Ancient graphite in the Eoarchean quartz–pyroxene rocks from Akilia in southern West Greenland I: Petrographic and spectroscopic characterization

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Geochimica et Cosmochimica Acta 74 (2010) 5862–5883 www.elsevier.com/locate/gca

Ancient graphite in the Eoarchean quartz–pyroxene rocks from Akilia in southern West Greenland I: Petrographic and spectroscopic characterization Dominic Papineau a,*, Bradley T. De Gregorio b, George D. Cody a, Marc D. Fries a,c, Stephen J. Mojzsis d, Andrew Steele a, Rhonda M. Stroud b, Marilyn L. Fogel a a Geophysical Laboratory, Carnegie Institution of Washington, DC 20015, USA Materials Science and Technology Division, Naval Research Laboratory, DC 20375-5320, USA c Jet Propulsion Laboratory, NASA, Pasadena, CA 91109, USA d Department of Geological Sciences, University of Colorado at Boulder, Boulder, CO 80309, USA b

Received 19 October 2009; accepted in revised form 19 May 2010; available online 1 June 2010

Abstract Because all known Eoarchean (>3.65 Ga) volcano-sedimentary terranes are locked in granitoid gneiss complexes that have experienced high degrees of metamorphism and deformation, the origin and mode of preservation of carbonaceous material in the oldest metasedimentary rocks remain a subject of vigorous debate. To determine the biogenicity of carbon in graphite in such rocks, carbonaceous material must be demonstrably indigenous and its composition should be consistent with thermally altered biogenic carbon as well as inconsistent with abiogenic carbon. Here we report the petrological and spectroscopic characteristics of carbonaceous material, typically associated with individual apatite grains, but also with various other minerals including calcite, in a >3.83 Ga granulite-facies ferruginous quartz-pyroxene unit (Qp rock) from the island of Akilia in southern West Greenland. In thin sections of the fine-grained parts of Akilia Qp rock sample G91-26, mapped apatites were found to be associated with graphite in about 20% of the occurrences. Raman spectra of this carbonaceous material had strong G-band and small D-band absorptions indicative of crystalline graphite. Three apatite-associated graphites were found to contain curled graphite structures, identified by an anomalously intense second-order D-band (or 2D-band) Raman mode. These structures are similar to graphite whiskers or cones documented to form at high temperatures. Raman spectra of apatite-associated graphite were consistent with formation at temperatures calculated to be between 635 and 830 °C, which are consistent with granulite-facies metamorphic conditions. Three graphite targets extracted by focused ion beam (FIB) methods contained thin graphite coatings on apatite grains rather than inclusions sensu stricto as inferred from transmitted light microscopy and Raman spectroscopy. TEM analyses ˚ for apatite-associated graphite, of graphite in these FIB sections showed a (0 0 0 2) interplanar spacing between 3.41 and 3.64 A ˚ which is larger than the spacing of pure graphite (3.35 A) and may be caused by the presence of non-carbon heteroatoms in interlayer sites. Samples analyzed by synchrotron-based scanning transmission X-ray microscopy (STXM) also confirmed the presence of crystalline graphite, but abundances of N and O heteroatoms were below detection limit for this method. Graphite in the Akilia Qp rock was also found to occur in complex polyphase mineral assemblages of hornblende ± calcite ± sulfides ± magnetite that point to high-temperature precipitation from carbon-bearing fluids. These complex mineral assemblages may represent another generation of graphitization that could have occurred during the amphibolite-facies metamorphic event at 2.7 Ga. Several observations point to graphitization from high-temperature fluid-deposition for some of the Akilia graphite and our results do not exclude a biogenic source of carbon in graphite associated with apatite, but ambiguities remain for the origin of this carbon. Ó 2010 Elsevier Ltd. All rights reserved.

*

Corresponding author. Tel.: +1 202 478 8908; fax: +1 202 478 8901. E-mail address: [email protected] (D. Papineau).

0016-7037/$ - see front matter Ó 2010 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2010.05.025

Petrography and crystal structure of Akilia graphite

1. INTRODUCTION Eoarchean volcano-sedimentary sequences occur as supracrustal enclaves in highly deformed granitoid gneiss terranes in southern West Greenland (Isua Supracrustal Belt and Akilia association; Nutman et al., 1996). The supracrustal rocks are dominated by amphibolites that are associated at the outcrop scale with lithologies such as banded magnetite + quartz rocks (banded iron formations, or BIFs) and quartz + biotite + garnet schists (“metapelites”). The rocks of probable metasedimentary origin in these supracrustal enclaves are significant because if their petrogenesis and age can be constrained, they hold the potential to provide direct information about the nature of the earliest environments in which life could have existed. For instance, Mojzsis et al. (1996) reported geochemical and isotopic evidence interpreted to represent biological activity more than 3.8 billion years ago, based on 13C-depleted graphite associated with apatite in a quartz-pyroxene (Qp) rock from the island of Akilia in the southernmost part of the Eoarchean Itsaq Gneiss Complex in West Greenland (Nutman et al., 1996; Manning et al., 2006; Cates and Mojzsis, 2006). As a consequence of this interpretation, the Akilia Qp rock has been the subject of intense scrutiny and disagreements remain regarding the protolith of this rock, its emplacement age, and the existence in it of apatite-associated graphite. The original interpretation that the protolith of the Akilia Qp rock formed from precipitation in seawater (McGregor and Mason, 1977; Mojzsis et al., 1996; Nutman et al., 1997; Friend et al., 2002; Mojzsis and Harrisson, 2002a, b; Dauphas et al., 2004, 2007; Manning et al., 2006; McKeegan et al., 2007; Nutman and Friend, 2006) contrasts with the view that this rock is a result of the metasomatic alteration of an ultramafic protolith (Fedo and Whitehouse, 2002a, b; Whitehouse and Fedo, 2003; Bolhar et al., 2004; Andre´ et al., 2006; Fedo et al., 2006; Whitehouse et al., 2009). A detailed analysis of the evidence for a chemical sedimentary origin of ferruginous Qp rocks on Akilia was previously presented by Manning et al. (2006). In terms of mineralogical composition, the Akilia Qp rock is consistent with a sedimentary protolith: it is dominated by quartz, clinopyroxene (hedenbergite), orthopyroxene (ferrosilite), amphibole (grunerite–hornblende), and magnetite. Indeed, pyroxenes and amphiboles in highly metamorphosed BIFs are expected from high-temperature reactions between ferrous carbonate and quartz, and other silicates or magnetite and quartz with metasomatic fluids (French, 1964; James, 1966; Schreyer et al., 1978; Klein, 1978, 2005). A chemical sedimentary protolith for the Akilia Qp rock is further supported by the presence of mass-independently fractionated sulfur isotopes in chalcopyrite-pyrrhotite associations (Mojzsis et al., 2003), isotopically heavy iron (Dauphas et al., 2004, 2007), oxygen isotopes (Manning et al., 2006), and major-, minor-, trace-elements and rare earth element (REE) distributions (Mojzsis and Harrison, 2002a, b; Friend et al., 2002; Nishizawa et al., 2005; Manning et al., 2006; Johannesson et al., 2006). Alternatively, REE patterns and trace-element abundances have been used to argue instead for a metasomatized igneous ultramafic

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protolith for the same rock (Fedo and Whitehouse, 2002a, b; Bolhar et al., 2004). The sedimentary protolith interpretation has been further questioned based on the inability to verify that mass-independently fractionated sulfur isotopes are present in the Akilia Qp rock (Whitehouse and Kamber, 2005), Si-isotope distributions similar to metamorphic silica (Andre´ et al., 2006), and Fe-isotopes interpreted to have been transported at a later time into the Akilia Qp rock (Fedo and Whitehouse, 2007). Despite these alternative interpretations, Eiler (2007) concluded in a recent review of this debate that the most reasonable interpretation for the origin of the Akilia Qp rock is that it is was originally a chemical sedimentary precipitate. It is also worth noting that several occurrences of highly metamorphosed Precambrian BIFs exist elsewhere in the world with mineralogical and geochemical characteristics similar to those of the Akilia Qp rock (Klein, 2005). Age constraints on the Qp rock were initially determined to be >3.85 Ga based on U–Pb zircon geochronology on neighboring orthogneisses to the enclave (Nutman et al., 1997), but this also has been the subject of debate (Whitehouse and Kamber, 2005). Depth profile analyses of single zircons from a structurally transecting orthogneissic sheet on Akilia revealed evidence for several phases of zircon growth at 3.83, 3.73, 3.65, and 2.66 Ga, which lends support to the interpretation that the zircons record the emplacement age (3.83 Ga) of the tonalitic dykes (Mojzsis and Harrison, 2002a, b; Manning et al., 2006). This Eoarchean (3.83 Ga) orthogneissic sheet was mapped as cross-cutting an ultramafic unit in contact with and synmetamorphic to the Akilia Qp rock (Manning et al., 2006). Conversely, it has been suggested that the U–Pb zircon geochronology data from Akilia orthogneisses indicate that 3.83 Ga zircons were captured in a 3.62 Ga tonalitic melt that was subsequently metamorphosed around 2.7 Ga and again around 1.7 Ga, and that all zircon ages older than ca. 3.65 Ga are inherited (Whitehouse and Kamber, 2005; Whitehouse et al., 2009). However, ages for the zircons summarized in Manning et al. (2006) are consistent with the metamorphic history of other rocks in the vicinity of Akilia (Kinny, 1986; Nutman et al., 1996; 2002) and collectively record a complex protracted igneous and metamorphic history for the Itsaq Gneiss Complex that began at ca. 3.83 Ga. The common association of 13C-depleted graphite with apatite in the Akilia Qp rock was proposed as a biological signature (Mojzsis et al., 1996). While there have been difficulties in verifying these common occurrences of apatite + graphite in sample G91-26 (Lepland et al., 2005; Nutman and Friend, 2006), McKeegan et al. (2007) confirmed the presence of these mineral associations in the same sample, argued that graphite occurred as inclusions inside apatite, and also reproduced the in situ measurement of 13C-depleted graphite. Age arguments based on U–Pb geochronology of Akilia apatites have been raised to argue that the apatites in the Akilia Qp rock formed only around 1.7 Ga (Whitehouse et al., 2009; Sano et al., 1999). However, the closure temperature for Pb in apatite is around 450 °C (Cherniak et al., 1991; Mojzsis et al., 1999) and the metamorphic history of the Itsaq Gneiss Complex includes a well-documented 1.7 Ga metamorphic events in the region

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(Baadsgaard et al., 1986). Hence the age of the apatites most likely records the age of this 1.7 Ga thermal event, which reset the U–Pb geochronometer of apatites. In light of the debate over the petrogenesis of this rock and the documentation of apatite-associated graphite, now is the time to take an integrated approach to identify and study the occurrences of graphite in quartz–pyroxene rocks from Akilia. In this work, we present comprehensive petrographic surveys of apatite + graphite in the Akilia Qp rock and report correlated in situ analyses of the structural and crystallographic properties of graphite occurrences with the goal of better constraining their possible origins and further assessing the possibility that biological information could be preserved in this rock. 2. ANALYTICAL METHODS 2.1. Optical microscopy Optical microscopic surveys of apatite-associated graphite were performed in thin sections from two samples, including the original sample G91-26 (also designated ANU-92197) (Mojzsis et al., 1996) and SM-9711 (Mojzsis et al., 2003). An Olympus Bx61 microscope with 10, 40, and 100 long working-distance objectives was used without immersion oil to map occurrences of apatite-associated graphite. 2.2. Raman spectroscopy Confocal laser Raman spectroscopy was performed with a WITec a-SNOM imaging system at the Carnegie Institution of Washington. Excitation was obtained by scanning the stage on targets with a frequency-doubled solid-state YAG laser (k = 532 nm) operating at between 1 and 8 mW output power as measured at the sample focal plane under 1000 magnification. A 100 lm diameter optic fiber was used to deliver Raman scattered photons to the Peltiercooled EMCCD detector. Pixel size was around 360 nm2 and each spectrum was acquired for between 0.6 and 2 seconds depending on signal/noise for a particular measurement. A 600 lines/mm spectrometer was used for all analyses and gave a spectral resolution of about 2.5 cm1. Hyperspectral images of specific mineral phases were produced by mapping the peak intensity of parabolic peak fits from spectral datasets and these images can be assembled to include information on chemical composition and latent crystal strain (Steele et al., 2007; Fries and Steele, 2008; Bernard et al., 2008). Individual spectra shown in this work represent averages of selected regions with similar spectral characteristics. 2.3. Scanning electron microscopy Scanning electron microscopy (SEM) in secondary electrons (SE) and back scattered electron (BSE) imaging modes was used to locate graphite at the surface of the thin section and to characterize the composition of apatite by energy dispersive X-ray spectroscopy (EDS). These analyses were performed at the Carnegie Institution of Washington with

a JEOL 6500F field-emission SEM with a fully automated EDS and BSE detection system. Standard operating conditions for SEM imaging and EDS analysis were a 15 kV accelerating voltage potential and an electron beam current of 0.5 or 1 nA. Polished thin sections were prepared with aluminum oxide and cerium oxide polishing powder for the last 3 and 0.5 lm polishing steps, respectively. The polished sections were then manipulated with gloves, cleaned with deionized H2O, stored in muffled aluminum foil, and coated with gold prior to SEM analysis. 2.4. Focused ion beam milling Focused ion beam (FIB) milling allowed for site-specific extraction of small samples appropriate for transmission electron microscopy (TEM) and scanning transmission Xray microscopy (STXM). Sections extracted using FIB methods were orthogonal to the surface, which allowed for direct visualization of the contact relationships between apatite and graphite in the Akilia Qp samples. In this way, true graphite inclusions (i.e. graphite completely enclosed within a mineral like apatite) could be confidently distinguished from coatings. FIB extraction was performed using a FEI Nova 600 DualBeam FIB-SEM at the Naval Research Laboratory. Once the selected graphite-apatite had been characterized below the surface by transmitted light microscopy and Raman spectroscopy, the target was located by SEM imaging and a protective Pt “strap” was deposited on the surface. A focused 30 keV Ga+ primary beam was then used to sputter away material from around the deposited Pt mask, leaving a 1 lm thick vertical section. A custom-shaped copper micro-tweezer or tungsten needle was moved towards the sample and attached with a Pt weld. The bottom and sides of the section were then removed by the Ga+ beam. Once the section was free from the matrix, further sputtering of the sides thinned the section down to 100 nm. During this phase, care was taken to simultaneously reduce the Ga+ beam current from 20 nA to less than 50 pA, which also decreased the spot size of the ion beam to 10 nm. This was done in order to achieve sections with high-quality surfaces (Zega et al. 2007; Wirth 2009) and to minimize the amount of Ga+ implantation in the section. The apatite + graphite section, attached to the tungsten needle, was then transferred to a Cu or Mo TEM half-grid before the final thinning steps were performed, while those held by Cu microtweezers were thinned without transfer, since the microtweezers could be folded into a TEM-appropriate sample holder. 2.5. Ultramicrotomy A graphite particle was also extracted ex situ from the thin section (Supplementary Electronic annex 1) and sectioned by ultramicrotomy. Approximately 1 ll of 28% HF was micro-pipetted on top of the target surface in the thin section to dissolve apatite and the surrounding quartz, leaving the acid-insoluble graphite intact. The free graphite particle was then manipulated with a micro-needle by electrostatic force and was inserted into a molten bead of

Petrography and crystal structure of Akilia graphite

sulfur, which crystallized around the graphite as it cooled to room temperature. The sulfur bead was subsequently glued onto an epoxy stub and microtomed with a diamond knife into 125 nm slices. The microtome sections of graphite and surrounding sulfur were transferred to 200 mesh, thin-bar, Cu TEM grids coated with a silicon monoxide support film. The sulfur was removed from the sample by exposing the grid to 70 °C air for a few minutes over a hot plate, during which the sulfur sublimed. 2.6. Transmission electron microscopy Samples were analyzed using a JEOL 2200FS field-emission transmission electron microscope (TEM) at the Naval Research Laboratory at an operating voltage of 200 keV. High-resolution TEM imaging was used to visualize the crystalline relationships between graphite and apatite, as well as to identify additional nanoscale minerals in the sample. Crystal structure orientations were derived from selected area electron diffraction patterns with a camera calibrated to polycrystalline aluminum standards. Mineral compositions were determined using an EDS system attached to the TEM. Full EDS hyperspectral images were obtained with the 2200FS operated in scanning (STEM) mode with a nominal probe size of 0.7 nm in order to map the elemental composition of the sections. High-angle annular dark-field (HAADF) imaging was used to quickly distinguish graphite, apatite, and quartz domains due to the large differences in atomic mass contrast. Sections prepared by FIB milling were inserted directly into the TEM without a conductive coating. However, ultramicrotomed sections placed on SiO-coated TEM grids required a light amorphous carbon film to mitigate charging effects during analysis. Depending on the grid material (Cu or Mo), a small Cu Ka or Mo La emission peak was present (8.05 or 2.29 keV, respectively) in EDS spectra due to stray electron fluorescence. 2.7. Synchrotron-based scanning transmission x-ray microscopy Synchrotron-based scanning transmission x-ray microscopy (STXM) analyses were performed at the Canadian Light Source (CLS) on beamline 10ID-1 and at the Advanced Light Source (ALS) on beamline 5.3.2. X-ray absorption near-edge structure (XANES) spectra of the C K-edge, N K-edge, and O K-edge were acquired on two FIB sections containing graphitic material and one microtome section from a graphite particle that had been removed from its associated apatite grain (Supplementary Electronic annex 1) prepared on a silicon-monoxide-coated TEM grid. Beamline 10ID-1 at the CLS uses X-rays spanning the range 130–2500 eV that are generated with an elliptically polarized undulator (EPU) inserted in the 2.9 GeV, 250– 100 mA, synchrotron storage ring. A monochromated X-ray beam was achieved by varying the EPU gap distance and refined using a Ni-coated grating monochromator with 250 lines/mm, providing an energy resolution greater than 0.1 eV (Kaznatcheev et al., 2007). In the case of beamline

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5.3.2 at the ALS, soft X-rays were generated via a bending magnet while the electron current in the storage ring was held constant in “topoff mode” at 500 mA at a storage ring energy of 1.9 GeV. Dispersive and non-dispersive exit slits were set at 25 lm. After sample insertion in either STXM at the CLS or at the ALS, the chamber was evacuated to 100 mTorr and back-filled with He. For both STXM instruments, the monochromatic X-ray beam is focused on the sample using a Fresnel zone plate objective and an ordersorting aperture yielding a focused X-ray beam spot of 30–40 nm. Both ALS and CLS beamlines used similar STXM instrumentation and control software. STXM data are typically acquired as hyperspectral images (or stacks), from which XANES spectra of regions of interest may be extracted (Jacobsen et al., 2000). The energy resolution of K-edge XANES spectra acquired in this way are determined by the energy step between sequential X-ray absorption images (Supplementary Electronic annex 6). Photon transmission from multiple pixels in a selected region of interest is summed to provide high quality X-ray absorption spectra with good signal-to-noise. XANES optical density (OD) spectra are calculated by calibrating the transmitted intensity (I) of the sample to the background transmission (I0) as OD = log(I/I0) (Sto¨hr, 1992). Principal component analysis and clustering algorithms are applied to the hyperspectral image datasets to automatically identify regions with similar spectral features (Lerotic et al., 2004). 3. RESULTS 3.1. Associations of graphite with apatite Optical microscopy mapping of the fine-grained areas of sample G91-26, corresponding to regions that have experienced relatively low degrees of silica mobility (Nutman et al., 1997), revealed carbonaceous associations with apatite in roughly 20% of apatite crystals (Table 1). One optical map of a 1  2 cm thin section (90 lm thick) of G91-26 revealed 623 apatite grains, among which 112 (18%) had graphite coatings (Fig. 1a). Surveys on other thin sections from samples of the same outcrop showed similar distributions of apatite grains and associations with graphite (Fig. 1b), although the coarse-grained portions contain significantly less apatite-associated graphite (Fig. 1c, d; Table 1). Apatites counted for these surveys were between 3 and approximately 100 lm in size and graphite associations varied between 1 and 45 lm. Examples of typical apatite + graphite associations are shown in Fig. 2. Small micron-size fluid inclusions that occasionally contained opaque material were also often present within the quartz near apatite + graphite associations (Fig. 2c–h). While these inclusions were not studied systematically, only a few were observed to be fluid-bearing. Most apatite + graphite associations were found to be present in quartz and only a few were found in ferruginous silicates. In 55 apatite + graphite associations analyzed by Raman spectroscopy, the carbonaceous material appears to be generally highly graphitized, and no other opaque minerals (e.g. magnetite or sulfides) were detected. A selection

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Table 1 Grain count statistics for thin sections in Fig. 1. Number of grains G91-26b

GR9707b

Fine-grained (a)

Fine-grained (b)

Coarse-grained (c)

Coarse-grained (d)

Apatite grains (euhedral >3 lm) Apatite grains in quartz Apatite grains in Fe-silicatesa

623 488 135

405 296 109

59 59 0

28 26 2

Apatite with graphite coatings Apatite with graphite in quartz Apatite with graphite in Fe-silicatesa

112 104 8

27 24 3

7 7 0

1 1 0

Graphite + hornblende + calcite ± magnetite ± pyrrhotite ± chalcopyrite ± pentlanditec

11

13

n.d.

n.d.

a

The Fe-silicate matrix includes pyroxenes and amphiboles. Letters in parentheses refer to the thin sections in Fig. 1. c These graphite associations were mapped in other sub-samples of G91-26, but from the same band corresponding approximately to the area of the thin sections in Fig. 1a and b. b

of hyperspectral Raman images of apatite-associated graphite is shown in Fig. 3a–d and typical averaged Raman spectra are presented in Fig. 3e. Raman spectral characteristics of the full dataset are listed in Table 2. Raman spectra of graphite associated with apatite often exhibit a small range of D-band intensities relative to strong G-bands showing small variations in the graphite. Strong 2D overtone Raman modes around 2700 cm1 were observed in 1 lm-sized sub-domains within three graphite grains (Fig. 4), and these regions have Raman spectral characteristics similar to those of graphite “whiskers” or cones (Tan et al., 2001; Shang and Jiang, 2005), implying that they are due to unique, localized graphitic structures within the grains. These structures tend to occur in small groups of three to four graphitic subdomains with sizes varying between 500 and 1000 nm. Secondary electron images revealed that graphite flakes are exposed on the surface of sample G91-26 (Supplementary Electronic annex 3), likely disrupted by polishing during sample preparation. This mechanical surface damage of graphite was also observed in Raman spectra of the sample surface by a significantly more intense D-band than that of the original, unpolished graphite in the rock sample (Supplementary Electronic annex 2; Pasteris, 1989). Spectral artifacts due to surface damage were avoided in the present study by limiting Raman analyses to material at least one micron below the surface of the thin section. 3.2. Graphite associated with calcite and other minerals Graphite in sample G91-26 occasionally occurs in complex, polyphase, mineral assemblages that included magnesian ferrohornblende ± calcite ± magnetite ± pyrrhotite ± chalcopyrite with occasional micrometer-size pentlandite zones (Fig. 6; Table 1). These complex assemblages were often located near the edge of large magnesian ferrohornblende (referred hereafter as hornblende) grains and tended to form outgrowths into quartz (Fig. 6a, c and e). Calcite in the Akilia Qp rock occurred as rare patches a few tens of micrometers in size and contained

variable amounts of Mg and Fe (typically less than 1%). Imaging by BSE revealed that when calcite was present, graphite was also present and sometimes, although not always, it was in direct contact with the calcite (Fig. 6b, d, h). Notably, apatite has so far not been found in any of the analyzed two-dozen or so occurrences of such mineral assemblages. 3.3. Structural characteristics of graphite in the Akilia Qp rock Raman spectra collected on carbonaceous material associated with apatite had a small D-band intensity relative to the corresponding G-band intensity, indicative of graphite (Table 2), whereas graphite associated with calcite and ferruginous minerals had a slightly higher D/G peak intensity ratio and resolvable D2-bands (around 1620 cm1) that indicate slightly less crystalline order. The positions of the D- and G-bands in graphite associated with either apatite or calcite were similar but varied between 1344 and 1358 cm1 for the D-band and between 1573 and 1581 cm1 for the G-band (Fig. 7a and b). The range of full width at half maximum (FWHM) for the G-band in these spectra was also relatively narrow (between 22 and 29 cm1), but the range of FWHM for the D-band (between 10 and 46 cm1) was slightly larger. The most notable distinctions between graphite spectra was in the intensity of their D-bands relative to their G-bands and the different characteristics of their 2D-bands (Fig. 7). While the D/G band intensity ratios were generally smaller for graphite associated with apatite compared to those of graphite associated with calcite and ferruginous minerals, no distinctions between these graphites were observed in their 2D/G intensities, 2D positions, and 2D FWHM (Fig. 7c and d). In graphite sub-domains showing increased 2D-band intensity, the 2D-band position ranged between 2685 and 2700 cm1 and was positioned at lower wavenumbers than the 2D-band of typical graphite, which varied between 2698 and 2710 cm1 (Fig. 7c). Two of these structures had particularly large D/G intensity ratios between 0.44 and 0.52 (Fig. 7d).

Petrography and crystal structure of Akilia graphite

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Fig. 1. (a) A 90 lm thick thin section of the Akilia Qp rock showing the locations of euhedral apatite grains larger than 3 lm (blue dots) and occurrences with graphite (numbered red dots). Thin sections in a) and (b) are fine-grained parts of sample G91-26, while (c) is a coarsegrained part of sample G91-26, and (d) is a coarse-grained portion of sample GR9707.

In optical microscopy and Raman surveys of the sample, about 30% of the graphite + apatite associations appeared to be “inclusions”, while about 11% appeared to be “invaginations” (Table 1). However, these distinctions are hampered by incomplete knowledge of the spatial relationship between apatite and graphite in the viewing direction. In order to directly observe the spatial relationships of these associations, FIB in situ lift-out sections were

prepared orthogonal to the thin section surface from three selected apatite + graphite targets (Fig. 5). In each of these extracted sections, the carbonaceous material formed a thin coating on the apatite grains rather than occurring as true inclusions or invaginations as interpreted from optical microscopy or Raman spectroscopy (Supplementary Electronic annex 5). Our observations of graphite coatings on apatite grains that appeared as inclusions

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Fig. 2. Example of apatite grains associated with carbonaceous material in sample G91-26 by transmitted light microscopy where (a and b) are typical occurrences in clear quartz, and (c–h) are occurrence near fluid inclusions in the surrounding quartz. Graphite coatings on apatite are represented by (a, c, d, f–h), while graphite associations that appear as “inclusions” in apatite are shown in (b and e).

imply that “inclusions” or “invaginations” of graphite on or in apatite are difficult to distinguish from coatings by analyzing the sample from a single, fixed viewing direction.

TEM analysis of one of these FIB sections revealed a sharp contact boundary between graphite and apatite (Supplementary Electronic annex 4c). The apatite grain was a single-crystal, as confirmed by selected area electron diffrac-

Petrography and crystal structure of Akilia graphite

(a)

(b)

(c)

(d)

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(e)

G91-26C_29 G91-26C_37

G91-26C_66 G91-26C_112 G91-26D_08

G91-26D_09

G91-26D_15

G91-26F_I_gra04b

G91-26F_I_gra03c G91-26F_I_gra06

1200

1500

1800

2100

2400

2700

Wavenumber (cm-1) Fig. 3. Transmitted light photomicrographs (a–d) of apatite grains associated with graphite with corresponding Raman hyperspectral images. In the hyperspectral images the intensity of the main Raman vibrational excitation mode in quartz is represented in blue, while that of apatite and graphite are represented in turquoise and in red, respectively. (e) Representative average Raman spectra of graphite associated with apatite or calcite.

tion (SAED), HRTEM imaging, and the presence of continuous Bragg-diffracted bend contours in lower-magnification TEM images (Supplementary Electronic annex 4a). Fast Fourier transform (FFT) analysis of lattice fringes in the apatite (Supplementary Electronic annex 4e) revealed ˚ (2 0 1) lattice planes perpendicular to the apatite– 3.54 A graphite interface. FFT analysis of the corresponding

graphitic lattice fringes close to the interface (Supplementary ˚ spacing, conElectronic annex 4d) showed a 3.57 to 3.64 A sistent with expanded (0 0 0 2) planes in the graphite. The graphite away from the grain boundary was polycrystalline with a <200 nm crystallite size. Sub-micron inclusions of chalcopyrite and chlorapatite were present within the poorly crystalline graphite (Supplementary Electronic annex 4b).

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Table 2 Raman spectral characteristics of graphite in the Akilia Qp rock. Mineral association

D pos (cm1)

D FWHM (cm1)

G pos (cm1)

G FWHM (cm1)

D/G (intensity)

2D pos (cm1)

2D FWHM (cm1)

2D/G (intensity)

T (°C) Beyssac 

T (°C) CD Cody 

T (°C) CG Cody 

G91-26C_4 G91-26C_19 G91-26C_20 G91-26C_21 G91-26C_22 G91-26C_23 ave# G91-26C_24 G91-26C_26 G91-26C_27 G91-26C_29 G91-26C_31 G91-26C_37 G91-26C_38a G91-26C_38b G91-26C_57 G91-26C_66 G91-26C_74 G91-26C_112 G91-26D_01 ave# G91-26D_03 ave# G91-26D_05 G91-26D_08 G91-26D_09 G91-26D_10 G91-26D_14 G91-26D_15 G91-26C_23 W1 G91-26C_23 W2 G91-26C_23 W3 G91-26D_01 W4 G91-26D_01 W5 G91-26D_01 W6 G91-26D_03 W7 G91-26D_03 W8 G91-26F_I_01 G91-26F_II_03a G91-26F_II_03b G91-26F_II_03c G91-26F_II_06

ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap ap cal, cal, cal, cal, cal,

1358 1358 1346 1348 1353 1350 1351 1353 1347 1346 1353 1349 1348 1355 1350 1348 1356 1348 1346 1349 1350 1350 1350 1346 1352 1349 1348 1350 1353 1344 1345 1345 1347 1345 1346 1345 1344 1346 1346

46.4 – 24.0 36.1 42.0 41.3 – – 33.6 38.2 35.7 41.8 42.6 – 45.2 40.4 40.4 44.4 40.5 42.1 42.9 44.9 44.4 38.3 30.7 41.0 33.0 29.8 33.2 35.3 24.8 36.7 25.9 37.0 40.4 40.8 40.9 41.1 42.8

1579 1574 1575 1573 1577 1577 1579 1576 1575 1578 1577 1576 1578 1578 1580 1580 1580 1581 1576 1576 1577 1577 1577 1577 1577 1577 1578 1578 1578 1573 1576 1576 1577 1576 1577 1576 1576 1577 1577

24.7 22.5 26.3 26.4 24.9 28.9 22.0 26.7 24.6 26.7 22.2 26.3 26.5 24.5 22.6 22.5 24.0 23.7 22.3 25.5 24.7 23.7 25.3 23.7 22.8 24.6 27.5 25.7 27.6 24.8 23.3 21.8 27.0 25.9 22.5 23.7 22.9 22.0 22.6

0.09 0.05 0.20 0.23 0.19 0.17 0.12 0.16 0.09 0.11 0.05 0.09 0.20 0.04 0.09 0.07 0.21 0.01 0.06 0.12 0.07 0.05 0.16 0.07 0.11 0.07 0.52 0.19 0.13 0.24 0.11 0.20 0.44 0.13 0.57 0.59 0.61 0.51 0.57

2704 2703 2704 2698 2704 2703 2707 2700 2703 2703 2705 2702 2706 2707 2710 2710 2704 2709 2704 2703 2706 2705 2705 2704 2705 2706 2694 2697 2700 2685 2700 2692 2697 2692 2700 2701 2698 2701 2700

76.6 64.7 62.1 74.8 71.4 72.8 68.9 86.0 72.7 69.7 67.3 69.9 71.2 71.2 69.7 70.6 76.5 75.9 69.8 72.0 69.4 71.3 72.2 74.2 75.4 70.1 43.0 31.7 56.5 45.2 60.1 51.2 48.0 46.6 70.6 72.7 77.5 75.6 75.2

0.33 0.30 0.26 0.44 0.26 0.37 0.33 0.30 0.32 0.42 0.32 0.35 0.54 0.41 0.30 0.19 0.31 0.28 0.30 0.43 0.34 0.33 0.37 0.33 0.33 0.34 1.40 2.28 0.58 0.92 0.37 0.63 1.09 1.01 0.24 0.31 0.29 0.18 0.24

570 609 588 567 559 582 590 597 601 576 625 599 607 627 602 612 571 635 615 594 612 617 579 582 570 603 489 569 572 546 599 572 588 562 452 456 439 464 452

764 – 829 793 776 778 – – 801 787 795 777 775 – 767 781 776 770 781 776 774 768 769 787 809 779 802 812 802 796 826 792 823 791 774 776 780 779 774

1125 1164 1099 1097 1122 1054 1173 1090 1128 1091 1170 1097 1095 1130 1163 1168 1115 1135 1168 1111 1126 1143 1116 1144 1159 1127 1078 1108 1075 1124 1151 1177 1085 1106 1162 1139 1158 1174 1162

#

mag, ccp, po, hbl hbl hbl hbl mag, po ccp, pn, hbl

These are average graphite spectra that exclude the area around curled graphite structures.

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Spectrum name

Petrography and crystal structure of Akilia graphite

(b)

(a)

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(c)

W6

W3

W5

W2

W8

W4

W7 W9

(d) W1 W2

Relative intensity (arbitrary units)

avg. a)

W4 W6

avg. b)

W7 W8

avg. c)

1200

1500

1800

2100

2400

2700

Wavenumber (cm-1) Fig. 4. (a–c) Raman hyperspectral images of three graphite coatings on apatite crystals (in transmitted light) containing sub-domains of curled graphite structures. Representative average spectra are shown in (d) for the curled graphite structures and average graphite (see also Table 2 for spectral parameters).

A second FIB section (Fig. 8a) was extracted in situ from graphite that contained sub-domains generating intense 2D peaks in Raman spectra (Fig. 4a). The extracted section contained a flat band of graphite between the quartz and apatite, and a rounded carbonaceous structure within the quartz (Fig. 8b–e). Graphitic lattice fringes in the band of poorly crystalline graphite were measured to ˚ near the outer edges of the band and 3.63 A ˚ be 3.41 A

near the center of the band. These lattice spacings are similar to the polycrystalline graphite in the previous FIB section. A 100 nm thick band of Fe-rich material was present between the graphite coating and the quartz (Fig. 8f). Another micron-sized mineral phase was located in the quartz adjacent to the rounded carbonaceous structure (Fig. 8c). This silicate mineral contained Al, Mg, K, and Fe, as detected by EDS (Fig. 8d). However, we were

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Fig. 5. (a–f) Three targets of graphite associated with apatite extracted by the FIB-SEM method. Transmitted light images (a, c, and e) were used to target graphite + apatite associations from below the surface, since they could not be visualized directly by SEM. Different stages of the FIB process (b, d, and f) revealed that in all cases so far, graphite occurs as a 500 nm coating at the interface of apatite and quartz.

not able to obtain sufficient quality SAED patterns for mineral identification, and the sample was too thick for HRTEM imaging. The section was returned to FIBSEM for further thinning, after which the quartz and apatite appeared amorphous as a result of FIB radiation damage, and lattice fringes could only be observed in the graphite. This re-thinning step removed excess quartz covering the rounded carbonaceous structure, and the overall size of the structure decreased by approximately 50% (Fig. 8c and e), which is consistent with a rounded, three-dimensional structure within the section. This feature was a possible source of the 2D-band signature observed in the bulk sample, although subsequent Raman spectroscopy of the FIB section did not confirm the presence of a large 2D-band. The lack of observable 2D band signature in the extracted FIB section is likely due to the small volume of the 2D-producing feature relative to the size of the Raman laser spot but may also be affected by surface amorphization by the Ga+ ion beam or hydrocarbon deposition during SEM and TEM imaging.

STXM analyses of graphite at the C K-edge in microtome and FIB sections revealed a high degree of crystallinity. C-XANES spectra of graphite in microtome and FIB sections showed similar characteristics with a prominent p* peak around 285.3 eV and a sharp p* exciton peak around 291.4 eV (Fig. 9c). Notably, no unambiguous molecular functional groups were detected in the energy range between the 1s ? p* and 1s ? r* transitions. However, spectra from microtome sections (Fig. 9d) and the FIB section differed slightly in the structure of the region between 288 and 292 eV. In C-XANES spectra from microtome sections, the r* peak rose sharply and exhibited a well-defined exciton absorption at 291.4 eV, while in spectra from the FIB section, the r* absorption was broad and contained only minimal exciton intensity. The lack of a strong r* exciton from graphite in the FIB section is consistent with the sub-micron poorly crystalline graphite and carbonaceous material observed in this section. In addition, the rounded carbonaceous structure observed in the FIB section appeared to be connected to the main graphite band

Petrography and crystal structure of Akilia graphite

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Fig. 6. Back-scattered electron images of examples of regions where graphite is associated with calcite and other minerals. These mineral assemblages are typical for hydrothermal skarns in metamorphosed rocks and they are shown in their larger petrographic context. The following mineral abbreviations were used: quartz = Qtz; hornblende = Hbl; pyroxene = Pyr; magnetite = Mag; graphite = Gr; calcite = Cal; pyrrhotite = Po; pentlandite = Pn; chalcopyrite = Ccp.

in STXM images acquired between 284.4 and 290.2 eV (Fig. 9b). Since a connecting bridge of carbonaceous matter between the two structures was not observed by TEM, it is possible that this feature is a STXM artifact, possibly due to strain fields or defects, within the quartz. This feature

does align with an annealed fracture present within the apatite grain on the opposite side of the graphite band (Fig. 8b). N-XANES and O-XANES spectroscopy of the graphite band and rounded carbonaceous structure failed to detect significant abundances of these elements.

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(b)

(a)

(c)

(d)

Fig. 7. Plots of Raman spectral characteristics for graphite associated with apatite (white boxes) or with calcite and other ferruginous minerals (grey disks) and for curled graphite structures (black boxes) in the Akilia Qp rock: (a) D-band FWHM (full width at half maximum) vs. D-band position, (b) G-band FWHM vs G-band position, (c) 2D position vs 2D FWHM, and (d) 2D/G vs. D/G intensity ratio.

4. DISCUSSION 4.1. Petrographic distribution of graphite in the Akilia Qp rock In the Akilia Qp rock, apatite grains are associated with graphite in roughly 20% of cases. This result unequivocally resolves the confusion created by previous failed attempts to verify the common occurrence of apatite + graphite associations in Akilia Qp rock sample G91-26 (Lepland et al., 2005; Nutman and Friend, 2006). The reason for these difficulties may have been partly due to the methods of grain separation used by these groups to find graphite associated with apatite grains, rather than in situ petrographic studies of thin sections. Graphite associated with apatite observed by optical microscopy and Raman imaging spectroscopy often appeared as “inclusions” or “invaginations” (Figs. 2–4), but FIB of three possible candidate targets has not revealed graphite completely enveloped by apatite (Fig. 5). These observations suggest that “inclusions” or “invaginations” of graphite in apatite cannot be unambiguously distinguished from coatings with twodimensional microscopic imaging methods applied in a single, fixed, viewing direction, as was previous argued (Mojzsis et al., 1996). Three-dimensional reconstructions of these graphite associations with apatite by confocal Raman spectroscopy (McKeegan et al., 2007) also are

challenging due to the very strong absorption of graphite, which can still be detected if the “confocal plane” analyzed is beneath a graphite coating, which yields a graphite signal in a focal plane where there is no graphite. This is illustrated in Supplementary Electronic annex 5 where four different “confocal planes” spanning 1.5 lm all included graphite absorptions, even though the graphite occurred as a coating less than 0.5 lm in thickness (Fig. 8). Optical microscopy revealed that, while most micronsized fluid inclusions in the Akilia Qp rock were empty, they were often located within the quartz near apatite + graphite associations (Fig. 2c–h). Some of these fluid inclusions contain graphitic material and thus fluids percolating through the rock may have been carbon-bearing (van Zuilen et al., 2009). In addition, the mere presence of fluid inclusions near apatite + graphite associations raises the possibility that some graphite in these associations may have been in contact with or formed from fluid-precipitation. The close association of graphite with complex mineral assemblages, that excluded apatite but included hornblende ± calcite ± magnetite ± pyrrhotite ± chalcopyrite ± pentlandite (Fig. 6), further suggests fluid-precipitation of this graphite under high-temperature conditions. The common contacts between calcite and graphite in these complex mineral assemblages also suggests co-precipitation from fluid-deposition. These graphitic areas were often enveloped by hornblende and tended to form outgrowths into quartz (Fig. 6 a,

Petrography and crystal structure of Akilia graphite

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Fig. 8. TEM data for the FIB section of graphite associated with apatite shown in Fig. 5f: (a) bright field TEM image revealing a 500 nm graphite coating on apatite, (b) EDS hyperspectral maps of relevant elements, (c) high-resolution image (200,000) of the contact between graphite and apatite showing locations of d) EDS spot analyses of various materials within the section, and HAADF images of (e) the rounded carbonaceous structure and (f) graphite band with Fe-rich lamella after the section was re-thinned in the FIB-SEM.

c, and e) that probably developed during retrograde metamorphism. 4.2. Implications of the crystal structure of the Akilia graphite In carbonaceous materials, relative intensities of the characteristic G-band (graphite or ordered) at 1590 cm1 and D-band (disordered or amorphous) at 1340 cm1 are correlated to several factors including the presence of heteroatoms, the temperature of formation

(Busemann et al., 2007; Cody et al., 2008), and the maturity of carbonaceous material with respect to the metamorphic grade of the host rock (Tice et al., 2004; Beyssac et al., 2002; Wopenka and Pasteris, 1993). The Raman spectral characteristics of the Akilia graphite indicated a generally high degree of crystallinity with the occasional disordered sub-domains in coatings on apatite grains, which point to graphitization when the rock was exposed to high temperatures, probably during peak metamorphism (Griffin et al., 1980). While the possible effects of retrogression on the crystal structure of graphite have not been thoroughly

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Fig. 9. Synchrotron-based STXM data for graphite associated with apatite shown in Fig. 4a (a–c) and in Fig. 4b (d–e): STXM images of (a) the entire FIB section imaged at 284.0 eV and (b) the graphite coating and nearby carbonaceous structure imaged at various energies showing a bridge or connection (arrow) to the graphite coating at 285.0 eV, (c) C-XANES spectra at the K-edge for different spots shown in a and d, (d) 300 eV STXM images of microtomed graphite extracted with a micro-needle attached to a micro-manipulator and laid onto a Si–O coated Cu-TEM grid, and (e) the same microtomed graphite particle imaged at 401.0 eV.

documented, the degree of crystallization of carbonaceous material is generally considered to represent the peak conditions of the metamorphic history of the rock (Beyssac et al., 2002). Raman spectra for graphite associated with apatite and for graphite associated with calcite and/or ferruginous minerals are consistent with crystallization during granulite- and/or amphibolite-facies metamorphism. Raman image scans offered the visualization of sample heterogeneities and structurally unique sub-domains within graphite grains. Most notable of these are sub-micron regions producing an enhanced 2D Raman feature compared to bulk graphite in the Akilia Qp rock, which are similar to curled graphite structures identified by Raman spectroscopy and visualized by SEM in CV3 carbonaceous chondrite meteorites (Fries and Steele, 2008) and to similar structures from a graphite-bearing pegmatite (Jaszczak et al., 2007). The enhancement of the 2700 cm1 Raman feature compared to that of bulk graphite (Fig. 7d) has also been associated with spiral growth patterns in synthetic carbon (Dong et al., 2001). Curled graphite structures in the Akilia Qp rock are similar to graphite whiskers, which are needle-like structures consisting of rolled-up sheets of graphite that can be thermodynamically favorable at high temperatures (Sears, 1959), and such structures have been synthesized in high-temperature experiments in excess of 2000 K (Tan et al., 2001; Dong et al., 2001). These unusual graphite structures are attributed to spiral and conical growth patterns. Graphene monolayers can also exhibit this kind of strong 2D overtone and experiments have shown that partial hydrogenation of graphene sheets can decrease the 2D intensity relative to the G-band intensity and signif-

icantly increase the D-band intensity (Elias et al., 2009). The naturally-occurring curled graphite structures attributed to spiral growth and conical patterns have been reported from alkaline pegmatites in the Kola Peninsula are associated with various mineralogically complex nepheline rock-types of enigmatic origin (Jaszczak et al., 2007). Graphite in this pegmatite occurs as coatings on crystals of aegirine, strontian fluorapatite, natronite, K-feldspar, micas, and other minerals. Raman spectra of curled graphite structures in the Akilia Qp rock are similar to those observed in the Kola pegmatite. This comparison is also significant because the Kola occurrences represent one of very few reported occurrences of graphitic material associated with apatite. The curled graphite structures found in graphite coatings on apatite grains in the Akilia Qp rock may have formed when the rock was exposed to the highest temperature in its history (i.e. the granulite-facies metamorphic event around 3.65 Ga). However, the lower temperature limit at which curled graphite structures can form is unconstrained and it is therefore also possible that these structures formed during later, lower grade, metamorphic events at 2.7 (amphibolite-facies) or even at 1.7 Ga. Formation during lower temperature metamorphic events is supported by the discovery of similar curled graphite structures in graphite associated with apatite from a Paleoproterozoic hematite BIF from the Vichadero Formation in Uruguay, which was metamorphosed to the upper amphibolite-facies (Papineau et al., 2010). All curled graphite structures found in sample G91-26 were found in graphite coatings on apatite (Fig. 4) and the cause of this distribu-

Petrography and crystal structure of Akilia graphite

tion is unclear. In fact, the mechanism for the 2D enhancement is still not fully explained and all known syntheses of graphite whiskers have been performed by gas-phase deposition, which may not necessarily be an applicable mechanism to describe the formation of curled graphite structures in the Akilia Qp rock. Further work on defining the growth conditions of graphite whiskers is needed to fully assess the geological implications of the curled graphite structures in the Akilia Qp rock. The c-axis lattice parameter in graphitic material is also related to the metamorphic grade of the rock. With increasing metamorphic grade, the (0 0 0 2) interplanar spacing of graphite decreases as H and other atoms are expelled from interlayer sites between graphite sheets (Luque et al., 1998; Landis, 1971). We observed lattice spacings between 3.41 ˚ in graphite associated with apatite in the Akilia and 3.64 A ˚ for Qp rock, which is larger than the spacing of 3.35 A pure, well-crystallized graphite. This interlayer expansion is most likely due to interlayer atoms or molecules (Rietmeijer, 1991) and may be the source of non-carbon heteroatoms observed in Akilia Qp graphite by nanoSIMS (Papineau et al., 2010). Oberlin et al. (1980) observed that complete graphitization of carbonaceous material can be limited by its initial heteroatom content. Each of the Akilia samples analyzed by STXM indicated the presence of crystalline graphite, as evidenced by the intense 1s ? p* photoabsorption at 285.4 eV and the visible 1s ? r* exciton at 291.4 eV. The former transition is caused by highly-symmetric aromatic carbon (C@C) in graphite monolayer sheets, while the latter is characteristic of graphitic structural domains or highly conjugated aromatic domains (Cody et al. 2008). Furthermore, identifiable peaks between the p* and r* due to common organic functional groups such as carbonyl (C@O) or nitrile (C„N) were not observed. These observations are consistent with Raman spectroscopy indicating a generally high degree of crystallinity. 4.3. Temperature estimates for graphitization in the Aklilia Qp rock Because the degree of ordering in carbonaceous material is an indicator of metamorphic grade, Raman spectra can be used to estimate crystallization temperatures (Wopenka and Pasteris, 1993; Beyssac et al., 2002; Cody et al., 2008). The Raman characteristics of carbonaceous material have been calibrated against metasedimentary mineral assemblages to use as a geothermometer for peak metamorphism in the temperature range between 330 and 640 °C (Beyssac et al., 2002). The crystallization temperature was derived from peak areas of the G, D, and D2 bands in Raman spectra using the equation: T ð CÞ ¼ 445  ½Darea =ðDarea þ D2area þ Garea Þ þ 641 The estimated crystallization temperatures obtained for graphite associated with apatite in the Akilia Qp rock (excluding the curled graphite structures) was between 559 and 635 °C, while that for graphite in complex ferruginous mineral assemblages was between 439 and 464 °C (Table 2). These temperature estimates suggest that there may have

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been more than one graphitization event, and that graphite associated with apatite crystallized at higher temperatures than the rest of the carbonaceous material in the rock. However, it is possible that this temperature range for graphite associated with apatite under-estimated the actual crystallization temperature, since the equation of Beyssac et al. (2002) has not yet been calibrated for granulite-facies metasedimentary mineral assemblages. Another thermometer based on Raman spectral characteristics of acid-insoluble carbonaceous material from primitive chondrite meteorites has been calibrated against the 1s – r* exciton of C-XANES spectra (Cody et al., 2008). This geothermometer implicitly assumes that only temperature is responsible for changes in the structure and bonding of carbonaceous material, and therefore it over-estimates the crystallization temperature of carbonaceous material in terrestrial rocks, where pressure also has a large effect. In crustal environments, deviatoric stresses during tectonic shearing tend to accelerate graphitization. It has been observed that for increased deformation, temperatures and exposure times required for graphitization decrease (Nover et al., 2005). Two empirical equations relate Raman spectral parameters, one for the FWHM of the Raman D-band and a second for the FWHM of the G-band (not used here), to the temperature implied by the intensity of the 1s – r* exciton: T ð CÞ ¼ 899:9  3:0  DFWHM þ 1:4  103  DFWHM 2 This equation yielded estimated graphitization temperatures for the Akilia Qp rock between 764 and 830 °C, with an error of 120 °C (2r). These calculated temperatures are interpreted to reflect maximal values and we thus infer that graphitization in the Akilia Qp rock occurred between approximately 635 and 830 °C. This estimate is reconcilable with the well-known granulite-facies metamorphic event, estimated at more than 650 °C (Griffin et al., 1980), which likely represents the last re-equilibration temperature experienced by the rock during slow cooling from peak conditions (Pattison et al., 2003) at 3.65 Ga that affected the southern Itsaq Gneiss Complex. These conclusions do not preclude the possibility that there may have been graphitization events later in the history of the Akilia Qp rock and that one of these events may have produced the graphite associated with complex ferruginous mineral assemblage and/or calcite. 4.4. Non-biological pathways of carbon to form graphite Graphite can precipitate directly from CAOAH fluids during high-temperature fluid–rock interactions. Fluid-precipitated graphite typically consists of medium- to coarsegrained flakes or needles, although metamorphic graphite can have similar characteristics (Luque et al., 1998). Examples of fluid-deposited graphite have been reported in metasedimentary and meta-igneous rocks from New Hampshire (Rumble et al., 1986; Rumble and Hoering, 1986), South India (Farquhar et al., 1999), and other localities (Luque et al., 1998). Fluid-deposited graphite also generally has a high degree of crystallinity, due to its direct precipitation from a fluid rather than through a transfor-

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mation from sedimentary carbonaceous material, and only a few examples with minor degree of disorder have been reported from auriferous quartz veins and from laboratory experiments (Luque et al., 1998, 2009). Detailed petrographic analyses of fluid inclusions in metasomatically-altered Eoarchean meta-basalts from Isua have shown that microscopic graphite-carbonate mineral associations formed from H2O–CO2–(CH4)–NaCl fluids early in the history of the rock (Heijlen et al., 2006). It has recently been shown that fluid-deposited graphite can form at temperatures as low as 500 °C so that high temperatures are not necessarily required for this process (Luque et al., 2009). The co-existence of graphite associated with complex mineral assemblages (Fig. 6) that commonly include calcite, hornblende, pyrrhotite, chalcopyrite, pentlandite, and magnetite points to high-temperature precipitation from carbon-bearing fluids. The common (close) association of graphite with calcite further suggests that these two phases precipitated synchronously and out of redox equilibrium. If so, these occurrences may be most easily explained by a coprecipitation from high-temperature fluids, rather than the decarbonation of siderite (see below). Such mineral imbrications are often surrounded by a hornblende envelope that likely formed during the 2.7 Ga amphibolite-facies retrograde metamorphic event. The petrographic textures and mineral assemblages shown in Fig. 6 suggest a precipitation sequence starting with graphite and calcite, sometimes followed by magnetite, sulfides (chalcopyrite + pyrrhotite ± pentlandite), and more graphite, subsequently enveloped by a hornblende shell, and finally embedded in recrystallized quartz. This proposed sequence of precipitation/recrystallization from CO2–CH4 bearing high-temperature fluids may have happened as a result of highly variable oxygen fugacity in micro-environments of the rock during metamorphism (Pasteris, 1988). Sources of these carbon compounds in high-temperature fluids may have included mantle gases, adjacent younger rocks, and/or devolatilized biological carbonaceous material from the original Akilia depositional environment. Fluid inclusions, occasionally containing opaque material, are common in the Akilia Qp rock and these often occur near apatite-associated graphite (Fig. 2c–h). Depending on the composition of metamorphic fluids and, to a lesser extent, temperature and pressure, metasomatized apatites may contain monazite and xenotime inclusions (Harlov et al., 2002). For instance, monazite inclusions in fluorapatites from granulite-facies metapelites formed during the partial dissolution or dissolution–reprecipitation of fluorapatite when exposed to high-temperature metasomatism (Harlov et al., 2007). Since temperature estimates for graphite associated with apatite in the Akilia Qp rock are consistent with the granulite-facies metamorphism, the chemical composition of the apatite may have been altered during the high-grade metamorphic event at 3.65 Ga. Akilia apatites contain relatively high levels of rare Earth elements (up to 0.2 wt.%; Nishizawa et al., 2005), but they contain no REE-phosphate inclusions as may be expected from such metamorphic conditions. Chlorapatites in the Akilia Qp rock may have formed from the replacement of original fluorapatite, which can occur during anionic

exchanges in metasomatic fluids (Harlov et al., 2006). If these fluids were carbon-bearing and responsible for graphite precipitation onto apatite grains, the high-temperature estimates of graphitization indicate that this carbon source was 3.65 Ga or older. However, it remains unclear how these common, discrete, and specific mineral associations in the Akilia Qp rock could be explained by fluid-deposition alone. Graphite formation can also occur through the aqueous decarbonation of iron-bearing carbonate minerals at temperatures above approximately 300 °C, which can produce a variety of organic compounds. For example, experiments on siderite decomposition have shown that a wide variety of moderate molecular weight organics can be synthesized according to the overall reaction: 3FeCO3 þ wH2 O ! Fe3 O4 þ xCO2 þ yCO þ zH2 þ organic compounds Organic compounds synthesized from this reaction are dominated by alkylated and hydroxylated single ring aromatic compounds (McCollom, 2003). If appreciable amounts of these compounds were formed during the metamorphic decarbonation of iron-bearing carbonates in the Akilia Qp rock, then subsequent reactions under granulite-facies metamorphism might also yield graphite, but presumably in close association with magnetite. Natural samples, in which non-biological carbonaceous material is associated with magnetite and carbonate, have been reported from terrestrial and martian mantle rocks (Steele et al., 2007). Such carbonaceous material in igneous rocks has been attributed to both the decarbonation of siderite and primary formation due to the reduction of CO2 or biocarbonate (Steele et al., 2007). Decarbonation reactions have also been proposed as the mechanism by which graphite, sometimes associated with apatite, formed in Eoarchean metacarbonate rocks from Isua (van Zuilen et al., 2002, 2003, 2005). This suggestion was based on the common association of carbonate and graphite in these rocks and on carbon isotopic compositions. However, a plausible abiogenic mechanism has not been proposed for the association of graphite specifically with apatite because the decarbonation of siderite is not expected to form apatite as a by-product. The complex mineral assemblages of graphite + hornblende ± calcite ± magnetite ± sulfides (Fig. 6) may have formed from mineral reactions such as the decarbonation of siderite, but these assemblages do not always contain magnetite, which would be expected from such reactions. In summary, while the possibility of decarbonation reactions producing complex graphitic mineral assemblages cannot be rejected, the co-occurrence of sulfides and the common envelopes of retrograde hydrated hornblende (as detected by Raman) more strongly support a fluiddeposited origin for the graphite associated with calcite and/or ferruginous minerals in the Akilia Qp rock. Fischer–Tropsch type (FTT) synthesis of organic compounds is a process that involves the reaction of CO, CO2 or HCO3- and H2 on catalytic mineral surfaces (such as magnetite, hematite, olivine, Fe–Ni-, and Fe–Cr-minerals) and produces various short-chained aliphatic hydrocarbons,

Petrography and crystal structure of Akilia graphite

alcohols, and carboxylic acids (CH4, C2H6, C3H8, etc.) (McCollom and Seewald, 2007; Foustoukos and Seyfried, 2004; McCollom, 2003; and references therein). FTT synthesis produces saturated hydrocarbons, but it is conceivable that thermal cracking and dehydrogenation reactions under granulite-facies metamorphic conditions could yield graphite upon hydrogen loss. High-pressure and temperature experiments have shown that relatively simple polycyclic aromatic organic molecules can be transformed into graphite (Davydov et al., 2004). Carbonaceous solids in Paleoarchean black cherts have been proposed to have formed from FTT synthesis (Brasier et al., 2002, 2005; Lindsay et al., 2005), but the structure and composition of this carbon is also consistent with a biogenic origin (De Gregorio et al., 2009). While the graphite in the Akilia Qp rock could have formed from such abiogenic processes, the common association of graphite with apatite in this rock is not predicted by a FTT- synthetic origin. It is thus concluded that an important process of graphitization for the Akilia Qp rock was by fluid-deposition. 4.5. Biological carbon from depositional environment metamorphosed to graphite In younger rocks, the metamorphic conversion of biogenic kerogen into graphite is a common process of graphitization. Organic matter in younger BIFs is expected to be biogenic, although this origin may not hold in all cases. If microorganisms were involved in the formation of BIFs, carbonaceous remains would likely be present in these rocks. However, organic carbon in Precambrian BIFs is generally low in abundance, highly mature, and partly graphitized. Furthermore, organic biomarkers have not yet been reported from these rock types. Graphite crystallized from biological carbon has been reported to occur as disks from amphibolite-facies quartzites from the Archean-Paleoproterozoic Wutai Complex in North China (Schiffbauer et al., 2007) and these graphites exhibit Raman spectra with similar D- and G-band characteristics to graphite in the Akilia Qp rock. The occurrence of a 100 nm thick Fe–C band at the interface of a graphite coating on apatite and the enveloping quartz matrix (Fig. 8f) may have formed as an exosolution from Fe-rich carbonaceous material and appears analogous to the electron-dense material between graphite layers reported from graphite disks in the Wutai Complex (Schiffbauer et al., 2007). Carbon-XANES spectra for the Akilia graphite are also similar to those of graphitic carbon from Middle Triassic plant spores metamorphosed to the blueschist-facies (Bernard et al., 2007). The common association of 13C-depleted graphite with apatite in the Akilia Qp rock has been proposed as a biological signature (Mojzsis et al., 1996). This mineral association has been observed in a wide variety of sedimentary rocks and is thought to arise from the diagenetic maturation of organic matter. In the process of sedimentation and early diagenesis of BIFs, anaerobic heterotrophic oxidation (e.g. fermentation, sulfate reduction, denitrification, etc.) of organic carbon can result in the liberation of phosphate and bicarbonate in pore water solutions, which

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contributes to the formation of diagenetic carbonate and apatite (Berner, 1990). Carbonaceous material associated with phosphatic minerals is rather common in modern phosphorites, but is also seen in carbonaceous shales, cherts, sedimentary carbonates, and BIFs (Papineau et al., 2010). Thus, it is possible that apatite + graphite associations in the Akilia Qp rock represent biogenic carbonaceous material incorporated into phosphates during diagenesis. However, it must be emphasized that fluid-deposition mechanisms may also produce such mineral associations and that these have yet to be evaluated. There may very well be more than one generation of graphite in the Akilia Qp rock, especially in light of: (i) fluid inclusions (most of them now empty) near some apatite + graphite associations (Fig. 2); (ii) the presence of fluid-deposited graphite associated with hornblende ± calcite ± Fe–Cu–Ni sulfides ± magnetite (Fig. 6); and iii) the observed near match of the graphite (0 0 0 2) and apatite (2 0 1) lattice planes near the interface of the two minerals (Supplementary Electronic annex 4). In fact, these observations can be interpreted to mean that apatite + graphite coprecipitated from high-temperature fluids. During fluiddeposition, dissolved carbon can be derived from either biogenic or abiogenic sources, or a mixture of both. Metamorphic devolatilization reactions of biological organic matter produce variable proportions of CO2 and CH4 in fluids, depending on oxygen fugacity, which may subsequently mix with non-biological sources of these compounds and/ or re-precipitate as graphite. Based on our exhaustive study of these rocks, we conclude that the source of carbon for graphite associated with apatite and/or calcite included metasomatically remobilized CH4 and CO2 of exogenous and/or indigenous provenance. 5. SUMMARY AND IMPLICATIONS Detailed petrographic observations of Akilia Qp sample G91-26 confirmed that graphite is commonly associated with apatite in this rock. Raman spectroscopy indicated an overall high degree of crystallization of the graphite that is consistent with formation at relatively high temperatures, estimated at approximately between 635 and 830 °C, consistent with granulite-facies metamorphic conditions. Raman hyperspectral imaging has also revealed occasional sub-domains of curled graphite structures in apatite-associated graphite, which most likely formed during the granulite-facies metamorphic event at 3.65 Ga or perhaps during the amphibolite-facies event at 2.7 Ga. Thin sections extracted by dual beam FIB-SEM contained thin graphite coatings on apatite grains as opposed to inclusions sensu stricto as inferred from transmitted light microscopy and Raman spectroscopy. TEM and STXM analyses of these FIB sections offered unprecedented details of these mineral associations. The (0 0 0 2) lattice spacing of graphite coat˚ spacing of highly-orings were expanded from the 3.35 A dered pure graphite indicating the presence of interlayer hydrogen or other atoms or molecules. Graphite in the Akilia Qp rock was also found to occur in complex mineral assemblages of hornblende ± calcite ± sulfides ± magnetite that point to high-temperature precipitation from carbon-

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bearing fluids. Graphite in these mineral assemblages may represent another generation of graphitization that may have occurred during the later amphibolite-facies metamorphic event at 2.7 Ga. The crystallization of biogenic carbonaceous material during prograde metamorphism and fluid-deposition of graphite are the two most important processes that produce graphite in metasedimentary rocks. The presence of generally empty fluid inclusions near apatite + graphite associations may be consistent with fluid-deposition, although if this were the case, the carbon source of these fluids was 3.65 Ga or older. Abiogenic processes could have produced reduced carbon, either during sedimentation or after deposition, which may have eventually been converted into graphite. However, it remains unclear how apatite + graphite associations could be explained by fluid-deposition alone. Here, correlated micro-analyses of graphite in the Akilia Qp rock lead to the interpretation that at least some of it formed from fluid-deposition, although these new datasets do not exclude a biological origin of the carbon in graphite associated with apatite (Papineau et al., 2010). Sources of carbon for fluid-deposited graphite in the Akilia Qp rock included fluid-remobilized CH4 and CO2 from exogenous and/or indigenous provenance, which may have included devolatilized biogenic carbonaceous material from the original depositional environment. The combined use of multiple micro-analytical techniques to characterize ancient carbonaceous material is a significant step towards determining the origin(s) of the Akilia graphite. TEM and STXM analyses of FIB sections from targets identified by optical microscopy and mapped by Raman spectroscopy offer unprecedented details of this material. Raman spectroscopy of carbonaceous material from the Archean Apex chert (Schopf et al., 2002, 2005; Brasier et al., 2002, 2005), which has experienced relatively low-grade greenschist-facies metamorphism, was insufficient to determine its origin unambiguously (Pasteris and Wopenka, 2003). However, correlated analysis of microtomed sections of this carbonaceous material using TEM and STXM demonstrated that it was structurally similar to kerogen from the 1.878 Ga Gunflint chert, even if an FTT origin could not be completely ruled out (De Gregorio et al., 2006; De Gregorio and Sharp, 2009). In this study, correlated micro-analyses of graphite in the Akilia Qp rock led to the interpretation that at least some of the graphite formed from fluid-deposition, but that the origin of graphite associated with apatite is ambiguous. Rather, like the complicated and protracted metamorphic history of this 3.83 Ga Akilia Qp rock, it is more likely that a combination of pathways was responsible for the origin of carbon and the formation of graphite in the Akilia Qp rock. New micro-analytical techniques like those used in this work and tested on terrestrial rocks of known provenance constitute a useful baseline approach to pave the way in the search for evidence of life on other ancient planetary surfaces. ACKNOWLEDGEMENTS We would like to thank M. Obst and D. Kilcoyne for help and advice on STXM analyses. Discussions with D. Rumble, R. Hazen,

T. McCollom, and F. McCubbin, as well as reviews from two anonymous reviewers and T. Chacko helped to improve the manuscript. This work was supported by the NASA Exobiology and Evolutionary Biology Program (Grant # NNX08AO16G), the NASA Astrobiology Institute, and the Geophysical Laboratory of the Carnegie Institution of Washington. DP also acknowledges the Fond que´be´cois de la recherche sur la nature et les technologies for support. SJM acknowledges grants from the NASA exobiology (#NAG5-13497) and the NSF LExEn (EAR0228999) programs. The synchrotron-based STXM work described in this paper was partly performed at the Canadian Light Source, which is supported by NSERC, NRC, CIHR, and the University of Saskatchewan, and at the Advanced Light Source, which is supported by the Director of the Office of Energy Research, Office of Basic Energy Sciences, Materials Sciences Division of the US-DOE.

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