid-Atlantic Ridge hydrothermal fluid is from Bau and Dulski (1999). Also shown for reference are metasomatic dolomites of Hecht et al. (1999).

Fig. 5. PAAS (Post-Archean-Australian-Shale; McClennan, 1989) normalized REE + Y patterns (REE = Y(SN) ). (a) IF-G Isua BIF (Bolhar et al., 2004) and metadolomite JHA170728 (this paper), modern Mg-calcite microbialite from Heron Island, Great Barrier Reef (Webb and Kamber, 2000) and a Warrawoona ∼3.45 Ga lowtemperature dolomitized stromatolite (van Kranendonk et al., 2003). (b) Isua dolomitic calc-silicate rocks and BIF with a dolomitic component. (c) Metasomatic dolomites and host dolomitic marble (Hecht et al., 1999) and Pilbara ∼3.45 Ga hydrothermal siderite deposited from hydrothermal waters (van Kranendonk et al., 2003).

tant feature is that hydrothermal fluids and carbonates precipitated from them show no Y anomaly (or it is extremely muted – Fig. 5c). Conversely this is a consistent feature of seawater, some groundwaters and the sedimentary proxies deposited from them. Analytically the Y anomaly is particularly useful, given the higher abundance

of Y and the neighboring reference Dy and Ho, compared to the LREE that in seawater proxies occur at an order of magnitude lower abundances. Thus for the Isua samples we consider here, their REE + Y(SN) signature indicates pre-metamorphic low-temperature deposition rather than from hydrothermal fluids. Deposition could have occurred in a range of surficial (directly from seawater) to nearsurface settings (from groundwater during diagenesis). Therefore we contend from the REE + Y data that the Isua supracrustal belt contains a diverse suite of chemical sediments whose chemistry is controlled by deposition from seawater or during near-surface diagenetic processes. Their composition ranged from almost pure dolomite (JHA170728), to clean chert (G91/75, Friend et al., 2007) to magnetite-rich BIF (IF-G, Bolhar et al., 2004 and G04/85, Friend et al., 2007), and are united in their origin by showing identical REE + Y(SN) patterns (Fig. 5a and b). 3.3. Major element variation of Isua chemical sedimentary rocks The low Al2 O3 and TiO2 samples including data from the literature and this study (Table 1), show a range in SiO2 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 into molecular 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 same lithological unit (such as the locality discussed in most detail in the east of the belt where weathered pillow basalts are capped by chemical sedimentary rocks – Fig. 1b), there is a spread in bulk compositions from close to ferroan dolomite (sample G04/54 – GPS

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Fig. 7. Molar Ca, Mg, Fe proportions of Isua chemical sedimentary rocks and pillow interstice G05/23 with low Al2 O3 and TiO2 . Analyses of samples are in Table 1, Dymek and Klien (1988) and Bolhar et al. (2004). Pilbara interpillow dolomite-rich sample is from Yamamoto et al. (2004).

65◦ 10.760 N, 49◦ 48.153 W and sample IS625-13B of Dymek and Klien, 1988; Table 1), to other more iron rich compositions (sample G04/55), 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 units of definite and proposed chemical sedimentary rocks throughout the Isua supracrustal belt (Figs. 1 and 7a, Table 1). The ends of this Ca,Mg–Fe array are represented by the sample JHA170728 of massive dolomite rock interlayered with metabasaltic amphibolites and the geochemical standard BIF sample IF-G (Fig. 7). We note that although JHA170728 contains dolomite, this will have been completely recrystallised during superimposed metamorphism. For the purpose of description, we have divided this compositional array into four groups of rocks; dolomites (JHA170728); dolomitic calcsilicate rocks (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). The REE + 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–Fe array is a low-temperature, pre-metamorphic, non-metasomatic feature. We interpret these compositions to have originally contained both iron rich phases (iron hydroxides and/or siderite) + dolomite, related either to a diagenetic phenomenon or to high strain telescoping of chemical sedimentary rock layers of different Fe/Mg ratio into the scale of a single hand specimen. Similar compositional ranges are found in younger, better-preserved (lower metamorphic grade and undeformed) Precambrian chemical sedimentary sequences such as the Transvaal Supergroup of southern Africa, where there is both sedimentary interlayering of BIF, cherts and carbonate rocks, and diagenetic dolomite growth (Beukes, 1987). 3.4. Chert + dolomite protolith of an Isua pillow interstice Sample G05/23 (GPS 65◦ 10.765 N, 49◦ 48.211 W) represents the interior of a pillow interstice. It is devoid of a foliation, and in decreasing modal abundance consists of quartz + tremoliteactinolite + calcite with accessory chalcopyrite in textural equilibrium (Fig. 3c). It was taken from metabasaltic rocks underlying the chemical sedimentary unit represented by samples G04/54 and G04/55. It has low Al2 O3 (0.28 wt%) and TiO2 (<0.01 wt%; Table 1), showing it is an infilled void, containing only minor volcanic material. It has Ca–Mg–Fe proportions close to ferroan dolomite (Fig. 6). Given the low amphibolite facies metamorphic grade of the rocks at this locality (derived from biotite-garnet thermometry and garnet zoning studies; Rollinson, 2003), we consider the quartz + tremolite + calcite assemblage to have formed by progres-

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sive metamorphism of silica + dolomite in the protolith, with likely maximum metamorphic temperature of ca. 550 ◦ C. Interstice sample G05/23 has a REE + Y(SN) pattern characterised by positive La and Y anomalies, and a strongly downward-bowed depletion of the light REE compared to the heavy REE (Fig. 6). In these respects the pattern has similarities to dolomitised Pilbara stromatolites, modern microbialites (Figs. 5 and 6), seawater and groundwaters. The only departure from this similarity with ancient seawater-like proxies is its anomalously low Eu abundance. On the other hand, the trace element signature is markedly different from that of the adjacent volcanic rocks (e.g.G05/22; Table 1, Fig. 6), hydrothermal fluids such as 364–35 ◦ C ones at the modern Mid-Atlantic Ridge (Douville et al., 1999; Fig. 6) and from 3.46 Ga Warrawoona Group siderite + jaspilite rocks formed from mixed hydrothermal fluids and shallow seawater (van Kranendonk et al., 2003). The reconstruction of interstice sample G05/23’s original pre-metamorphic silica + dolomite assemblage from its petrography and chemistry then suggests the interstice was filled by low-temperature silica and dolomite precipitated from sea or groundwater (not hydrothermal fluids) prior to the superimposed low amphibolite facies metamorphism.

4. Discussion – biogenic origin of low-temperature dolomite in Isua rocks 4.1. REE + Y(SN) signatures of low-temperature dolomite in post-Eoarchaean sedimentary rocks and basalts In the Phanerozoic, times of extensive dolomite precipitation broadly correlate with diverse indicators of decreased oxygen levels in the atmosphere and oceans (Burns et al., 2000). If this has been generally true throughout Earth’s history, dolomite precipitation should have been widespread in the early Precambrian, where there is diverse and consistent evidence of low oxygen levels (e.g. Kasting, 1993). In accord with this, there is widespread dolomitization in early Precambrian carbonate sedimentary sequences, which has been demonstrated to be diagenetic, and not formed later (Beukes, 1987; Wright and Altermann, 2000). Some of the world’s oldest domains of essentially nonmetamorphosed and undeformed chemical sedimentary rocks and pillow lavas are in the 3460–3430 Ma Warrawoona Group, Pilbara Craton (Western Australia) (van Kranendonk et al., 2007). 3440 Ma biotic stromatolite structures from the Strelley Pool Chert Formation are widely dolomitic. In some cases dolomite forms crystal fans formed under low-temperature diagenetic conditions, originally after aragonite or gypsum (van Kranendonk et al., 2003; see also Gandin et al., 2005; Gandin and Wright, 2007 for examples from the Neoarchaean of South Africa). The Pilbara dolomitic stromatolites all preserve seawater-like REE + Y(SN) patterns with positive Y, Eu, Ce and La anomalies (Fig. 5). These signatures resemble those in modern and Paleozoic microbial carbonates unpolluted by terriginous material (Fig. 5; Webb and Kamber, 2000; Nothdurft et al., 2004) and modern groundwaters (Johannesson et al., 2006). In the stratigraphically higher Euro Basalt Formation of the Warrawoona Group, Yamamoto et al. (2004) report pillow interstices rich in dolomite that formed by replacement of calcite. The REE(SN) pattern of these dolomites (Yamamoto et al. did not report Y abundance) indicates sea- or groundwater was present at dolomitization, and again suggests that it occurred at a low temperature, probably shortly after eruption, as seen in modern oceanic basalts (Rouxel et al., 2008). One modern analogue process could be microbial precipitation of dolomite on basalt in anoxic groundwater (Roberts et al., 2004). Thus 3460–3430 Ma Warrawoona Group rocks afford the oldest glimpse of the ancient marine-biosphere system, where there has been minimal modification by superimposed high tem-

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perature metamorphism, metasomatism and ductile deformation. It demonstrates that dolomite precipitation was occurring in different low-temperature settings, whose modern analogs involve the action of anaerobic bacteria. There is also strong independent evidence for life in the Warrawoona sequences including preserved microfossils and undisturbed stable isotope signatures (reviewed by Schopf, 2006). 4.2. Significance of REE + Y(SN) trace element patterns in Isua dolomites The seawater-like REE + Y pattern recurs in chemical sedimentary rocks throughout the geological record from Isua in the Eoarchaean (this paper and Bolhar et al., 2004; Friend et al., 2007) up to the Holocene, where it is present in low-temperature settings such as microbialite carbonate cements (Webb and Kamber, 2000) or groundwaters (Johannesson et al., 2006). Therefore, we contend that the Isua supracrustal belt contains a diverse suite of chemical sediments whose trace element chemistry and the presence of dolomite in the protoliths was controlled by seawater or near-surface diagenetic processes. 4.3. Low-temperature dolomite formation mechanisms 4.3.1. Background Dolomite, broadly defined here to as including ferroan varieties with some substitution of Mg2+ by Fe2+ , is a common low-temperature sedimentary and diagenetic mineral in carbonate sediments from sabkha to deep ocean environments (see Wright and Wacey, 2005 and references therein) and occurs in interstices within basalts (Burns et al., 2000; Yamamoto et al., 2004). Despite over two centuries of intense research, numerous low-temperature physico-chemical experiments, including one running for >30 years at 1000 times oversaturation, have failed to precipitate dolomite (Machel and Mountjoy, 1986; Hardie, 1987; Wright, 1997; Purser et al., 1994; Land, 1998; Wright, 2000; Wright and Wacey, 2004). This is all the more surprising given that thermodynamic considerations dictate that in the marine environment, not only should dolomite precipitate spontaneously, but any solid calcium carbonate should also be immediately dolomitized (Lippmann, 1973). However, many observed chemical trends in saline solutions do not follow thermodynamic predictions (e.g. Kelleher and Redfern, 2002). Although seawater is supersaturated with respect to dolomite, no spontaneous precipitation of the mineral has ever been observed, and it is extremely doubtful that dolomite could have spontaneously precipitated from normal seawater in the past (e.g. Wright, 2000; Wright and Altermann, 2000). Precipitation of dolomite from a fluid or by replacement of a CaCO3 precursor has only 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 the Precambrian. 4.3.2. Solution chemistry of dolomite Clearly, the precipitation of dolomite from supersaturated seawater solution under earth-surface conditions is inhibited in some way, and research has shown that this can be attributed to molecular kinetics, which affect the behaviour of the component ions of dolomite in solution. These kinetic barriers have been identified as: (1) The high hydration energy of the Mg2+ ion (Lippmann, 1973; Dasent, 1982; Slaughter and Hill, 1991). Cations in aqueous solutions form hydration shells, which enhance solubility. The magnesium cation is strongly attached to the oxygen atoms of

six water molecules that act as electric dipoles whose removal requires energy. (2) The extremely low concentration and activity of CO3 2− (e.g. Garrels and Thompson, 1962; Lippmann, 1973; Slaughter and Hill, 1991). (3) The presence of even very low concentrations of sulphate, which causes ion complexing including the formation of strongly bonded neutral MgSO4 0 and CaSO4 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 Mg2+ ion, which significantly in experiments is always the last to precipitate out from evaporated brines (e.g. Borchert and Muir, 1964, personal observation). In saline solutions, ion pairs form due to short-range interactions of adjacent ions, attracted by coulombic forces. This complexing reduces the ions’ activities below their molarities, making precipitation of carbonate minerals unlikely. For example more than 90% of total CO3 2− in seawater is complexed with hydrated metal cations, chiefly magnesium (Garrels and Christ, 1965). 4.3.3. Sulphate and sulphate reducing bacteria The requirement to overcome these inhibitors for inorganic precipitation led some researchers to investigate settings where modern sedimentary dolomite is forming, specifically to understand how the kinetic barriers might be overcome in the natural environment. A common link was subsequently observed between microbial mediation, organic degradation, raised carbonate alkalinity, removal or absence of sulphate and dolomite formation in both modern 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; Wright and Oren, 2005; Wacey et al., 2007; Kenward et al., 2009; Oliveri et al., 2010). Sulphate is an abundant component of seawater, and forms neutral 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 sulphate can inhibit calcite precipitation, and this is entirely consistent with dolomite inhibition – the same kinetics are involved, with sulphate forming neutral ion pairs with Ca2+ , but the hydration shell surrounding the Ca2+ ion is less tightly bound than that surrounding the Mg2+ ion (1926 kJ per mole at infinite dilution). Sulphate reducing bacteria (SRB) can remove all kinetic inhibitors to dolomite formation through dissociation of MgSO4 0 ion pairs as a part of their metabolic activity, the formation of carbonate ions through buffering of bicarbonate during organic diagenesis (Slaughter and Hill, 1991; Wright, 1999), and the disruption of hydration shells at electrically charged surfaces of degraded organic material (Kafkafi et al., 1967; Slaughter and Hill, 1991; Balarew et al., 2001). Charged cell walls attract both Mg2+ and Ca2+ cations, with subsequent adsorption of carbonate ions and the formation of fine-grained carbonates, and with extracellular polysaccharides serve as nucleation centres for dolomite precipitation (e.g. van Lith et al., 2003). Hardie (1987) used the presence of sulphate in certain lake waters to argue against the effectiveness of sulphate as an inhibitor to dolomite formation. However, the presence of sulphate is necessary for SRB to metabolise, and dolomite is formed not in aerobic lake waters but in anoxic conditions associated with the zone of microbial sulphate reduction. The experiments of Usdowski (1967), Baker and Kastner (1981) and Morrow and Ricketts (1988), provide compelling evidence for sulphate inhibition of dolomite precipi-

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tation. Sulphate is of course present in seawater, but in modern deep-sea organic-rich sediments, where continual diffusion of seawater SO4 2− feeds and controls populations of SRB, dolomite forms in sulphate-free anoxic interstitial waters immediately beneath the zone 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 observations provide the necessary evidence that the removal of sulphate by SRB from a solution in which it was originally present can lead to dolomite formation. Observations and data reported in Wright (1999, 2000), Wacey (2003), Wacey et al. (2007) and in experiments by Wright and Wacey (2004, 2005) also demonstrate that primary dolomite forms where bacterial sulphate reduction has removed sulphate in the presence of organic matter. For example, high counts of SRB recorded from Coorong distal ephemeral lake sediments (up to of 3.74 × 106 /ml in dolomitic lakes) and enrichment of 34 S in residual lakewaters, indicate flourishing microbial populations and bacterial fractionation of sulphate during sulphate reduction. Sulphate concentrations were extremely high (>20,000 mg/l) in the lakewater samples, but declined dramatically and progressively with depth through the sulphate reduction zone in dolomitic sediment pore-water samples. By the end of the evaporative cycle, sulphate was entirely removed. These observations, data and experiments indicate that it is bacterial sulphate reduction and the consequent biochemical interactions that drive dolomite formation. As long ago as 1928, Georgii Nadson, a Soviet biologist, recognized an association between microbial mediation and carbonate (including dolomite) formation, associated with bacterial sulphate reduction. Other early studies in which bacteria were implicated in dolomite 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’ while Davies and Ferguson (1975) established a ‘causal relationship’ between dolomite and decaying organic matter in experiments. Methanogens and dissimilatory iron reducing bacteria have also been strongly implicated in dolomite formation (e.g. Roberts et al., 2004). 4.3.4. Abiological mechanisms The organogenic approach dating back to Nadson was supplanted by a series of theoretical physico-chemical models that became popular but remained unsupported by experimental evidence (for reviews, see Hardie, 1987; Tucker and Wright, 1990; Wright, 2000). When interpreting dolomite formation in the context of these ‘conventional’ models, consideration of fundamental chemical constraints has often been avoided or neglected. Most of these models involve dolomitization of limestone by circulating brines 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’ was invoked to explain massive dolomitization of carbonate platforms by the circulation or tidal pumping of normal seawater through carbonate platforms (e.g. Carballo et al., 1987; Land, 1991). However, seawater is unable to act as a dolomitizing solution because of the high enthalpy of hydration of the magnesium ion, the low activity of the carbonate ion and the presence of sulphate – all effective inhibitors to dolomite formation. Lippmann (1973) argued that for dolomitization of calcite to even begin in normal seawater, all Ca2+ must first be removed by precipitation, and a significant increase in carbonate alkalinity would be required for it to proceed – changes in chemical composition that are impossible to maintain in normal seawater. A similar case can be made against the mixing zone or ‘Dorag’ model (Badiozamani, 1973), in which mixed marine and meteoric

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waters produce a solution both undersaturated in calcite and supersaturated in dolomite. Subsequent high Mg:Ca ratios are unlikely to lead to dolomitization because there is no mechanism to breach the high hydration barrier of the magnesium ion, nor increase carbonate ion activities. This model also predicts that dolomitization would be concomitant with the dissolution of calcite, but as Hardie (1987) reports, this does not occur in modern mixing zones. Smith et al. (2003) provides an example of dolomite formation by SRB in a modern mixing zone. Overriding kinetic problems are also inherent in models invoking an evaporative mechanism, including seepage reflux (Adams and Rhodes, 1960) and evaporative pumping (Hsu and Siegenthaler, 1969). Lippmann (1973) showed that evaporation would not favour dolomitization despite consequent high Mg:Ca ratios, because of the accompanying decrease in the activity of the CO3 2− free ion. The already low activity of the CO3 2− ion would be further suppressed due to its complexing to form neutral ion pairs, giving MgCO3 0 , while Mg2+ would also complex with sulphate, as MgSO4 0 . Experimental evaporation of brines in beakers has not produced dolomite in the sequence of precipitated minerals (e.g. Borchert and Muir, 1964, personal observation) clearly indicating that simple evaporation does not produce dolomite. Proposals that dolomite is an evaporative mineral (e.g. McKenzie et al., 1980), are thus shown to be untenable. Where dolomite occurs beneath the Abu Dhabi sabkha, it does not cross-cut facies, as might be expected from seepage reflux or evaporative pumping, but is located within horizontal beds associated with cyanobacterial mats (McKenzie et al., 1980; Wright, 2000) – suggesting that microbial mediation may be involved. More recent work by Bontognali et al. (2010) attributes the Abu Dhabi dolomite formation to a microbially mediated origin. 4.3.5. Documentation of anaerobic microbial mediation In contrast, anaerobic microbial mediation has been observed to cause the precipitation of dolomite in modern sediments and groundwater-basalt systems, with several different metabolic pathways and waters of different composition (Moore et al., 2004; Roberts et al., 2004; Douglas, 2005; Wright and Oren, 2005 and references therein; Wright and Wacey, 2005). These cases of dolomite formation have been replicated under controlled conditions in the laboratory (e.g. Vasconcelos et al., 1995; Roberts et al., 2004; Wright and Wacey, 2004, 2005), making a compelling case that microbial mediation is essential for low-temperature precipitation of dolomite in saline solutions (Douglas, 2005; Wright and Oren, 2005; Wacey et al., 2007). The idea that dolomite needs thousands or even millions of years to form is manifestly false, since recent experiments have conclusively demonstrated that dolomite can be formed in days through microbial mediation (Wright and Wacey, 2004, 2005). On the antiquity of microbial sulphate reduction, Shen et al. (2001) and Shen and Buick (2004) have used sulphur isotope data to argue that sulphate reducing microbes had evolved by the early Archaean. The large spread of 34 S values of microscopic pyrites aligned along growth faces of former gypsum crystals in the ∼3.47-Ga North Pole barite deposit of northwestern Australia provide the oldest evidence of microbial sulphate reduction and the earliest indication of a specific microbial metabolism. Microbial ecosystems dominated the depositional environments of Archaean, Proterozoic and some 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 et al., 2006; Gandin and Wright, 2007). Such ecosystems thus acted not only as the ‘engine house’ of early carbonate production, but also provide a ‘process analogue’ for dolomite formation. In conclusion, although seawater is supersaturated with respect to dolomite, the mineral is prevented from precipitating under normal earth-surface conditions by kinetic inhibitors. Numer-

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ous attempts by researchers undertaking many physico-chemical experiments over 200 years have failed, and ‘conventional’ models for dolomite formation lack experimental proof and thorough empirical support. The only proven process for overcoming the kinetic inhibitors to dolomite precipitation at sedimentary temperatures and pressures in the natural environment is through microbial mediation, which has operated since the Archaean. This provides compelling evidence that sedimentary dolomite carries the essential fingerprint of life. 4.4. Dolomite in ≥3700 Ma Isua supracrustal rocks In more ancient rocks such as the Isua rocks discussed here, high-grade metamorphism with deformation has minimised the hope of preserving microfossils, and can also mean that stable isotopic signatures for early life might be corrupted or ambiguous. Therefore, on the proviso that research on modern sediments and volcanic rocks and laboratory experiments now shows that anaerobic microbial mediation is essential for low-temperature dolomite formation (see above and Burns et al., 1988; Douglas, 2005), then the evidence from Isua of Eoarchaean pre-metamorphic low-temperature dolomite precipitation from sea or groundwater in sedimentary and volcanic rocks can be used as a fingerprint for biological activity, despite the superimposed amphibolite facies metamorphism and deformation. 4.5. Dolomite signature as evidence for life by 3850 Ma, from Akilia association rocks? Elsewhere in the Eoarchaean Itsaq Gneiss Complex of southern West Greenland, more deformed and highly metamorphosed (up to granulite facies) orthogneisses contain inclusions of chemical metasedimentary rocks, which are grouped within the Akilia association of McGregor and Mason (1977). The age of different units of Akilia association rocks probably varies from ≥3850 Ma to ca. 3600 Ma (Nutman et al., 1996, 2000; Friend and Nutman, 2005). However, the oldest interpreted ages of ≥3850 Ma are disputed (see Nutman et al., 2000, 2004 and Manning et al., 2006 versus Whitehouse et al., 1999, 2009). Some Akilia association diopside–diopside-rich rocks can be silica-penetrated igneous cumulate rocks (McGregor and Mason, 1977; Nutman, 1980). However, other rocks are so low in Al2 O3 , TiO2 , Cr and Ni (Table 1) that it is impossible they were derived by metasomatic silica addition to mafic rocks (Nutman, 1980; Manning et al., 2006; Friend et al., 2007). Instead they were derived from BIF and carbonate + chert sedimentary/diagenetic mixtures, with the latter now occurring as essentially de-carbonated upper amphibolite to granulite facies quartz + diopside ± amphibole ± magnetite rocks (Nutman, 1980; Manning et al., 2006; Nutman and Friend, 2006). These chemical sedimentary rocks, exemplified by G91/26R and G91/25 from Akilia (Fig. 1 inset), show a similar Ca–Fe–Mg spectrum as the lower metamorphic grade Isua BIF with a dolomitic component, such as G07/22 and G93/28 (Fig. 6). Seawater-like REE + Y(SN) patterns (slightly modified during granulite facies metamorphism) are preserved in Akilia association iron rich BIF (Bolhar et al., 2004) and across the Ca–Fe–Mg array (e.g. G91/25 and G91/26R in Fig. 8; Friend et al., 2007). If the ≥3850 Ma age for the particular samples G91/26R and G91/25 is correct, then the whole rock geochemical evidence that their protoliths contained early low-temperature dolomite would extend this signature for life back to ≥3850 Ma. 4.6. Looking for early life – a broader discussion Previous searches for evidence of Eoarchaean life in Isua rocks largely focused on interpreting 13 C-depleted isotopic signatures and petrography of graphite (Schidlowski et al., 1979; Mozjsis et

Fig. 8. REE + Y(SN) plot for Akilia association rocks G91/26R and G91/25, with Isua rocks of similar bulk composition (interpreted as BIF with dolomitic component) shown for comparison.

al., 1996; Rosing, 1999). However, due to the superimposed metamorphism, the biogenicity of the graphite’s carbon and the true meaning of its isotopic signature have been debated (van Zuilen et al., 2002). Thus although some of this graphite evidence is very compelling, a consensus has yet to be reached that any of it is undisputable evidence for life. From another perspective, the very presence of the Isua BIFs with the possibility of microbial mediation in deposition of such rocks (Cloud, 1973; Garrels et al., 1973; Konhauser et al., 2002) suggests credible evidence for early life. Perhaps consistent with this view is the strong fractionation of iron isotopes in Isua BIF relative to terrestrial igneous rocks (Dauphas et al., 2004), which is a signature of oxidation of Fe(II) by anoxygenic photoautotrophic bacteria or abiotic photo-oxidation (Konhauser et al., 2002). This biotic origin of BIF is also supported by the documentation of ferrous iron oxidation by some modern anoxygenic phototrophic bacteria (Liss et al., 1993). However, biotic BIF deposition would have been within an anoxic marine ecosystem (Beukes, 2004), which disappeared in the early Paleoproterozoic (Cloud, 1973) and cannot be directly studied today. Unlike BIF precipitation, low-temperature dolomite formation still occurs today in several environments (Von Der Borch, 1976; Gieskes et al., 1982; Vasconcelos and McKenzie, 1997; Moore et al., 2004; Wacey et al., 2007). In the laboratory it has only been replicated at normal sedimentary temperatures and pressures by anaerobic microbial mediation rather than by abiotic physicochemical processes (Vasconcelos et al., 1995; van Lith et al., 2003; Roberts et al., 2004; Douglas, 2005; Wright and Wacey, 2004, 2005). Bolhar et al. (2004), using REE + Y trace element chemistry, first demonstrated that an Isua carbonate (sample SM-GR-97/26) had a sedimentary protolith precipitated from a low-temperature seawater-like medium. By the same method, more Isua carbonate + quartz bearing rocks have subsequently been demonstrated to be sedimentary in origin (Friend et al., 2007), countering the general conclusion of Rose et al. (1996) and Rosing et al. (1996) that Isua carbonates are entirely metasomatic and never have sedimentary protoliths. However, Bohlar et al. ended their discussion by stating “whether the similarity in REE + Y pattern of modern seawater and the source of IGB (Isua) BIFs itself is an indication for the presence of life on Earth remains to be established” (Bolhar et al., 2004, p. 58). By focusing on the major element chemistry of Isua rocks with REE + Y(SN) seawater-like signatures that contained dolomite in their protoliths (sedimentary rocks and pillow lava interstices), we have now established a link between the chemistry of these rocks and life. If this

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interpretation of the process of low-temperature dolomite growth in some Isua rocks is correct, then by 3700 Ma, life was already well established in a range of ecological niches, including in a sub-surface volcanic pillow interstice environment. Such an environment would afford life protection from any surface-sterilizing massive impacts at that early stage of Earth’s history (summarized by Koeberl, 2006), and from whence surficial environments could be readily colonized again. 5. Conclusions (1) Dolomite was a pre-metamorphic mineral in some Isua chemical sedimentary rocks and pillow interstices. (2) The Isua dolomite (e.g. sample JHA170728) has a REE + Y(SN) seawater-like signature identical to international BIF standard IF-G from Isua, showing it formed in a pre-metamorphic lowtemperature environment, during deposition or diagenesis. (3) Recent low-temperature dolomite precipitation in sedimentary systems and intra-basalt ground waters apparently requires microbial mediation, and this has been replicated under controlled laboratory conditions. On the other hand, years of experimentation has failed to precipitate low-temperature dolomite by abiotic means. Given our results and these observations, we propose that geochemical evidence for premetamorphic Isua low-temperature dolomite is simple, direct evidence for microbial life by 3700 Ma. (4) If some Akilia association BIF with a dolomitic component and seawater-like REE + Y(SN) signatures really are ≥3850 million years old, then the dolomite component in these rocks is evidence of life at the start of Earth’s known sedimentary record. Acknowledgements This research was funded by Australian Research Council grant DP0342794, 2006–2009 support from the Chinese Academy of Geological Sciences and currently an operating grant from the University of Wollongong. The Geological Survey of Denmark and Greenland is thanked for permission to publish this paper. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.precamres.2010.08.006. References Adams, J.E., Rhodes, M.L., 1960. Dolomitization by seepage refluxion. AAPG Bulletin 44, 1912–1921. Allaart, J.H., 1976. The pre-3760 m.y. old supracrustal rocks of the Isua area, central West Greenland, and the associated occurrence of quartz-banded ironstone. In: Windley, B.F. (Ed.), The Early History of the Earth. Wiley, London, pp. 177–189. Altermann, W., Kazmierczak, J., Oren, A., Wright, D.T., 2006. Cyanobacterial calcification and its rock-building potential during 3.5 billion years of Earth history. Geobiology 4, 147–166. Appel, P.W.U., Moorbath, S., Myers, J.S., 2003. Isuasphaera isua (Pflug) revisited. Precambrian Research 126, 309–312. Badiozamani, K., 1973. The Dorag dolomitization model – application to the Middle Ordovician of Wisconsin. Journal of Sedimentary Petrology 43, 965–984. Baker, P., Kastner, M., 1981. Constraints on the formation of sedimentary dolomite. Science 213, 214–216. Balarew, C., Tepavitcharova, S., Rabadjieva, D., Voigt, D., 2001. Solubility and crystallization in the system MgCl2 –MgSO4 –H2 O at 50 and 75 ◦ C. Journal of Solution Chemistry 30, 815–823. Bau, M., Dulski, P., 1999. Comparing yttrium and rare earths in hydrothermal fluids from the Mid-Atlantic Ridge: implications for Y and REE behaviour during nearvent mixing and for the Y/Ho ratio of Proterozoic seawater. Chemical Geology 155, 77–90. Beukes, N.J., 1987. Facies relations, depositional environments and diagenesis in a major early Proterozoic stromatolitic carbonate platform to basinal sequence, Campbellrand Subgroup, Transvaal Supergroup, Southern Africa. Sedimentary Geology 54, 1–46. Beukes, N.J., 2004. Early options in photosynthesis. Nature 431, 522–523.

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