Preservation of biosignatures in metaglassy volcanic rocks from the Jormua ophiolite complex, Finland

Precambrian Research 136 (2005) 125–137

Preservation of biosignatures in metaglassy volcanic rocks from the Jormua ophiolite complex, Finland Harald Furnesa,∗ , Neil R. Banerjeea,b , Karlis Muehlenbachsb , Asko Kontinenc b

a Department of Earth Science, University of Bergen, Allegt. 41, 5007 Bergen, Norway Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alta., Canada T6G 2E3 c Geological Survey of Finland, P.O. Box 1237, FIN-70211 Kuopio, Finland

Received 21 April 2004; accepted 20 September 2004

Abstract Evidence of microbially-mediated alteration of basaltic glass is preserved in originally glassy basalts (rims of pillow lavas, hyaloclastite breccias, and chilled margins of dykes) from the well-preserved 1.95 Ga Jormua ophiolite complex (JOC) of Northeastern Finland. Although textural evidence of microbial alteration is commonly observed in relic glass from recent oceanic crust and some ophiolites, these textures have been destroyed during greenschist to lower amphibolite facies regional metamorphism and deformation of the JOC. However, another robust biosignature is found in the generally depleted ␦13 C values of disseminated carbonate extracted from originally glassy basalts, relative to crystalline samples. The same distribution of ␦13 C values is well documented in samples from recent oceanic crust as well as ophiolites of Phanerozoic age. This characteristic contrast in the ␦13 C values of disseminated carbonate is interpreted to result from microbe-induced fractionation during oxidation of organic matter. X-ray mapping of initial alteration zones has identified residual carbon associated with highly-concentrated S that is unrelated to carbonate. We attribute these biosignatures to microbially-mediated alteration of originally glassy material prior to ophiolite emplacement. © 2004 Elsevier B.V. All rights reserved. Keywords: Finland; Early Proterozoic; Ophiolite; Volcanic glass; Bioalteration; ␦13 C

1. Introduction It is now well-documented that basaltic glass from in-situ upper oceanic crust, as well as from pillow lavas of well-preserved ophiolites, hosts extensive mi∗

Corresponding author. Fax: +47 55 315729. E-mail address: [email protected] (H. Furnes).

0301-9268/$ – see front matter © 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2004.09.009

crobial activity. The strongest evidence of this process is provided by microbially-induced alteration textures that are generated during etching of glass along fractures that allow seawater penetration, as colonizing microbes utilize the glass as a source of energy/nutrients. Ross and Fisher (1986) were the first to describe etching of volcanic glass by microorganisms. Thorseth et al. (1992) demonstrated that this process had been

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operative in basaltic glass fragments in a subglacial hyaloclastite in Iceland and suggested a mechanism whereby microbes controlled local changes in pH. Later, microbial etching was experimentally demonstrated (Thorseth et al., 1995a; Staudigel et al., 1995), and Welch and Ullmann (1996) showed that organic acids are more effective in dissolving silicates than inorganic acids. In the presence of bacteria Brantley et al. (2001) demonstrated that during dissolution of hornblende metals were preferentially released, components that bacteria commonly use in their enzymes. Microbes on basalt glass from the Knipovich Ridge, suspected to be the iron-oxidizing bacterium gallionella have been described by Thorseth et al. (2001), and Emerson and Moyers (2002) observed obligate microaeriophilic Fe oxidizers in hydrothermal vents from Loihi seamount. Oxidation of reduced iron may hence be one of the major forms of energy these microbes may obtain from the glass. The biogenicity of these textures is supported by (1) elevated levels of biologically important elements such as C, N, P, K, and S. X-ray mapping of areas of biogenerated textures invariably display significant increase in carbon content at their alteration fronts. In the absence of associated calcium the carbon is hence considered to be of organic origin (e.g. Furnes et al., 2001b). (2) Presence of fossilized organic remains such as filaments and/or biofilms. In a study of bio-generated textures in the pillow lavas from the Troodos ophiolite (Cyprus) fossilized filaments attached to the fresh glass were observed (Furnes et al., 2001c). (3) depletion of carbon isotope ratios of disseminated carbonate by as much as –20‰, suggesting microbial fractionation during metabolic oxidation of organic matter. All studies of bioaltered glassy margins of pillows show that these parts in general have significantly lower ␦13 C than the crystalline interior of pillows (e.g. Furnes et al., 2001a). (4) identification of microbes inhabiting altered glassy margins of recent pillow lavas. On the glassy margin surface of pillow lavas from the Knipovich Ridge incipient microbial colonization, represented by various microbial morphologies has been described by Thorseth et al. (2001). All the abovementioned features have been comprehensively dealt with in a number of recent papers (Thorseth et al., 1995; Furnes et al., 1996; Fisk et al., 1998; Torsvik et al., 1998; Furnes and Staudigel, 1999; Furnes et al., 1999, 2001a, 2001b; Thorseth et al., 2001;

Banerjee and Muehlenbachs, 2003; Staudigel et al., 2003; Thorseth et al., 2003). Research into microbially-mediated alteration of insitu oceanic crust is a rapidly growing field. However, the investigation is beset by many problems such as sample recovery, depth of penetration during drilling, lateral correlation, and a very limited age range in the Earth’s history (Furnes et al., 2001b). Many of these problems are mitigated by investigation of wellpreserved ophiolitic volcanic sequences. Hence, investigation of bio-tracers in the volcanic succession of ophiolites is complementary to the study of microbial activity within recent oceanic crust, and may indeed provide answers that can only be documented in the volcanic rocks of ophiolites. Despite the great potential ophiolites represent in the search for fossil evidence of microbial alteration, only very limited information so far exist in this respect, i.e. one account from the Cretaceous Troodos ophiolite complex (Furnes et al., 2001c), and another account on the Late Ordovician Solund–Stavfjord ophiolite complex (Furnes et al., 2002). In this paper we present new evidence for bioalteration of originally glassy material from the Early Proterozoic Jormua ophiolite complex (JOC), Northeastern Finland and discuss the data in light of the present body of knowledge concerning microbiallymediated alteration of in-situ oceanic crust and ophiolites.

2. Regional setting of the Jormua ophiolite complex The JOC occupies the central part of an Early Proterozoic (2.3–1.92 Ga) metasedimentary sequence surrounded by Archean basement rocks (Fig. 1A; Kontinen, 1987; Peltonen et al., 1998). The JOC contains all the components of a typical Penrose-style ophiolite including pillow lavas and volcanic breccias, a sheeted dyke complex, mafic cumulates, and mantle peridotites (Fig. 1B). Although it has been tectonically disrupted into several blocks, reconstruction of the sequence has been proposed as shown in Fig. 1C (Kontinen, 1987; Peltonen et al., 1996). The thickness of the JOC varies and in places the lava sequence rests directly upon the mantle rocks (Fig. 1C). The crustal and mantle character of the various blocks indicate formation at slow-spreading conditions during continental

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Fig. 1. (A) Simplified geological map of the Jormua ophiolite complex (JOC) and adjacent lithologies of Northeastern Finland. (B) Detailed map of part of the JOC and location of sampling sites. (C) Reconstructed profile of the JOC and sampling sites. All illustrations are modified versions based on the work of Peltonen et al. (1989).

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break-up and initiation of a Red Sea type oceanic crust (Peltonen et al., 2003). The age of the ophiolite has been determined to be 1953 ± 3 Ma, based on a concordant U–Pb zircon age from a gabbroic dyke cutting the mantle tectonite (Peltonen et al., 1996). The JOC thus represents one of the oldest examples of crust formed by ocean floor spreading in which all components are present.

3. Sampling locations The thickness of the metabasaltic volcanic sequence varies between locations up to a maximum of 500 m. Sampling was carried out in four locations, representing different stratigraphic levels within the ophiolite (Fig. 1B and C). Location 1 consists of pillow lavas that may be close-packed (Fig. 2A) or, in most cases, contain thin inter-pillow hyaloclastite breccias (Fig. 2B). The pillows are invariably non-vesicular, demonstrating eruption in relatively deep water (probably >1000 m). Remnants of variolitic textures can still be seen in the outer ∼1 cm of the chilled pillow margins. The stratigraphic position of location 1 is estimated to be ∼125 m below the top of the volcanic sequence. Location 2 consists mainly of hyaloclastite breccias (Fig. 2C) interbedded with layers of pillow lavas. Some of the pillows exhibit well-defined drainage structures, thus defining reliable way-up indicators for the sequence. The fine-grained to glassy fragments of the volcanic breccias, mostly highly irregular in shape, range in size from less than 1 mm up to ∼10 cm long (Fig. 2C). The stratigraphic position of location 2 is estimated to be ∼350 m below the top of the volcanic sequence. Location 3 consists of a mixture of pillow breccias/hyaloclastites and pillow lavas, intruded by dykes (subordinate relative to the volcanic rocks). This location is thought to represent the transition zone between the sheeted dyke complex and the volcanic sequence. The stratigraphic position of location 3 is estimated to be ∼100 m below the top of the volcanic sequence, showing that the dykes, in places, reach a high stratigraphic level within the volcanic succession. Location 4 represents a 100% sheeted dyke complex (Fig. 2D). The dykes are subparallel with well-defined chilled margins along younger dykes that split the older

Fig. 2.

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dykes (Fig. 2D). The pseudostratigraphic position of the dyke sequence is difficult to determine, but as a conservative estimate, we propose a depth of between 500 and 1000 m below the top of the volcanic sequence.

4. Petrography All of the volcanic rocks and the dykes from the JOC have been metamorphosed and foliated during upper greenschist to lower amphibolite facies conditions. The peak of metamorphism has been estimated at ∼500 ± 20 ◦ C (Peltonen et al., 1996). This metamorphism and deformation has completely recrystallized the rocks and erased any trace of primary magmatic textures. Despite the pervasiveness of this metamorphic overprint, the originally glassy rims of pillows, hyaloclastite breccias, and dyke margins are easily recognized in outcrop (Fig. 2) and hand specimen. The typical mineral assemblage in the metabasaltic rocks is oligoclase–andesine + actinolitic hornblende– hornblende + Fe-oxides ± chlorite ±epidote ± quartz ± biotite. Relic glass is commonly replaced by similar mineral assemblages. The common equilibrium metamorphic mineral assemblage in the mantle serpentinites is antigorite + Cr-magnetite ± tremolite.

5. Sample preparation and analytical methods Samples were collected as mini cores (2.5 cm wide by 5–10 cm long) with the use of a rock drill (Fig. 2A). Individual samples commonly cut different lithologies and were divided into sub-samples that have the same sample number but different names (e.g. 7-J01 pillow rim and 7-J-01 interpillow hyaloclastite, see Table 1). For purposes of this study, we separated samples that originally contained glass from those that were originally crystalline. Glassy samples include sam-

Fig. 2. (A) Location of drill cores in a pillow interior (PI), pillow rim (PR), and interpillow hyaloclastite (IH) from location 1. Subsamples were made from cores that cut more than one lithology. (B) Pillow lavas surrounded by hyaloclastite breccia from location 1. Note the white chilled margin and variolitic zones (up to 2 cm thick) around the pillows. (C) Hyaloclastite breccias from location 2. (D) Sheeted dyke complex at location 4. The formerly glassy chilled margins (white) are clearly visible. Lens cap is 5 cm.

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ples of chilled pillow rims (including the variolitic zones), fragments and matrix of hyaloclastite breccias, as well as samples from the chilled margins of dykes within the sheeted dyke complex. The crystalline samples represent the interiors of pillows and dyke rocks. A total of 51 glassy and 11 crystalline samples were analysed for this study. Sampling locations and their relative stratigraphic positions are shown in Fig. 1. Carbonates were analysed for their C-isotopic composition by extracting CO2 under vacuum from wholerock powders after treatment with 100% phosphoric acid for 1 h at 25 ◦ C (McCrea, 1950). The exsolved CO2 was analysed on a Finnigan MAT 252 mass spectrometer. The errors in calculated carbonate yields range from ∼±1 to ∼±15% for samples rich and poor in carbonate, respectively. The errors on isotopic analyses for carbon are better than ±0.1‰. The isotope data are reported in the usual delta-notation with respect to VPDB for carbon and VSMOW for oxygen (Craig, 1957, 1961). X-ray mapping of iridium-coated thin sections was carried out using a JEOL JXA-8900R microprobe with an accelerating voltage of 15 kV and probe current of 3 × 10−8 A at the University of Alberta (Banerjee and Muehlenbachs, 2003). Scanning electron microscopy (SEM) observations were performed on a JEOL JSM6301FXV instrument connected to a Princeton Gamma Tech IMIX energy-dispersive spectrometer system at the University of Alberta. The analyses were performed at an accelerating voltage of 20 kV and a working distance of 15 mm.

6. Results Petrographic analysis of 30 thin sections of formerly glassy volcanic rocks from the JOC failed to reveal any textural evidence of microbial alteration. Despite the lack of textural evidence we conducted carbon isotope analyses and X-ray element mapping to search for biosignatures. These techniques have previously been successfully applied in studies of recent oceanic crust and ophiolites, including ∼3.5 Ga pillow basalts that have been metamorphosed under greenschist facies conditions (Furnes et al., 2004). We thus decided to build on previous studies and determine if evidence for microbial activity was preserved despite the per-

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Table 1 Carbon isotope analyses of carbonates from glassy and crystalline samples Sample #

Type

␦13 C

Loc. 1. Upper volcanic zone Glassy material 2-J-01 PR −3.4 4-J-01 PR −9.1 5B-J-01 PR −8.2 5C-J-01 PR −13.4 6-J-01 PR −4.7 7-J-01 PR −10.6 9-J-01 PR −13.3 10-J-01 PR −8.5 12-J-01 PR −5.4 14-J-01 PR −12.4 3-J-01 IH −7.5 7-J-01 IH −7.3 13-J-01 IH −6.2 4-J-01 VZ −10.8 6-J-01 VZ −3.3 10-J-01 VZ −6.6 12-J-01 VZ −3.9

5A-J-01 8-J-01 8A-J-01 8B-J-01 8C-J-01 14A-J-01 14B-J-01 14C-J-01

Crystalline material PI −3.4 PI −6.9 PI −6.7 PI −5.5 PI −6.9 PI −8.6 PI −5.9 PI −8.5

␦18 O

wt.% cc

10.9 8.8 22.5 18.4 24.2 17.2 14.6 24.1 16.0 18.5 21.5 19.0 23.0 6.2 16.9 16.5 21.5

0.010 0.005 0.008 0.026 0.005 0.005 0.002 0.021 0.006 0.001 0.006 0.006 0.005 0.001 0.002 0.003 0.007

19.9 21.8 17.0 19.2 18.6 16.0 18.0 20.8

0.011 0.007 0.025 0.023 0.016 0.022 0.022 0.008

Loc. 3. Volcanic-dyke transition zone 38-J-O1 37-J-01 41-J-01 44-J-01 45-J-01 35-J-01 36-J-01 44-J-01

Glassy material PR −11.2 H −10.6 H −9.5 H −8.9 H −10.8 CDM −8.5 CDM −9.7 CDM −9.2

␦13 C

␦18 O

wt.% cc

Loc. 2. Lower volcanic zone Glassy material 26-J-01 PR −8.7 28A-J-01 PR −6.5 28B-J-01 PR −9.2 30A-J-01 PR −14.1 30B-J-01 PR −10.6 30C-J-01 PR −12.6 26-J-01 IH −4.2 28-J-01 IH −6.2 15-J-01 H −13.1 16-J-01 H −8.8 17-J-01 H −8.9 18-J-01 H −7.4 19-J-01 H −6.3 20-J-01 H −5.5 22-J-01 H −10.5 23-J-01 H −6.5 15-J-01 HF −5.9 16-J-01 HF −10.1 17-J-01 HF −9.3 21-J-01 HF −10.6 22-J-01 HF −13.9

18.4 15.6 15.1 18.1 19.9 20.5 19.0 22.5 20.7 20.2 23.6 22.3 22.0 26.4 15.3 23.9 22.1 21.5 18.6 21.8 21.6

0.017 0.022 0.017 0.020 0.011 0.014 0.011 0.009 0.003 0.006 0.009 0.005 0.010 0.007 0.005 0.014 0.006 0.012 0.007 0.030 0.014

19.8

0.020

Sample #

30-J-01

Type

Crystalline material PI −6.8

Loc. 4. Sheeted dyke complex 9.2 13.0 14.5 10.0 13.5 9.6 16.7 8.2

0.007 0.002 0.012 1.188 13.680 0.004 0.009 8.400

47-J-01 49-J-01 50-J-01

Glassy material CDM −10.2 CDM −10.7 CDM −9.8

19.0 12.7 17.6

0.027 0.006 0.030

46-J-01 50A-J-01

Crystalline material DI −5.8 DI −8.8

18.9 17.9

0.010 0.048

wt.% cc = weight percentage carbonate; PR = pillow rim; PI = pillow interior; IH = interpillow hyaloclastite; VZ = variolitic zone; H = hyaloclastite; HF = fragment in hyaloclastite: CDM = chilled dyke margin; DI = dyke interior.

vasive upper greenschist to lower amphibolite facies metamorphic overprint of the JOC. 6.1. Carbon isotopes The ␦13 C values and concentrations of disseminated carbonate from the various glassy components (pillow rims, variolitic pillow rims, inter-pillow hyalo-

clastites, hyaloclastite breccias, fragments in the hyaloclastite breccias, and chilled margins of dykes) and crystalline material are listed in Table 1. In general the originally glassy components have depleted ␦13 C values relative to the crystalline material. The ␦13 C values of glassy material vary from −3.3 to −14.1‰ (average = −8.8‰), whereas the crystalline samples vary from −3.4 to −8.8‰ (average = −6.7‰). Fig. 3A

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tion) contain the highest carbonate contents (>1 wt.%) but still exhibit depleted ␦13 C values (from −8.9 to −10.8‰). The lowest ␦13 C value (−14.1‰) comes from a pillow rim at location 2 within the lower volcanic zone. 6.2. Element distribution Although clear textural evidence of microbial alteration is lacking in the JOC, we selected samples displaying fine-grained alteration textures after glass for X-ray element mapping to prospect for elevated concentrations of elements commonly associated with organic material. The distribution of C, S, K, Ca, and Fe from two formerly glassy pillow margins are shown in Fig. 4. Areas of elevated carbon concentration occur within amphibole and plagioclase crystals, at the boundary between amphibole and plagioclase crystals, and associated with chlorite (Fig. 4). Carbon was routinely measured on two different spectrometers to monitor the reproducibility of observed signals. This ensured topographic effects were not responsible for the elevated values. Only one carbon map is shown for each sample in Fig. 4 but both carbon maps for the same areas show the same distribution. Maps of Al, Cl, N, Na, P, Mg, Si, and Ti were also routinely measured. Most of these elements do not show enrichments beyond what might be expected mineralogically and argue against the possibility of carbon highs due to inorganic carbonate material (e.g., Ca, Mg, Fe) or epoxy (e.g. Cl). The elevated carbon values are thus taken to represent a pure carbon component, interpreted as residual organic material. Associated with the areas of high carbon concentrations are also high concentrations of S (Fig. 4).

7. Discussion Fig. 3. Relationship between ␦13 C and wt.% carbonate for the glassy and crystalline samples of (A) the Jormua ophiolite complex, (B) Phanerozoic ophiolites, and (C) recent oceanic crust.

shows the relationship between ␦13 C ratio and weight percentage carbonate, calculated from the proportion of exsolved CO2 , for the glassy and crystalline samples. Glassy samples poor in carbonate (< 0.01 wt.%) typically show ␦13 C values less than −10‰. Three glassy samples from location 3 (volcanic–dyke transi-

Direct evidence of biomediated alteration is commonly provided by textural relationships. These features, represented by coalesced spherical bodies generally less than 1 ␮m in diameter (referred to as granular texture), and straight to curved tubular bodies attaining lengths up to > 100 ␮m (referred to as tubular texture), are commonly considered as the most powerful evidence of bioalteration (Furnes and Staudigel, 1999; Furnes et al., 2001b; Banerjee and Muehlenbachs, 2003). Such structures are, however, only clearly pre-

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Fig. 4. Series of backscattered electron images (BSE) and X-ray maps (C, S, K, Ca, and Fe) from two areas within the formerly glassy chilled pillow rim of sample 7-J-01 from location 1. Increasing order of elemental abundance black–blue–green–yellow–red–pink. The letters (A to D) in the top carbon map refer to Fig. 5. Am = amphibole; Ap = apatite; Cl = chlorite; FeOx = iron oxide; Pl = plagioclase; Ti = titanite. Scale bar is 50 ␮m.

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served when fresh glass is still present, or if the recrystallized material has no foliation (Furnes et al., 2002). In the metamorphically recrystallized and foliated material of the Jormua ophiolite complex there is no trace left of such structures, and hence, all evidence of bioalteration must be based on other lines of evidence. We suggest that carbon isotope data and the distribution of carbon and other biologically essential elements represent powerful biosignatures that have withstood the metamorphism and deformation of these rocks. 7.1. Interpretation of carbon isotope data Disseminated carbonate in the various originally glassy rocks from the JOC is isotopically lighter (average −8.8‰) than in crystalline rocks (average −6.7‰). Also, the relic glassy samples extend to much lower ␦13 C values (−14.1‰) than the crystalline rocks (−8.8‰). Crystalline interior samples are relatively poor in carbonate (< 0.048 wt.%) and typically display ␦13 C values comparable with primary mantle CO2 (−5 to −7‰; Alt et al., 1996; Hoefs, 1997). Such isotopic contrasts are also seen in pillow lava rims from Phanerozoic ophiolites (Fig. 3B) and recent oceanic crust (Fig. 3C) where the generally low ␦13 C values of disseminated carbonate are attributed to metabolic byproducts formed during microbial oxidation of dissolved organic matter in pore waters (Furnes et al., 2001a; Banerjee and Muehlenbachs, 2003; Furnes and Muehlenbachs, 2003). Assuming that all the basaltic samples, prior to alteration and metamorphism, had initial magmatic ␦13 C values in the range −5 to −7‰ (Alt et al., 1996; Hoefs, 1997) and that inorganically precipitated marine carbonate had ␦13 C values close to zero (Alt et al., 1996) we can make some predictions regarding the observed distribution of ␦13 C values in the glassy and crystalline samples. Fourteen of the 49 (∼29%) formerly glassy samples and 8 of the 11 (∼73%) crystalline samples have ␦13 C values between –7 and 0‰. These values are best explained as being derived from some combination of inorganically precipitated carbonate and magmatic values. We then have to explain the large variation in the ␦13 C values of the glassy samples, i.e. those <−7‰. The lightest ␦13 C values occur in relic glassy samples with very low carbonate abundances (0.001–0.025 wt.%). Hence we consider it unlikely that Rayleigh fractionation, a process that may fractionate

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carbon isotope ratios during basalt degassing and subsequent alteration (Hoefs, 1997), is responsible for the pronounced spread in the ␦13 C values of the glassy components. The samples with ␦13 C values lower than −7‰ most likely contain an amagmatic carbon component with low ␦13 C. It has recently become well-established that microbes live on the surface of, and along fractures within, basaltic glass. We predict that this microbial activity should be reflected in the carbon isotope ratio of carbonates produced from metabolic byproducts. The carbon isotopes (12 C and 13 C) may be fractionated by biotic processes (Oremland, 1988). Most microorganisms obtain carbon and energy from the oxidation of organic matter. During respiration, organic material is oxidized to 12 C-enriched CO2 that may subsequently be precipitated as carbonate minerals with depleted ␦13 C values. We interpret the isotopically light ␦13 C values of carbonate in the formerly glassy JOC samples, particularly those samples that contain low abundances of carbonate and ␦13 C values < −10‰, to be the result of microbial fractionation. If the proportion of inorganically precipitated calcite (with ␦13 C ∼ 0‰) exceeds that resulting from precipitation of carbonate from microbially-produced CO2 during oxidation of organic carbon, the ␦13 C value of a sample may be greater than −7‰. This is commonly seen for samples that are relatively rich in carbonate (Fig. 3A; Table 1). It is interesting that the three glassy samples with the highest carbonate contents (from 1 to 14 wt.%), have ␦13 C values (−8.9 to −10.8‰) below those expected for mantle CO2 . Preliminary carbon isotope analyses of metasomatic carbonate rocks produced during regional metasomatism of the mantle sequence have ␦13 C values mostly between 0 and −3‰, with some values as low as −17‰ (J. Karhu, unpublished data). The ␦18 O values of these metasomatic carbonates (+10 to +13 ‰) are also similar to the ␦18 O values of the three samples with high carbonate contents (Table 1). Given the fact that all of these samples have high carbonate contents, their ␦18 O values are similar to those from the mantle metasomatic carbonates, and they all occur within the volcanic-dyke transition zone, we suggest they may have formed from fluids similar to or mixed with those responsible for the precipitation of the carbonates within the mantle sequence. However, it is important to point out that the distribution of carbon isotope values between glassy and

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crystalline samples cannot be explained by interaction with the fluids that produced the mantle carbonates since no systematic difference in the ␦13 C values between glassy and crystalline samples would be expected. 7.2. Distribution of bio-elements The presence of carbon, unrelated to carbonates, is interpreted to represent residual organic remains in the rock. The organic remains may have been preserved in various ways, either (1) by having been enclosed by the secondary mineral phases as inclusions, or (2) having become part of the crystalline structure. In this context it is pertinent to understand what happens to a microbial cell when it becomes fossilised. Bacteria have the ability to bind metals, (Fe, Al)-silicates and pure silica to their highly reactive surfaces and thus act as favourable interfaces for mineral development (e.g. Ferris et al., 1987; Beveridge, 1989; Urrutia and Beveridge, 1994; Douglas and Beveridge, 1998; Ehrlich, 1999). Westall et al. (1995) demonstrated experimentally that the cell wall of a bacterium, as well as the cytoplasm and extracellular polysaccharide substances produced by Bacteria, act as nucleation sites for mineralization. During this process organic remains may be trapped as inclusions within or between crystals. This mechanism has been proposed for the preservation of organic matter in fracture fillings of authigenic minerals of the Columbia River basalts (McKinley et al., 2000) and in the 3.3–3.5 Ga cherts of the Barberton greenstone belt (Westall et al., 2001). A similar mechanism was proposed for the preservation of organic remains in the greenschist metamorphic pillow lava rims of the Late Ordovician Solund–Stavfjord ophiolite complex (Furnes et al., 2002). It is also known that organic molecules may become part of crystals during diagenesis and recrystallization (Collins et al., 1995; Van Lith et al., 2001). We have performed SEM imaging (Fig. 5) of the areas of high carbon content illustrated in Fig. 4. Each of these areas occurs either within a secondary mineral phase (inclusion) or along the grain boundaries between minerals. Within these areas we have identified mineralized features that demonstrate the elevated carbon and sulphur contents are not due to contamination from debris (Fig. 5). The mineralized features are of approximately the correct size for microbial or-

Fig. 5. SEM images of the areas of high carbon content identified in Fig. 4 (labelled A to D). Arrows indicate material interpreted as mineralized organic remains. Scale bar is 5 ␮m for A to C and 10 ␮m for D.

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ganic remains but are morphologically equivocal due to their simple rounded nature. At the relatively high temperatures (∼500 ◦ C) and pressures these rocks have suffered, we suggest these features are likely recrystallized and have lost their original form. The carbon and sulphur within these features has not been remobilized and incorporated into secondary metamorphic mineral phases but rather has remained within inclusions or along grain boundaries. 7.3. Timing of bioalteration The most active stage of bioalteration in the oceanic crust takes place within a few millions of years after emplacement at the ridge crest. During this stage high water circulation effectively cools the crust to deep levels (Alt et al., 1986), making living conditions for microorganisms possible well into the sheeted dike complex, as suggested for the Costa Rica Rift (Furnes et al., 1999). The available data suggest that the biosignatures associated with alteration were introduced prior to the regional amphibolite facies metamorphism between 1.87 and 1.85 Ga (Tuisku, 1997), and probably closer to 1.95 Ga (the magmatic age of the JOC). Evidence for this is provided by 1) the difference in the ␦13 C values of the glassy and crystalline components, and 2) the organic remains encrusted by amphibole and plagioclase. The ␦13 C data show that the original glassy components preserve isotopic signatures consistent with microbial activity while the crystalline material does not (Fig. 3). Experimental work has demonstrated that microbes prefer to colonize glass rather than crystalline material (Brekke, 1998). It is unclear, and beyond the scope of this paper, what causes the preferential colonization but it is most likely driven as a result of chemotactic and/or electrostatic factors, processes that likely commence immediately after formation of the pillow lavas. Hence, it is reasonable to consider the observed biosignatures as premetamorphic, since one should not expect any difference in the ␦13 C values between “glassy” and “crystalline” components if these were imposed penecontemporaneously with the regional metamorphism. In addition, textural relationships indicate the organic remains were present in the glassy rims of the pillow lavas prior to the formation of the metamorphic minerals (Fig. 4). We therefore suggest that these biosignatures were preserved in the glassy volcanic rocks while

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at the spreading ridge, at the stage when hydrothermal process provided vigorous water circulation through the crust, and hence the availability of nutrients was at a maximum. 7.4. Important perspectives Microbial activity in glassy pillow lavas and hyaloclastites from the upper oceanic crust is potentially very important because it might involve a very large biomass (Gold, 1992). This process likely extends from mid-ocean ridges and ridge flanks, into the ocean basins, and possibly into the ocean trenches where oceanic crust is subducted (Staudigel et al., 2004). Bioalteration of oceanic basaltic glass may thus be regarded as a fundamental geochemical process as it undoubtedly influences the chemical exchange between oceans and the oceanic crust as well as chemical transport in mid-ocean ridge hydrothermal systems at temperatures that allow life to exist (Staudigel et al., 1998). It has been suggested that the birthplace for life may have been connected to volcanic environments, such as deep-sea hydrothermal vents within early oceanic crust (e.g., Russel and Hall, 1997). An important task is to search for evidence of bioalteration in ancient pillow lavas and other glassy subaqueous volcanic rocks in which microbial life may have thrived. It is pertinent to mention that clear evidence of several types of biosignatures have been found in the ∼3.5 Ga pillow lavas of the Barberton Greenstone Belt (Furnes et al., 2004). This magmatic sequence has been interpreted to represent Mesoarchean oceanic crust, though a typical sheeted dyke complex is absent (De Wit et al., 1987). From the present study it appears that biosignatures such as organic carbon as well as sulphur get trapped during authigenic and metamorphic mineral growth that replace the original glassy rocks (selvages of pillows and hyaloclastite breccias). In addition the C-isotopic signatures in carbonate that reflect microbial metabolism during bioalteration of the glassy components appear to be also preserved. These positive indications of microbial interaction with the formerly glassy rocks of the 1.95 Ga Jormua ophiolite complex, representing one of the oldest known examples of a Penrose-style ophiolite, extends the evidence for bioalteration of indisputable oceanic crust back to Early Proterozoic times.

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8. Conclusions

References

Despite the lack of textural evidence, the preset study confirms that geochemical evidence for microbial alteration of formerly glassy rocks within the Jormua Ophiolite, Finland is preserved. The average ␦13 C value of disseminated carbonates within originally glassy components (−8.8‰) is lower than that of mantle carbonates (−5 to −7‰), whereas the average ␦13 C value within crystalline material (−6.7‰) is comparable. Only 29% of the ␦13 C values of the disseminated carbonates in glassy samples fall between mantle values and inorganically precipitated marine carbonate (∼0‰), which represent the two largest carbon reservoirs. Conversely most of the crystalline samples fall within this range. In addition, the disseminated carbonates in glassy samples extend to much lower ␦13 C values (−14.1‰) than the crystalline samples. We interpret these results to be the result of biological fractionation of carbon isotopes by Bacteria during colonization of the formerly glassy basalts. Furthermore, the presence of carbon (unrelated to carbonates) and sulphur, trapped as inclusions and between metamorphic minerals, associated with mineralized structures of possible organic origin, provides evidence of organic remains in the originally glassy volcanic rocks. We suggest that the bioalteration process took place early, when hydrothermal activity provided vigorous water circulation (and maximum availability of nutrients) through the oceanic crust, soon after formation at the ridge axis. This study extends the evidence for microbial alteration of oceanic crust back to the Early Proterozoic.

Alt, J.C., Honnorez, J., Laverne, C., Emmermann, R., 1986. Hydrothermal alteration of a 1 km section through the upper oceanic crust, DSDP Hole 504B: the mineralogy, chemistry, and evolution seawater–basalt interactions. J. Geophys. Res. 91, 10309–10335. Alt, J.C., Laverne, C., Vanko, D.A., Tartarotti, P., Teagle, D.A.H., Bach, W., Zuleger, E., Erzinger, J., Honnorez, J., Pezard, P.A., Becker, K., Salisbury, M.H., Wilkens, R.H., 1996. Hydrothermal alteration of a section of upper oceanic crust in the eastern equatorial Pacific: a synthesis of results from site 504 (DSDP Legs 69, 70, and 83, and ODP Legs 111, 137, 140 and 148). In: Alt, J.C., Kinoshita, H., Stokking, L.B., Michael, J.P. (Eds.), Proceedings of the Ocean Drilling Program, Sci. Results 148, College Station, TX (Ocean Drilling Program), pp. 417–434. Banerjee, N., Muehlenbachs, K., 2003. Tuff life: Bioalteration in volcaniclastic rocks from the Ontong Java Plateau. Geochem. Geophys. Geosyst. vol. 4 (no. 4), 10.1029/2002GC000470. Beveridge, T.J., 1989. Role of cellular design in bacterial metal accumulation and mineralization. Annu. Rev. Microbiol. 43, 147–171. Brantley, S.L., Lierman, L., Bau, M., Wu, S., 2001. Uptake of trace metals and rare elements from hornblende by a soil bacterium. Geomicrobiol. J. 18, 37–61. Brekke, I., 1998. Microbial degradation of minerals and glasses: the importance of iron content and crystallinity. Candidatus Scientiarum Thesis, University of Bergen, Norway. p. 70. Collins, M.J., Bishop, A.N., Farrimond, P., 1995. Sorption by mineral surfaces: Rebirth of the classical condensation pathway for kerogen formation? Geochim. Cosmochim. Acta 59, 2387–2391. Craig, H., 1957. Isotopic standards for carbon and oxygen and correlation factors for mass-spectrometric analysis of carbon dioxide. Geochim. Cosmochim. Acta 12, 133–149. Craig, H., 1961. Standards for reporting concentrations of deuterium and 18 O in natural waters. Science 133, 1702–1703. De Wit, M.J., Hart, R.A., Hart, R.J., 1987. The Jamestown ophiolite complex Barberton mountain land: a section through 3.5 Ga oceanic crust. J. African Earth Sci. 6, 681–730. Douglas, S., Beveridge, T.J., 1998. Mineral formation by bacteria in natural microbial communities. FEMS Microbiol. Ecol. 26, 79–88. Ehrlich, H.L., 1999. Microbes as geologic agents: their role in mineral formation. Geomicrobiol. J. 16, 135–153. Emerson, D., Moyers, C.L., 2002. Neutrophilic Bacteria are abundant at the Loihi seamount hydrothreaml vents and play a major role in Fe oxide deposition. Appl. Envir. Microbiol. 68, 3085–3093. Ferris, F.G., Fyfe, W.S., Beveridge, T.J., 1987. Bacteria as nucleation sites for authigenic minerals in a metal-contaminated sediment. Chem. Geol. 63, 225–232. Fisk, M.R., Giovannoni, S.J., Thorseth, I.H., 1998. Alteration of oceanic volcanic glass: textural evidence of microbial activity. Science 281, 978–979. Furnes, H., Thorseth, I.H., Tumyr, O., Torsvik, T., Fisk, M.R., 1996. Microbial activity in the alteration of glass from pillow lavas from Hole 896A. In: Alt, J.C., Kinoshita, H., Stokking, L.B., Michael,

Acknowledgements We thank J. Karhu for allowing access to his unpublished data; J. Ellingsen for help with the illustrations; O. Levner for help with the isotope analyses; S. Matveev for help with the X-ray mapping; and G. Braybrook for help with the SEM. This study was financed through the RCN (the Research Council of Norway) projects “SUBMAR” and “Tracing fossil bio-signals in young to ancient ophiolites”, an RCN postdoctoral fellowship to NB, grants from NSERC, and field support by the Geological Survey of Finland. We acknowledge constructive reviews by P. Schiffman and an anonymous reviewer.

H. Furnes et al. / Precambrian Research 136 (2005) 125–137 J.P. (Eds.), Proc. Ocean Drilling Program, Sci. Results 148. College Station, Texas (Ocean Drilling Program), pp. 191–206. Furnes, H., Muehlenbachs, K., Tumyr, O., Torsvik, T., Thorseth, I.H., 1999. Depth of active bio-alteration in the ocean crust: Costa Rica Rift (Hole 504B). Terra Nova. 11, 228–233. Furnes, H., Staudigel, H., 1999. Biological mediation in ocean crust alteration: how deep is the deep biosphere? Earth Planet Sci. Lett. 166, 97–103. Furnes, H., Muehlenbachs, K., Torsvik, T., Thorseth, I.H., Tumyr, O., 2001a. Microbial fractionation of carbon isotopes in altered basaltic glass from the Atlantic Ocean Lau Basin and Costa Rica Rift. Chem. Geol. 173, 313–330. Furnes, H., Staudigel, H., Thorseth, I.H., Torsvik. T., Muehlenbachs, K., Tumyr, O., 2001b. Bio-alteration of basaltic glass in the oceanic crust. Geochem. Geophys. Geosyst. 2, Paper number 2000GC000150. Furnes, H., Muehlenbachs, K., Tumyr, O., Torsvik, T., Xenophontos, C., 2001c. Biogenic alteration of volcanic glass from the Troodos ophiolite Cyprus. J. Geol. Soc. Lond. 158, 75–84. Furnes, H., Muehlenbachs, K., Torsvik, T., Tumyr, O., Shi, L., 2002. Bio-signatures in metabasaltic glass of a Caledonian ophiolite West Norway. Geol. Mag. 139, 601–608. Furnes, H., Muehlenbachs, K., 2003. Bioalteration recorded in ophiolitic pillow lavas. In: Dilek, Y., Robinson, P.T. (Eds.), Ophiolites in Earth History, Geol. Soc. Lond. Spec. Publ., vol. 218, pp. 415–426. Furnes, H., Banerjee, N.R., Muehlenbachs, K., Staudigel, H., de Wit, M.J., 2004. Early life recorded in Archean pillow lavas. Science 304, 578–581. Gold, T., 1992. The deep, hot biosphere. Proc. Natl. Acad. Sci. 89, 6045–6049. Hoefs, J., 1997. Stable Isotope Geochemistry. Springer, p. 201. Kontinen, A., 1987. An early Proterozoic ophiolite–the Jormua mafic–ultramafic complex, northeastern Finland. Precamb. Res. 35, 313–341. McCrea, J.M., 1950. On the isotope chemistry of carbonates and a paleotemperature scale. J. Chem. Phys. 18, 849–857. McKinley, J.P., Stevens, T.O., Westall, F., 2000. Microfossils and paleoenvironment in deep subfurface basalt samples. Geomicrobiol. J. 17, 43–54. Oremland, R.S., 1988. Biochemistry of methanogenic bacteria. In: Zehnder, A.J.B. (Ed.), Biology of Anaerobic Microorganisms. John Wiley & Sons, New York, pp. 641– 705. Peltonen, P., Kontinen, A., Huhma, H., 1996. Petrology and geochemistry of metabasalts from the 1.95 Ga Jormua ophiolite Northeastern Finland. J. Petrol. 37, 1359– 1383. Peltonen, P., Kontinen, A., Huhma, H., 1998. Petrogenesis of the mantle sequence of the Jormua ophiolite (Finland): melt migration in the upper mantle during Palaeoproterozoic continental break-up. J. Petrol. 39, 297–329. Peltonen, P., M¨antt¨ari, I., Huhma, H., Kontinen, A., 2003. Archean zircons from the mantle: the Jormua ophiolite revisited. Geology 31, 645–648. Ross, K.A., Fisher, R.V., 1986. Biogenic grooving on glass shards. Geology 14, 571–573.

137

Russel, M.J., Hall, A.J., 1997. The emergence of life from iron monosulphide bubbles at a submarine hydrothermal redox and pH front. J. Geol. Soc. Lond. 154, 377–402. Staudigel, H., Chastain, R.A., Yayanos, A., Bourcier, W., 1995. Biologically mediated dissolution of glass. Chem. Geol. 126, 147–154. Staudigel, H., Yayanos, A., Chastain, R., Davies, G., Verdurmen, E.A.T., Schiffman, P., Bourcier, R., De Baar, H., 1998. Biologically mediated dissolution of volcanic glass in seawater. Earth Planet. Sci. Lett. 164, 233–244. Staudigel, H., Furnes, H., Kelley, K., Plank, T., Muehlenbachs, K., Tebo, B., Yayanos, A., 2004. The Oceanic crust as a Bioreactor. In: Wilcock, W.S.D., DeLong, E.F., Kelley, D.S., Baross, J.A., Cary, S.C., (Eds.), The Subseafloor Biosphere at Mid-Ocean Ridges, Geophysical Monograph Series 144, in press. Thorseth, I.H., Furnes, H., Heldal, M., 1992. The importance of microbiological activity in the alteration of natural basaltic glass. Geochim. Cosmachim. Acta 56, 845–850. Thorseth, I.H., Furnes, H., Tumyr, H.O., 1995a. Textural and chemical effects of bacterial activity on basaltic glass: an experimental approach. Chem. Geol. 119, 139–160. Thorseth, I.H., Torsvik, T., Furnes, H., Muehlenbachs, K., 1995b. Microbes play an important role in the alteration of oceanic crust. Chem. Geol. 126, 137–146. Thorseth, I.H., Torsvik, T., Torsvik, V., Daae, F.L., Pedersen, R.B., 2001. Keldysh-98 Scientific Party Diversity of life in ocean floor basalt. Earth Planet. Sci. Lett. 194, 31–37. Thorseth, I.H., Pedersen, R.B., Christie, D.M., 2003. Microbial alteration of 0–30 Ma seafloor and sub-seafloor basaltic glasses from the Australian Antarctic Discordance. Earth Planet. Sci. Lett. 215, 237–247. Torsvik, T., Furnes, H., Muehlenbachs, K., Thorseth, I.H., Tumyr, O., 1998. Evidence for microbial activity at the glass-alteration interface in oceanic basalts. Earth Planet. Sci. Lett. 162, 165–176. Tuisku, P., 1997. An introduction to Palaeoproterozoic (Svecofennian) regional metamorphism in Kainuu and Lapland, Finland. In: Evins, P., Laajoki, K., (Eds.), Archean and Early Proterozoic (Karelian) Evolution of the Kainuu-Per¨apohja area, Northern Finland. A Guidebook for the Nordic Research Field Seminar Organised by the Universities of Oslo, Oulu and Turku, 2–10 June, 1997. Res. Terrae, Ser. A, No. 13, pp. 27–31. Urrutia, M.M., Beveridge, T.J., 1994. Formation of fine-grained metal and silicate precipitates on a bacterial surface (Bacillus subtilis). Chem. Geol. 116, 261–280. Van Lith, Y., Vasconceios, C., Warthmann, R., McKenzie, J.A., 2001. Organic matter trapped in sedimentary dolomite: A potential tracer to recognize microbial dolomite. European Union of Geoscience XI. J. Conf. Abstr., 56. Welch, S.A., Ullmann, W.J., 1996. Feldspar dissolution in acidic and organic solutions: compositional and PH dependence of dissolution rate. Geochim. Cosmochim. Acta 60, 2939–2948. Westall, F., Boni, L., Guerzoni, M.E., 1995. The experimental silicification of microorganisms. Palaeontology 38, 495–528. Westall, F., de Wit, M., Dann, J., van der Gaast, S., de Ronde, C.E.J., Gernecke, D., 2001. Early Archean fossil bacteria and biofilms in hydrothermally-influenced sediments from the Barberton greenstone belt. South Africa. Precamb. Res. 106, 93–116.