Journal of Petrology

Journal of Petrology Advance Access originally published online on November 19, 2004
Journal of Petrology 2005 46(2):291-318; doi:10.1093/petrology/egh075

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Journal of Petrology vol. 46 issue 2 © Oxford University Press 2004; all rights reserved

Assigning Dates to Thin Gneissic Veins in High-Grade Metamorphic Terranes: A Cautionary Tale from Akilia, Southwest Greenland

MARTIN J. WHITEHOUSE^1,* and BALZ S. KAMBER²

¹ SWEDISH MUSEUM OF NATURAL HISTORY, BOX 50007, SE-104 05 STOCKHOLM, SWEDEN
² ADVANCED CENTRE FOR QUEENSLAND UNIVERSITY ISOTOPE RESEARCH EXCELLENCE (ACQUIRE), UNIVERSITY OF QUEENSLAND, ST LUCIA, QLD 4072, AUSTRALIA

RECEIVED MAY 7, 2004; ACCEPTED SEPTEMBER 6, 2004

	ABSTRACT

TOP ABSTRACT INTRODUCTION REVIEW OF AKILIA GEOLOGY... SAMPLING ANALYTICAL METHODS RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
A granodiorite from Akilia, southwest Greenland, previously suggested to date putative life-bearing rocks to 3·84 Ga, is re-investigated using whole-rock major and trace-element geochemistry, and detailed cathodoluminescence image-guided secondary ion mass spectrometer analyses of zircon U–Th–Pb and rare earth elements. Complex zircon internal structure reveals three episodes of zircon growth and/or recrystallization dated to c. 3·84 Ga, 3·62 Ga and 2·71 Ga. Rare earth element abundances imply a significant role for garnet in zircon generation at 3·62 Ga and 2·71 Ga. The 3·62 Ga event is interpreted as partial melting of a c. 3·84 Ga grey gneiss precursor at granulite facies with residual garnet. Migration of this 3·62 Ga magma (or melt–crystal mush) away from the melt source places a maximum age limit on any intrusive relationship. These early Archaean relationships have been complicated further by isotopic reworking in the 2·71 Ga event, which could have included a further episode of partial melting. This study highlights a general problem associated with dating thin gneissic veins in polyphase metamorphic terranes, where field relationships may be ambiguous and zircon inheritance can be expected.

KEY WORDS: Archaean; geochronology; Greenland; secondary ion mass spectrometry; zircon

	INTRODUCTION

TOP ABSTRACT INTRODUCTION REVIEW OF AKILIA GEOLOGY... SAMPLING ANALYTICAL METHODS RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
A major source of controversy in the debate about the origin and evolution of life on Earth is the assignment of both relative and absolute ages to key lithologies. Indeed, the significance of actual evidence for biogenicity or for key steps of biological evolution is impossible to evaluate in cases where no consensus can be reached about relative and absolute age. Nowhere is this problem more pronounced than on the island of Akilia, southwest Greenland, where evidence for Earth's earliest life has been claimed on the basis of isotopically light carbon, occurring as graphite inclusions hosted by apatite, in a rock originally interpreted as a banded iron formation (Mojzsis et al., 1996). Since this claim was first published, several lines of evidence have emerged that question both the biogenicity and antiquity of the Akilia example, to the extent that most specialists now prefer to focus the discussion about earliest life on the much better preserved and unambiguous volcano-sedimentary sequences in the (near-) contemporaneous Isua Greenstone Belt and the somewhat younger Warrawoona Group in the East Pilbara Craton. Despite this sobering outcome of the Akilia debate, many lessons can still be learnt, e.g. regarding establishment of relative and absolute age relationships of key lithologies.

All lithologies on Akilia have experienced two major episodes of metamorphism at 3·60–3·65 and 2·71 Ga, reaching granulite and upper amphibolite facies respectively, as well as later thermal and minor structural overprints. In cases where age relationships are determined with granitoid rocks, such high-T metamorphism and high-strain deformation conspire to inextricably link relative and absolute ages. First, quartz and feldspar, the main constituents of granitoids, deform plastically under such conditions (e.g. Pryer, 1993) and, in the presence of an aqueous fluid, have potential to melt at temperatures exceeding 670°C (e.g. Clemens & Vielzeuf, 1987). Once an anatectic melt is present, the crust is further weakened (Hollister & Crawford, 1986), which facilitates and enhances deformation. As a result, partially molten rocks can easily lose their original relative field context and assume discordant relationships that preferentially reflect the latest preserved episode of deformation rather than original relationships. Second, although the U–Pb zircon geochronometer is very resilient, Pb-loss, recrystallization or even new growth occur typically under amphibolite- and granulite-facies metamorphism (e.g. Fraser et al., 1997; Vavra et al., 1999; Hoskin & Black, 2000; Hoskin & Schaltegger, 2003; Whitehouse & Platt, 2003). Memory loss and recrystallization are certainly aided by the presence of an aqueous fluid (Villa, 1998; Tomaschek et al., 2003) but crystallization of overgrowths on pre-existing grains or even entirely new grains is more likely in a partial melt (Roberts & Finger, 1997).

The combination of disruption of original relative age relationships with variable extent of zircon (re-)crystallization under high-grade tectonometamorphism inextricably links the tasks of establishing the relative and absolute ages of highly deformed and migmatized lithologies. In the Akilia example, the validity of cross-cutting field relationships is ambiguous at best. Here, we present new high-spatial resolution (c. 10–30 µm) U–Th–Pb isotopic and rare earth element (REE) abundance data for carefully characterized growth phases in polyphase zircon crystals from a granitoid whose intrusion age has been claimed, by some, to provide a minimum age for life on Earth. The complexity of field, structural, textural, chemical and isotopic relationships of these outcrops reflects the fact that the southwestern tip of the island is essentially a highly metamorphosed shear zone. The key to finding more convincing evidence for the establishment of life is to concentrate on low-strain outcrops, where unequivocal relative field relationships can be established and where absolute dates and isotopic integrity are preserved to the extent that the null hypothesis can be applied.

	REVIEW OF AKILIA GEOLOGY AND GEOCHRONOLOGY

TOP ABSTRACT INTRODUCTION REVIEW OF AKILIA GEOLOGY... SAMPLING ANALYTICAL METHODS RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
The lithologies on southwest Akilia that are key to inferences for early terrestrial life occur as enclaves of mafic–ultramafic rocks and putative metasediments [collectively defined as the Akilia association by McGregor & Mason (1977)] within the regional early Archean Amîtsoq gneisses. The rock types in these enclaves are not amenable to precise dating with methods available at present; hence, the debate about their absolute age has revolved around the nature of their relative age relationship to their host gneisses and, on Akilia itself, focuses on a single, small outcrop where highly strained granitoid gneisses occur in contact with the putative life-bearing enclave. Despite their obviously complex structural context, these granitoid gneisses have, none the less, repeatedly been interpreted to cross-cut and therefore provide a minimum age for the mafic–ultramafic enclave (Nutman et al., 1997, 2000, 2002; Mojzsis & Harrison, 2000, 2002). In their comprehensive treatment of field relationships on southwest Akilia, Myers & Crowley (2000) concluded that the high-strain nature of the outcrop made it impossible to determine with any degree of confidence the relative age relationships between gneissic protoliths. Whitehouse & Fedo (2003) have presented further structural data from this specific outcrop and conclude that, with regard to relative dating constraints, the critical contact is not only tectonized at the point at which discordance has previously been claimed, but also documents a truncation of structures in the gneiss sheet by the mafic–ultramafic enclave, yielding the complete opposite of previously claimed relative age relationships. Based on these studies, there is clearly little point in obtaining absolute age data from the gneisses in order to constrain the age of the mafic–ultramafic enclave. Indeed, more recent studies of the putative metasediments (Fedo & Whitehouse, 2002; Bolhar et al., 2004) that have been claimed to be dated by the supposedly discordant gneiss sheet leave, in our view at least, little doubt that the original evidence for traces of life itself must be in error. This view is further reinforced by a reported absence of apatite-hosted graphite particles in a reinvestigation (Lepland et al., 2005) of the original sample. None the less, for scientists who remain persuaded by the original geological and geochemical evidence (Mojzsis et al., 1996), it is important to test whether a coherent absolute age interpretation of the quartz diorite sheet is attainable. As a prelude to our own study, we review below all previous geochronological data obtained from this sheet.

The only data of relevance are for rocks from the contact between the Akilia association enclave and Amîtsoq gneisses (Fig. 1) on the northern side of the small peninsula that forms the southwestern tip of Akilia. This so-called ‘northern contact’ is where Nutman et al. (1997) collected their sample, G93-05, which they described as a ‘quartz-diorite sheet’. The rock apparently yielded plenty of zircon (Nutman et al., 1997, p. 2478) ‘which range from approximately 100 to 300 µm in length with most of them prismatic and the remainder being stubby-prismatic to ovoid-equant in habit’. These workers further claimed (p. 2478) that ‘structurally older inherited cores were not discerned in the photomicrographs of the analysed zircons’. As will become apparent when discussing our new data, it is important to recall Nutman et al.'s (1997) statement (p. 2478) that ‘in selecting grains of the main population for analysis, the clearest (generally lowest U) grains were predominantly selected’. The clear, low-U grains analysed by Nutman et al. (1997) define a statistically relevant population with an average age of 3865 ± 11 Ma (2), whereas the deeper-coloured, mostly ovoid grains yielded ages between 3630 and 3530 Ma. On the basis of the their data and observations regarding zircon morphology, size and colour, Nutman et al. (1997) concluded that the igneous protolith of the quartz-diorite gneiss sheet intruded before 3850 Ma.

View larger version (55K): [in this window] [in a new window]   Fig. 1. Location maps of the Archaean craton in southern Greenland and the location of Akilia in the Godthåbsfjord. Detailed map of southwest Akilia (main panel), modified after Myers & Crowley (2000), shows the location of sample SM/GR/97/8 (GPS coordinates 63°55·853'N, 51°41·018'W, WGS84 datum) close to the eastern margin of the folded Akilia mafic–ultramafic enclave.

 
Whitehouse et al. (1999) also published ion-probe U–Pb zircon data for a so-called quartz-diorite gneiss from the same general outcrop (their sample SM/GR/97/7). They presented cathodoluminescence (CL) images of the prismatic, elongate zircons selected for analysis, which showed oscillatory banding and a paucity of core structures. Ion-probe data for SM/GR/97/7 zircons did not directly correspond to Nutman et al.'s (1997) results, either in age distribution (younger) or U-content (higher). Whitehouse et al. (1999) concluded based on their own data and CL imaging that the bulk of the zircon consisted of igneous growth at 3650 Ma, with minor inherited cores in excess of 3·8 Ga. Not surprisingly, Nutman et al. (2001) commented on the new results and inferred that Whitehouse et al. (1999) had either mislabelled or otherwise confused a sample, as the differences between data for G93-05 and SM/GR/97/7 appeared too substantial to permit a common origin of both samples. In their response, Whitehouse et al. (2001) confirmed that their sampling site of SM/GR/97/7 was no more than 25 cm from where Nutman et al. (1997) claimed to have taken G93-05. Whitehouse et al. (2001) acknowledged the substantial difference in data for zircon extracted from samples in such close proximity but reiterated their original explanation that this could have been because of a bias in selecting grains for mounting and/or that there could be true sample heterogeneity on a dm scale—something that is indeed evident from the now numerous detailed maps and photographs of the outcrop (Myers & Crowley, 2000; Nutman et al., 2000, 2002; Whitehouse & Fedo, 2003).

Subsequently, Nutman et al. (2000) published retrospective CL images for zircons of G93-05, which confirmed the paucity of the 3·6–3·7 Ga CL-zoned thick rims that Whitehouse et al. (1999) described from their sample SM/GR/97/7. Nutman et al. (2000) also provided new U–Pb ion-probe dates for additional zircon grains from G93-05. Their new data confirmed that among the grains selected by Nutman et al. (1997), biased to large, low-U, clear, prismatic material, the pre-3·8 Ga oscillatory grains clearly dominate. The only exception is their grain 16 [shown in fig. 10 of Nutman et al. (2000)], which is a strongly zoned, but equant grain, that is clearly 3·65 Ga old.

The most recently published dataset (Mojzsis & Harrison, 2002) for zircon from another sample (GR9716) from this locality, claimed to be equivalent to G93-05, has added considerably to the complexity. Unfortunately, Mojzsis & Harrison (2002) did not provide CL images of their zircon. Instead, they obtained two datasets using two different experimental approaches. One dataset was obtained by depth profiling into the centre of a single, clear (presumably low-U) euhedral, prismatic zircon (grain GR9716_69) similar to the majority of zircons analysed by Nutman et al. (1997, 2000) from their sample G93-05. The profiling technique yielded clear evidence for this single grain of a 3828 ± 8 Ma core, mantled by a c. 3·6–3·65 Ga inner rim, which itself is mantled by a c. 2·7 Ga outer rim. This observation is entirely consistent with the general picture commonly obtained from tonalite–trondhjemite–granodiorite (TTG) suite gneisses in the area (e.g. Whitehouse et al., 1999; Nutman et al., 2002) when dates are obtained from suites of handpicked zircon grains selected for CL-image-guided U–Pb ion-probe dating. The second dataset presented by Mojzsis & Harrison (2002), however, is potentially more insightful than the depth profiling because it was obtained in situ in petrographic thin sections. Of the 71 grains analysed, only a single grain, embedded in biotite, yielded an age in excess of 3·8 Ga. This grain (g1-1) yielded a near-concordant ²⁰⁷Pb/²⁰⁶Pb age of 4079 ± 18 Ma and was interpreted by Mojzsis & Harrison (2002) as an inherited xenocryst at a relatively late stage of protolith emplacement. This is by far the oldest zircon yet reported from the North Atlantic Craton and has no correspondence to any other previously studied sample. However, the important observation from the in situ study is that almost all the grains analysed yielded ages between 3·5 and 3·7 Ga. This prompted Mojzsis & Harrison (2002) to propose that in a 3·6 Ga migmatization event, a significant fraction ‘(50%?)’ of the protolith zircon dissolved and re-precipitated via partial melting, dissolution–reprecipitation or Ostwald ripening. However, provided that the dated zircons in this section are representative of the rock's entire population (i.e. only one grain in 71 analysed is older than 3·8 Ga), the estimate of the fraction of reworked zircons ought to be moved upward from ‘(50%?)’ to 98%. Although the in situ method has obvious advantages for preserving the petrographic enviroment of the analysed zircon, it should be realized that unlike separated zircon grain mounts, where polishing is optimized to reveal as much internal structure as possible, the zircon surface exposed in a thin section is essentially random. Without additional control of CL (or BSE) imaging, application of this technique to potentially complex zircons will inevitably yield ambiguous results owing to mixing of different growth phases. Indeed, a plot of Th/U against apparent age [fig. 7 of Mojzsis & Harrison (2002)] shows poor correspondence between in situ analyses and results of an earlier (Mojzsis & Harrison, 2000) grain mount study on the very same sample. The apparent paucity of >3·8 Ga zircon growth zones in the grains exposed in the thin section studied by Mojzsis & Harrison (2002) could thus partly stem from overlapping the primary ion beam onto ubiquitous younger mantles around old cores.

When the entire zircon evidence from this locality is considered, two obvious implications for further studies (such as ours) emerge. First, there appears to be very significant sample heterogeneity on the dm-scale as far as the distribution of zircon with different age, structure and chemistry is concerned. Only this can explain why Nutman et al. (1997) obtained a maximum age of 3865 ± 11 Ma [later revised to 3841 ± 6 Ma by Nutman et al. (2000)], and Mojzsis & Harrison (2002) determined (from a single grain) a slightly younger age of 3828 ± 8 Ma, whereas Whitehouse et al. (1999) found only a single >3·8 Ga age from the core of a c. 3·65 Ga grain. As it is physically impossible to sample exactly the same material as previous workers (unless a concerted effort is made to obtain a single large sample that is later split into aliquots and distributed to different laboratories), we have reprocessed one of our samples (SM/GR/97/8, taken immediately adjacent to SM/GR/97/7) which we now believe corresponds to the original G93-05 [see discussion also by Whitehouse & Fedo (2003)] in an effort to obtain a higher yield of zircon for a new imaging and dating session. The second implication is that irrespective of the heterogeneity in zircon distribution, it is obvious that the selection of rare, clear (low-U), prismatic grains for ion-probe analysis very strongly biases to the oldest grains. The picture gained from such grains is not at all representative of the entire zircon population. Although this bias may be logical and intentional when directed towards estimating the potential age of a pre-metamorphic protolith, it can be misleading with regard to the geological history experienced by a sample. In the case of the much-discussed Akilia locality, the question is clearly that of how old the relative field relationships are (which some interpret as igneous cross-cutting), and it is, therefore, critically important to test whether these rocks underwent significant, pervasive partial melting that would have obliterated older structural relationships pertaining to their original, magmatic emplacement. For this reason, we have tried to select the complete variety of zircon present in these rocks, and to analyse the various growth and recrystallization domains observed in the CL images. Furthermore, building on our previous work (Whitehouse & Kamber, 2002, 2003), we have obtained REE compositional data for the dated zircon spots.

	SAMPLING

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The sample studied here, SM/GR/97/8, was collected from the northern contact locality (Fig. 1) by S. Moorbath in the company of the late V. R. McGregor and S. J. Mojzsis. It is closely related spatially to sample SM/GR/97/7 of Whitehouse et al. (1999), coming from <25 cm west across regional strike of that sample. Whitehouse & Fedo (2003) noted that the latter sample preserves gneissic banding in thin section, whereas SM/GR/97/8 is essentially homogeneous (Fig. 2), with equigranular quartz and feldspar mosaic texture implying largely static, post-deformational recrystallization. This led them to suggest that SM/GR/97/8 is the most likely equivalent of the so-called ‘homogeneous tonalite’ sheet worked on by Nutman et al. (1997, 2000) and Mojzsis & Harrison (2000, 2002), with SM/GR/97/7 representing the banded gneiss immediately adjacent to the east, which, they noted, in places, appears to be cross-cut by the homogeneous sheet [see fig. 4d of Whitehouse & Fedo (2003)].

View larger version (132K): [in this window] [in a new window]   Fig. 2. Thin section images comparing texture of (a) granodiorite, SM/GR/97/8 (scale bar represents 4 mm) and (b) adjacent gneiss, SM/GR/97/7 (scale bar represents 2 mm). Note strong millimetre-scale banding in (b) defined primarily by hornblende. (c) Photomicrograph (plane-polarized light) of SM/GR/97/8 showing equigranular, mosaic recrytstallization texture of quartz (qtz) and feldspar (pl), together with hornblende (hbl) and biotite (bi) laths, the latter defining a slight foliation; field of view is c. 2 mm x 3 mm.

 

	ANALYTICAL METHODS

TOP ABSTRACT INTRODUCTION REVIEW OF AKILIA GEOLOGY... SAMPLING ANALYTICAL METHODS RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
High spatial resolution U–Th–Pb and REE data were generated using a Cameca IMS1270 large-format ion microprobe at the Nordsim facility, Swedish Museum of Natural History. Detailed analytical methods have been described previously for U–Th–Pb (Whitehouse et al., 1997, 1999) and REE (Whitehouse & Kamber, 2003; Whitehouse & Platt, 2003). In all cases, a defocused O₂-primary beam was used to project the image of either a 100 µm or a 50 µm aperture onto the sample, generating elliptical, flat-bottomed craters of nominal c. 30 µm and c. 15 µm (long axis), respectively. For some analyses, the spatial resolution was further enhanced by use of the field aperture in the secondary ion optics, restricting the actual analysed region admitted to the mass spectrometer to a c. 11 µm square on the illuminated object, regardless of the size of primary probe. Complete U–Th–Pb analyses at a mass resolution (M/M) of c. 5000 were performed using a peak switching routine, with a single ion-counting electron multiplier (EM) as the detection device. An energy window of 60 eV was used throughout, with energy adjustments made using the ⁹⁰Zr₂¹⁶O peak. Precise mass calibration was maintained by using an automatic routine in the Cameca CIPS software to scan over large peaks and extrapolate the mass to B-field curve for peaks between these reference points (e.g. Pb-isotopes were calibrated by centring the ⁹⁴Zr₂¹⁶O peak at nominal mass 204 and the ¹⁷⁷Hf¹⁶O₂ peak at nominal mass 209). Pb/U ratios, elemental concentrations and Th/U ratios were calibrated relative to the Geostandards zircon 91500 reference, which has an age of 1065 Ma (Wiedenbeck et al., 1995). In some cases, particularly the highest spatial resolution analyses of very small zircon domains, only Pb-isotopic ratios were measured using either a short mono-collection peak-switching routine, or a multi-collection routine using four ion-counting EMs positioned at the appropriate mass spacing and operating at a mass resolution of c. 8000 (determined by the fixed exit slit width of the multicollector system). For the multicollector measurements, the relative efficiencies of the EMs were first matched as closely as possible using pulse height analysis to set appropriate high voltage and thresholds, and then calibrated by direct measurement on the NIST SRM610 standard, assuming the composition determined by TIMS (Belshaw et al., 1994). Decay constants follow the recommendations of Steiger & Jäger (1977).

Simultaneous high mass resolution ion counting was also used for most of the analyses of Light REE (LREE) (La–Eu) coupled with low mass resolution, energy filtered mono-collection measurements of the middle to heavy REE (MREE to HREE) following the method detailed by Whitehouse & Kamber (2003) and Whitehouse (2004). For a few analyses, selected high-abundance HREE (i.e. Sm, Gd, Dy, Er, Yb) were also analysed at high mass resolution using simultaneous ion counting, with a single peak switch to measure ⁹²Zr²⁸Si¹⁶O₂ at nominal mass 152 as a matrix reference peak. REE abundances were normalized relative to NIST SRM610 [working values of Pearce et al. (1996)] with the exception of the multi-collector HREE measurements where the Geostandards 91500 zircon was used, assuming the concentrations determined by Whitehouse & Platt [2003; these were themselves calibrated against NIST SRM610 in this earlier study, ensuring compatibility of the datasets generated by the different methods; furthermore, these concentrations are in good agreement with the ‘working values’ of Wiedenbeck et al. (2004)]. Tables 2 and 3 present the U–Th–Pb and REE analytical data, respectively, and also document the spatial resolution and method used for each analysis.

Cathodoluminescence images of zircon were obtained using a Hitachi S4300 scanning electron microscope at the Swedish Museum of Natural History.

Whole-rock major and trace-element analyses of SM/GR/97/8 were obtained from an aliquot that was milled in tungsten carbide at Oxford University. No data for W and Ta are reported because of contamination from the mill. Major element compositions were determined at the Department of Earth Sciences (UQ) by ICP–OES from a solution obtained by dissolving a fused bead. Results are reported in Table 1. Trace-element analysis was performed at the ACQUIRE laboratory (UQ) by ICP–MS from a solution obtained by digesting the sample in a steel-jacketed PTFE (Teflon), vessel. Data are reported in Table 1. Analytical details were the same as those reported by Kamber et al. (2003).

View this table: [in this window] [in a new window]   Table 1: Major and selected trace-element data from Akilia samples

 

	RESULTS

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Whole-rock major and trace-element geochemistry
At least a part of the confusion regarding equivalence of samples from this outcrop stems from the disparate petrographic names assigned by various workers. The original sample, G93-05, was called a ‘quartz-diorite’ by Nutman et al. (1997), a term also adopted by Whitehouse et al. (1999) for their sample, SM/GR/97/7, but was later referred to as a homogeneous ‘tonalite’ by Nutman et al. (2000). Mojzsis & Harrison (2002) referred to their supposedly equivalent sample GR9716 also as tonalite. Major and selected trace-element (Zr, Pb, U, Th) compositions of these samples are listed in Table 1 alongside that of our sample, SM/GR/97/8.

We note the close correspondence of all analyses from these samples, in accord with our belief that they do indeed originate from the same, homogeneous part of the gneiss sheet. The normative mineralogy calculated from bulk-rock composition also indicates that all three samples are in fact granodiorites according to the widely accepted IUGS classification of plutonic rocks (Streckeisen, 1976; le Maitre, 2002), richer in quartz (c. 27%) than quartz-diorite (5–20%) and lower in plagioclase as a proportion of total feldspar (c. 85%) than tonalite (90–100%). On this basis, we use granodiorite throughout this paper. We note, however, that although quartz-diorite is definitely an inappropriate name, use of other plutonic rock-classification schemes adds ambiguity to the distinction between granodiorite and tonalite. In an ab–an–or diagram (Barker, 1979), the percentage of normative orthoclase (c. 14%) means that all three analyses plot within the tonalite field. The use of normative mineralogy in classification schemes such as QAP or ab–an–or that are intended principally for modal mineralogy can introduce additional ambiguity when dealing with hydrated rocks (SM/GR/97/8 contains abundant modal biotite) because of the tendency of norms to assign K entirely to orthoclase and to consider only orthoclase as an alkali feldspar. In the only commonly used classification scheme that utilizes elemental abundances instead of normative or modal mineralogy, the R₁–R₂ cation proportion scheme of de la Roche et al. (1980), these analyses fall on the boundary between the granodiorite and tonalite field.

This confusion over precise petrographic name is far from trivial when set against the backdrop of arguments about zircon saturation and the ability of melts to retain inherited zircon that has long been a feature of the Akilia geochronology debate. As noted recently by Hanchar & Watson (2003), the actual Zr-concentration and cation ratio (Na + K + 2Ca)/(AlSi) value (M) have only a small influence on zircon saturation temperature (± a few tens of °C), the most significance factor being the assumed temperature of crystallization. The higher magmatic temperature expected for a tonalite implies a greater ability to dissolve xenocrystic zircon compared with a lower-temperature granodiorite, although, as we shall demonstrate below, several additional factors need to be taken into account in discussing zircon solubility, not the least of which is the presence of water, which will significantly lower the solidus of any granitic composition.

With regard to incompatible trace elements, sample SM/GR/97/8 has general characteristics similar to those of arc (sensu lato) magmas (Fig. 3a). The key features are over-enrichment (relative to similarly compatible elements) in Cs, Pb, Be, Li and relative depletion in Nb (note that the reported value is an overestimate as a result of likely contamination from the tungsten–carbide mill). Although these features do not necessarily imply an origin of this chemistry from supra-subduction zone melting, we emphasize that enrichment in Cs, Pb, Be and Li is a very clear indication that source melting was triggered by aqueous fluids. Nutman et al. (1999) and Kamber et al. (2002) have reported similar incompatible trace-element patterns for a variety of early Archaean TTGs from southwest Greenland. The former researchers commented on the fact that trace-element systematics such as those of sample SM/GR/97/8 cannot be generated without preferential element enrichment from aqueous fluids. The latter workers provided additional evidence for the role of aqueous fluids in generating the earliest preserved continental crust. In the context of the Akilia debate, it is very significant that trace-element distribution strongly argues in favour of hydrous rather than anhydrous melting because previous consideration of zircon solubility (Nutman et al., 1997; Mojzsis & Harrison, 2002) has assumed that the igneous precursor of the studied rock crystallized from a ‘dry’ melt which, together with inappropriate petrographic identification discussed above, implies a much higher magmatic temperature and, hence, ability to dissolve inherited zircon. This point will be discussed in more detail after presentation of the zircon data.

View larger version (28K): [in this window] [in a new window]   Fig. 3. (a) N-MORB normalized (Sun & McDonough, 1989) trace-element pattern of sample SM/GR/97/8 (solid circles) compared with average continental crust estimate (open circles) of Rudnick & Fountain (1995). Beryllium concentration estimate for N-MORB explained by Kamber et al. (2002). Note coherence in preferential Cs, Pb, Be and Li enrichment. (b) Chondrite-normalized La–Yb vs Yb diagram after Martin (1986). Note that there is considerable variation in La–Yb for >3·8 Ga tonalites and quartz-diorites shown as open circles [data from Nutman et al. (1999) and Kamber et al. (2000)] but with one exception, they all plot into the field of Archaean TTG. The more potassic 3·65 Ga Amîtsoq gneisses [shown as solid circles; data from Kamber et al. (2002)] show even more variability in LREE/HREE fractionation, with several samples plotting outside the Archaean TTG field. Sample SM/GR/97/8 (large open cross) plots at the lowermost limit of Archaean TTG.

 
A final aspect of the whole-rock geochemistry that is of relevance to the present study is the observation that the studied granodiorite does not show the strong LREE over HREE enrichment that is so characteristic of typical Archaean TTGs. In fact, the chondrite-normalized La/Yb of our sample is only 9·0, whereas the average of reported >3·8 Ga tonalite and quartz-diorite data from south of the Isua Greenstone Belt is 18—almost twice as fractionated as sample SM/GR/98/8. For this reason, this sample only just plots into the Archaean TTG field (Fig 3b) defined by Martin (1986). The more potassic 3·65 Ga Amîtsoq gneisses, reported by Kamber et al. (2002), show a bimodal distribution of chondrite-normalized La/Yb: one group with very high values of c. 50, the other with much lower values of 6–17. This implies that at least in terms of REE systematics, the studied sample from Akilia is rather atypical for an Archaean TTG but akin to more potassic 3·65 Ga granitoids in the area. The origin of extreme HREE depletion in the more typical TTG may be a function of fractional crystallization processes rather than direct source processes as suggested by Kleinhanns et al. (2003), possibly due in part to differing H₂O-contents (Muentener et al., 2001).

Zircon populations in SM/GR/97/8
CL imaging reveals complex, polyphase internal structures in most zircon grains from sample SM/GR/97/8 (Fig. 4). On the basis of this imaging, we identify three principal types of zircon growth, commonly more than one of which is represented within a single grain. Group 1a comprises oscillatory-zoned cores occurring in euhedral to slightly rounded prismatic grains with length/width ratio of c. 3–4. The oscillatory zoned part of the grain is generally CL-bright (subgroup 1a) and its zoning is commonly sharply truncated, in some cases (e.g. grain 2–34) showing evidence for fracturing. Several cores (subgroup 1b) exhibit considerably darker CL and less distinct zoning (e.g. grains 1–4, 1–9 and 1–16). Group 2 ‘soccer ball’ zircon grains are generally equant and multifaceted, showing very low cathodoluminescence, although vaguely zoned internal structures are occasionally visible (e.g. grain 2–13, Fig. 4). Group 3 zircon grains comprise CL-dark, sometimes zoned, rims or embayments that occur on both group 1a zircon grains (in which case, they are classified as subgroup 3a) and group 1b and 2 zircon grains (classified as subgroup 3b). This type of zircon varies from substantial tip overgrowths of prismatic grains (e.g. grain 2–4, Fig. 4) to thin rims developed on typical group 2 grains (e.g. grain 2–14, Fig. 4). A fourth group of zircon is defined as outer rims that are found on grains with two earlier phases of growth.

View larger version (258K): [in this window] [in a new window]   Fig. 4. Cathodoluminesence images of all zircon grains investigated in this study. Labelling nomenclature 1–12b, for example, indicates sample mount 1, grain number 12, analysis b. Ellipses show location of analysed region (U–Th–Pb or Pb-isotopes for geochronology and REE determinations). Where spatial resolution was further enhanced using field aperture blanking, this is indicated by square outline, which represents the region of the grain that was actually analysed; in this case, the primary beam size is indicated by a dashed ellipse. Note that dark areas in some of the analysed regions correspond to residual gold coating from later analytical sessions that was not entirely removed prior to post-analytical CL imaging (e.g. grain 2–2, 2–8, 2–10).

 
Zircon U–Th–Pb geochronology
SIMS U–Th–Pb analytical data for zircons from SM/GR/97/8 are presented in Table 2 and in a plot of ²⁰⁷Pb/²⁰⁶Pb dates in increasing order (Fig. 5). These are discussed below in the context of the grouping of zircon growth phases based on CL characteristics.

View larger version (25K): [in this window] [in a new window]   Fig. 5. Cumulative plot of 207Pb/206Pb ages, arranged in order of consecutive age and differentiated into different zircon groups, combined with a probability plot. Flat regions in the cumulative plot, corresponding to significant peaks in the probability plot, are interpreted to result from specific geological events for which weighted average ages have been calculated, in some cases including analyses from zircon belonging to different groups. Sloping regions in the cumulative plot and smaller peaks to the low-age side of the main peaks in the probability plot are interpreted to result from non-recent Pb-loss.

 

View this table: [in this window] [in a new window]   Table 2: SIMS U–Th–Pb data from Akilia sample SM/GR/97/8

 
Group 1 zircon cores exhibit a wide but bimodal range of ²⁰⁷Pb/²⁰⁶Pb dates from c. 2·7 to >3·8 Ga. The CL bright, oscillatory-zoned majority of this group (subgroup 1a) show a more restricted range of ²⁰⁷Pb/²⁰⁶Pb dates from c. 3·57 to 3·88 Ga. These analyses are concordant to slightly discordant with low-U concentrations (c. 70–220 ppm) and moderate Th/U ratios (0·14–0·70). The similar CL characteristics and Th/U systematics suggest that these grains could belong to a single igneous protolith with an age in excess of 3·8 Ga which has experienced variable amounts of ancient Pb-loss, probably during the high-grade events recorded in the Godthåbsfjord at c. 3·65 and c. 2·7 Ga, resulting in a spread of ages along concordia. In this context, the oldest coherent group of ²⁰⁷Pb/²⁰⁶Pb dates that define a prominent plateau on the cumulative age diagram (Fig. 5) yields a weighted average of 3842 ± 5 Ma (n = 7, MSWD = 0·62, 95% confidence). This interpretation is in accord with previous interpretations of similarly spread dates from other gneisses in this region (e.g. Nutman et al., 1996, 2000; Whitehouse et al., 1999), but the possibility that younger magmatic zircon is represented in this population cannot be entirely ruled out. A single analysis yielding a ²⁰⁷Pb/²⁰⁶Pb age of 3876 ± 26 Ma (2) is slightly older than the main c. 3·84 Ga group and might represent a genuinely older grain. Four analyses of CL-dark cores of subgroup 1b have high-U concentrations (1800–8300 ppm), and low Th/U ratios (<0·04). Only one concordant date has been obtained at c. 2·72 Ga, with the other analyses plotting between 27 and 47% discordant. One of the discordant analyses has a ²⁰⁷Pb/²⁰⁷Pb age of c. 2·78 Ga, suggesting that it cannot simply represent an equivalent of the concordant 2·72 Ga analysis that has undergone later Pb-loss but probably contains an inherited component or is a much older grain (i.e. early-Archaean) that has experienced extreme, possibly multiple, Pb-loss.

Group 2 ‘soccer ball’ zircon grains mostly yield a small range of ²⁰⁷Pb/²⁰⁶Pb dates from c. 3·57 to c. 3·62 Ga. Th/U ratios are very low (<0·03) and U-concentrations fall in the range c. 500–800 ppm. The spread to younger dates is attributed to minor Pb-loss from originally 3·62 Ga zircons, probably during the c. 2·71 Ga regional high-grade event (although still concordant at the 2 error level, the two analyses with youngest ²⁰⁷Pb/²⁰⁶Pb dates are displaced to the right of the inverse concordia). One zircon core (1–15a) that is provisionally classified into group 2 on the basis of its CL-dark, rounded appearance yields a concordant age of c. 2·71 Ga, with low Th/U (0·025) and high U-concentration (1100 ppm).

Group 3a zircon phases comprise distinct, rather thin rim, overgrowths or embayments developed on oscillatory zoned CL-bright zircon cores of group 1a. Pb-isotope ratios only were obtained from this group using a small primary beam spot (c. 15 µm), together with further enhancement of spatial resolution of c. 11 µm using a field aperture in the secondary ion beam (see Fig. 4). Despite the overall similarity of appearance in CL as well as consistently low Th/U ratios (<0·06), this group yields a distinctly bimodal set of ²⁰⁷Pb/²⁰⁶Pb dates in the range c. 3·70–3·50 and 2·72–2·58 Ga. In line with our earlier interpretations, we attribute this limited spread in zircon dates to Pb-loss affecting zircon grown originally at c. >3·60 and c. 2·71 Ga (Fig. 5). Group 3b zircon (rims on group 1b and group 2 zircon), together with group 4 zircon (rims on polyphase grains) show ²⁰⁷Pb/²⁰⁶Pb dates around 2·7 Ga or lower and consistently low Th/U ratios (<0·1). Six of the seven analyses in group 3b for which full U–Pb data are available lie on a regression with an upper intercept date of 2709 ± 13 Ma (MSWD = 2·2, 95% confidence); a subset of four of these analyses define a concordia age of 2709 ± 7 Ma (MSWD = 0·73, 2; decay constant errors included). A somewhat older, discordant age of c. 2·95 Ga from analysis 1–12a probably represents a mixed analysis between a c. 2·7 Ga rim and older core, as a ²⁰⁷Pb/²⁰⁶Pb age of 2741 ± 26 Ma (2) was obtained from a subsequent higher spatial resolution analysis.

The U–Th–Pb data that we obtain from sample SM/GR/98/8 are in good agreement, both in maximum age and age range, with data reported by Nutman et al. (1997) and Mojzsis & Harrison (2002) from their respective samples of the homogeneous tonalite from this locality, again suggesting that SM/GR/97/8 is the equivalent of these previously investigated samples.

Zircon REE geochemistry
REE concentration data are presented in Table 3, in a series of normalized [to C1 chondrite values of Anders & Grevasse (1989)] abundance diagrams (Figs 6–9), and in plots showing the degree of LREE and HREE fractionation against age (Fig. 10a and b). Zircon samples from group 1a, i.e. oscillatory zoned c. 3·6 to >3·8 Ga cores of grains, show a limited range of compositions in Fig. 6. Most analyses show typical ‘igneous’ zircon REE profiles with closely parallel steep slopes from La to Lu, marked positive-Ce and small negative-Eu anomalies that are very similar to the profiles reported from other early-Archaean zircon grains in the Godthåbsfjord (e.g. Whitehouse & Kamber, 2002, 2003). Two of these five profiles (2a–x and 10a–x) represent analyses of the same growth phase (based on CL imaging) that was analysed for U–Th–Pb but in a slightly different location. Analyses from the original dated location show markedly increased LREE concentrations (omitted from Fig. 6 but data are retained in Table 3) that might be related to later LREE mineralization along cracks, perturbing the original profile. The exact nature of this mineralization is unclear but, given the clearly LREE-depleted nature of the original zircon, only a very small amount of a strongly LREE-enriched phase (e.g. monazite or allanite) would be required to influence the REE profiles, whereas other characteristic enrichments (e.g. higher levels of Th in monazite) remain undetected. Analysis 2–5a exhibits a highly unusual REE profile that is enriched in both LREE and MREE. Because it is unlikely that zircon can incorporate REE in this pattern [by see discussion in Whitehouse & Kamber (2002)], we conclude that the REE profile of this particular grain must be controlled by inclusions (and/or microcrack fillings) of both LREE- and MREE-enriched minerals, such as monazite, allanite or apatite. This analysis is not considered further for the purposes of protolith characterization. In general, the REE profiles from group 1a, together with Th/U ratios, are typical of zircon crystallized in equilibrium with a primary melt.

View larger version (38K): [in this window] [in a new window]   Fig. 6. Chondrite-normalized (C1, Anders & Grevasse, 1989) abundances of REE from zircon growth belonging to group 1a (oscillatory zoned cores) and group 1b (CL-dark cores). Dashed lines connect Sm and Gd for HREE-only analyses in which Eu was not determined. Horizontal spacing corresponds to effective ionic radius in eightfold coordination (Shannon, 1976; indicated in pm on the upper x-axis).

 

View larger version (31K): [in this window] [in a new window]   Fig. 7. Chondrite-normalized abundances of REE from zircon belonging to group 2 (‘soccer-ball’ zircon). Normalization and axis spacing as in Fig. 6.

 

View larger version (32K): [in this window] [in a new window]   Fig. 8. Chondrite-normalized abundances of REE from zircon growth belonging to group 3a (overgrowths or embayments on oscillatory zoned cores of group 1a), differentiated into (a) early Archaean and (b) late Archaean. Dashed lines connect Sm and Gd for HREE-only analyses in which Eu was not determined. Normalization and axis spacing as in Fig. 6.

 

View larger version (41K): [in this window] [in a new window]   Fig. 9. Chondrite-normalized abundances of REE from zircon growth belonging to group 3b (late Archaean rims on group 1b zircon) and group 4 (outer late Archaean rims on polyphase zircon). Dashed lines connect Sm and Gd for HREE-only analyses in which Eu was not determined. Normalization and axis spacing as in Fig. 6.

 

View larger version (18K): [in this window] [in a new window]   Fig. 10. Temporal evolution of (a) extent of LREE depletion represented by (Pr/Sm)N, (b) HREE enrichment expressed as (Dy/Yb)N, and (c) Th/U ratios. Dashed vertical lines represent the age (in Ma) of defined events in the Godthåbsfjord region, as labelled in (a).

 

View this table: [in this window] [in a new window]   Table 3: SIMS REE, Y and Hf data from Akilia sample SM/GR/97/8

 
Group 1b zircon samples also represent cores, but these show markedly different characteristics from those of group 1a, particularly with regard to the LREE, which are consistently flat at chondrite-normalized values between c. 40 and 1000, with absent to small positive Ce-anomalies (Fig. 6). The HREE profiles of this group are either highly fractionated, similar to those of group 1a, or in one case (analysis 1–4b), flat at c. 100x chondritic.

Group 2 zircon grains exhibit an internally coherent set of REE profiles (Fig. 7) that are notably different from those of either group 1a or group 1b. The most prominent difference is the marked depletion of the HREE with flat or even concave-up profiles from Dy to Lu, irrespective of whether the analysed grain is early or late Archaean. These zircon grains show more variable fractionation of LREE relative to group 1a, with the early Archaean grains showing the steepest profiles. In general, these analyses show smaller-magnitude Ce-anomalies and larger negative Eu-anomalies. Interestingly, the least suppressed positive Ce-anomaly (analysis 14a2) is accompanied by the least enhanced negative Eu-anomaly—a feature that would be consistent with variable redox control of the behaviour of these two elements during the event generating the c. 3·6 Ga zircon. The characteristic HREE depletion of this group (and possibly in analysis 1–4b of group 1b) is the most important observation, as it suggests crystallization in an environment depleted in HREE, probably because of the role of garnet.

Group 3a zircon grains show a wide range of REE profiles (Fig. 8) that probably reflect different origins within this group. The similarities of individual analyses to those of either group 1a and 1b or group 2 are apparent in comparison of profiles in chondrite-normalized abundance diagrams and are particularly well defined in plots of (Pr/Sm)_N and (Dy/Yb)_N against age. The characteristic HREE depletion of group 2 zircon is evident in some of both the early and late Archaean analyses of this group, suggesting crystallization from a garnet-influenced medium, although fractionated HREE profiles are also seen in both age groups. LREE profiles show a range of fractionation in the early Archaean analyses from steep profiles similar to those of group 1a through to somewhat flatter (but still fractionated) profiles. The late Archaean analyses show only flatter LREE profiles.

Analyses of zircon in group 3b, all of which are late Archaean in age, exhibit a wide range of fractionation in both HREE and LREE (Fig. 9a). In two cases, the corresponding core (group 1b) and rim (group 3b) analyses may be compared. For zircon grain 1–4, the core (1–4b) and rim (1–4a) REE systematics are very similar, whereas for grain 1–16, the rim (1–16a) shows a similar degree of fractionation of HREE at lower concentrations relative to the core (1–16b), but a strongly reverse fractionated LREE profile.

Group 4 zircon analyses both show fractionated HREE, in one case (2–36c) at a low overall concentration level, whereas the LREE of analysis 2–23c show a flat profile (Fig. 9b). It is interesting to note that the HREE of analysis 2–23c contrasts with the flat profile seen in the early Archaean group 3a core overgrowth in this same zircon grain (analysis 2–23bx).

	DISCUSSION

TOP ABSTRACT INTRODUCTION REVIEW OF AKILIA GEOLOGY... SAMPLING ANALYTICAL METHODS RESULTS DISCUSSION CONCLUSIONS REFERENCES

 
Equivalence of studied samples
The first point to note about our new zircon geochronological data from SM/GR/97/8 is that they exhibit the same range of ages and, within specific age groups, Th/U ratios as data previously presented (Nutman et al., 1997; Mojzsis & Harrison, 2002) from a homogeneous granitoid sheet within the much debated c. 0·7 m wide composite gneiss layer on Akilia. Taken together with our whole-rock major- and trace-element data, which also show strong similarity to these earlier studies, we are, therefore, convinced that this sample is an appropriate representative of the homogeneous granodiorite sheet from which claims for the >3·83 Ga age for the mafic–ultramafic enclave on southwest Akilia have been derived. A previous claim (Whitehouse et al., 1999) that another sample from the composite gneiss sheet (SM/GR/97/7) was the equivalent, despite its younger c. 3·65 Ga zircons, was based on an incomplete understanding of the complexity of the gneiss sheet (Nutman et al., 1997), that was subsequently revealed by Myers & Crowley (2000), Nutman et al. (2002) and Whitehouse & Fedo (2003); this claim is now considered invalid, although data from this sample remain valid.

In this paper, we also reiterate the detailed field observations of Whitehouse & Fedo (2003), which indicate (1) that the homogeneous granodiorite sheet cannot be demonstrated to cross-cut the mafic–ultramafic enclave and hence cannot provide an age constraint, and (2) that the homogeneous granodiorite itself appears to cross-cut a banded gneiss (SM/GR/97/7) that contains typically igneous oscillatory-zoned zircons with an age of c. 3·65 Ga (Whitehouse et al., 1999). Thus, the possibility needs to be considered that the preserved discordance could indicate intrusive remobilization of the granodiorite 3·65 Ga. Against this background, >3·65 Ga oscillatory zoned zircons in the granodiorite could be remnants from a c. 3·84 Ga protolith that were not digested during the c. 3·62 and 2·71 Ga remelting and remobilization events.

Although it is beyond any doubt that 3·75–3·84 Ga granitoids existed in the Godthåbsfjord region, the critical issue on Akilia Island is not to provide evidence for relict original >3·8 Ga features (such as zircon cores), but to establish which aspects of the gneiss as it presents itself today can be attributed to known stages of the regional geology. The key question in this regard is to what extent the gneiss was remelted and deformed during the 3·6 and 2·71 Ga events. Only then can relative age relationships seen in the outcrop be dated. In the following sections, we discuss our new data together with previously published observations in the context of three possible emplacement scenarios.

Emplacement models consistent with zircon data and field observations
Model 1: 3·60–3·65 Ga emplacement of the granodiorite with inheritance of older zircon
As previously discussed for other TTG gneisses from the Godthåbsfjord, the distribution of zircon ages in SM/GR/97/8 and equivalent samples is formally consistent with a magmatic age of 3·60–3·65 Ga. This age is within error of that of the immediately adjacent sample SM/GR/97/7, which appears structurally older (Myers & Crowley, 2000; Whitehouse & Fedo, 2003). There is thus no conflict between relative and absolute age dates in this interpretation, in which >3·65 Ga zircon is viewed as inherited. The possibility that melts of tonalitic–granodioritic composition can entrain older zircon has long been a vexed issue in discussion of the protolith ages of rocks from the Gothåbsfjord region. Several workers, applying the zircon solubility arguments of Watson & Harrison (1983), considered that the assumed high temperature and low-Zr concentration of most TTG melts a priori preclude survival of older zircon (Nutman et al., 1997; Mojzsis & Harrison, 2002). We note that these studies universally ignore the factors considered by Watson (1996) to favour the survivability of inherited zircon, namely (1) presence of large zircon in the source rock or occlusion of zircon grains within major silicates; (2) a brief or cool magmatic event; and (3) failure of the melt to coalesce into a large body. In our view, two additional important issues that seriously question the validity of rejecting the hypothesis of inherited zircon concern the assumed solidus temperature for the studied rocks and the estimated fraction of zircon which actually survived as xenocrystic grains.

With regard to the homogeneous granodiorite sheet on Akilia, no evidence has ever been presented that it represents, as claimed (e.g. Mojzsis & Harrison, 2000), a high-temperature (c. 900°C) dry melt. Rather, the new trace-element data presented here for SM/GR/97/8, in accordance with published data for similar rocks (Nutman et al., 1999; Kamber et al., 2002), strongly argue against a dry (and hence hot) solidus. Experiments of wet melting of tonalite compositions have shown that very significant melt fractions (30–70%) are still present at temperatures below 750°C (e.g. Peters & Kamber, 1994; Sawyer, 1994). As pointed out by Kamber & Moorbath (1998), a large proportion of typical Amîtsoq gneisses will retain a ‘critical melt fraction’, sufficient still to be mobile, at c. 775°C and at this temperature, they would be Zr-oversaturated and hence capable of preserving inherited zircon. Under these conditions, the criteria noted by Watson (1996) of cool magmatic events and failure of melts to coalesce into large bodies would be met. Because temperature is the single most important parameter determining zircon saturation (Hanchar & Watson, 2003), advocates of a c. 900°C solidus for the studied rock have to produce conclusive supporting evidence, before discarding the possibility of inheritance.

Inheritance of xenocrystic zircon in exactly these rocks has been demonstrated with the >4·0 Ga grain analysed in situ by Mojzsis & Harrison (2002). These workers suggested, however, that survival of a significant proportion of inherited zircon was unlikely (assuming, of course, that their solidus temperature estimate of 900°C is correct). The question, then, is what should be regarded as a significant proportion of inherited grains? As discussed earlier, for understandable reasons, zircon separation, grain selection and ion-probe dating all bias towards large grains of high-quality (low-U) zircon. In our study, we have strived to date a representative selection of grains, although even with these efforts, the majority of zircon grains (many of which are too small for routine handling) have remained untouched. Although it is impossible to estimate reliably the true proportion of >3·8 Ga zircon in this rock, it is sufficient to note from the CL images (Fig. 4) that group 1a zircon grains are not necessarily a predominant component of the total zircon fraction.

We thus maintain that if the homogeneous granitoid sheet was emplaced as a crystal-rich magma, instead of as a completely liquid melt, the observation by Mojzsis & Harrison (2002) that the oldest zircon grains are found in the middle of feldspars and must therefore represent the protolith age is irrelevant, because they would merely represent the age of the source rock from which the homogeneous granodiorite sheet was ultimately derived and have no bearing on the emplacement age of the granodiorite sheet itself, which has been the singular important goal of these studies.

Although a magmatic age of 3·6–3·65 Ga appears plausible from the data discussed thus far, our new REE data for group 2 zircon grains (which best represent growth of zircon in this time period) suggest that this event more probably represented lower-crustal remelting of an older protolith (in the presence of garnet or another HREE-retaining phase). We investigate this in the following two models.

Model 2: remelting of 3·84 Ga gneiss at 3·62 Ga
Structural and geochronological constraints discussed by Whitehouse & Fedo (2003) preclude a >3·65 Ga magmatic age for the homogeneous granitoid sheet. It is therefore likely that the actual magmatic age is represented by one of the two remaining age groups identified: either c. 3·62 or c. 2·7 Ga. In this section, we investigate the constraints imposed by our REE data on these ages.

The distinctive HREE-depleted profiles seen in the group 2 zircon grains, as well as c. 3·62 Ga overgrowths on older cores (group 3a), are typical of zircon that has crystallized and/or recrystallized in equilbrium with a strongly HREE-enriched phase of which garnet is most likely (e.g. Rubatto, 2002; Whitehouse & Platt, 2003). As garnet is not present in this granodiorite, we propose that it was a residual phase during melting of an older protolith. Melting of a garnet-bearing mafic source, although capable of producing a quartzofeldspathic melt, is considered unlikely because a much steeper whole-rock REE profile similar to typical TTGs (Fig. 3a) would be expected (e.g. Martin, 1986). Furthermore, a non zircon-bearing mafic precursor would not yield older zircon, as seen in the cores of many of the SM/GR/97/8 grains. Instead, we propose melting of a tonalite or quartz–diorite protolith during granulite-facies metamorphism at c. 3·62 Ga. Several experimental studies have documented garnet as a reaction product formed by dehydration melting of biotite-bearing quartzofeldspathic rocks (e.g. Rutter & Wyllie, 1988; Skjerlie & Johnston, 1993, 1996; Patiño-Douce & Beard, 1996; Singh & Johannes, 1996a, 1996b). The study of Singh & Johannes (1996a, 1996b) specifically investigated melting of a tonalite and showed that garnet appeared in the melt residues at pressures above 10 kbar and temperatures exceeding c. 750°C, whereas the melts themselves were of granodioritic composition.

An intriguing aspect of the REE composition of 3·62 Ga zircon in this sample is that it appears to record different growth conditions, with the group 2 and some of the group 3a analyses recording the garnet influence on HREE described above, whereas others show fractionated HREE that appear to be unaffected by garnet. This is further emphasized in Fig. 11a, which plots a vector for pairs of analyses from the same grain, showing the change in degree of HREE fractionation (represented by (Dy/Yb)_N) and HREE concentration (represented by [Dy]). Correlation of REE profiles with CL images shows that within group 3a, there is no direct correspondence between morphological character and REE composition, as both obvious embayments (e.g. 2–8b and 2–36a) and more continuous, enveloping overgrowths (e.g. 2–23b and 2–20b) show both types of REE composition. A possible explanation for this could be that the zircon grains record prograde development from early metamorphic–anatectic conditions in which garnet was not present through to later stages when melt was produced in equilibrium with a garnet residue to produce both rims and entire grains (group 2) that are HREE-depleted.

View larger version (27K): [in this window] [in a new window]   Fig. 11. HREE vectors for pairs of analyses from individual zircon grains in (Dy/Yb)N vs Dy concentration (ppm) space for: (a) early Archaean rims of group c. 3a (i.e. developed on c. 3·8 Ga cores) and (b) late Archaean rims of group 3a (developed on c. 3·8 Ga cores), group 3b (developed on late Archaean group 1b cores), and group 4 (outer rims on polyphase grains). Symbol represents composition of zircon on which the rim is developed and the position of the arrowhead represents the composition of the rim itself (note that for grain 2–34 in (a), the change in composition is smaller than the size of the symbol and so the arrowhead has been omitted). Shaded area represents the total compositional range of group 2 zircon.

 
Model 3: remelting of early Archaean gneiss at 2·71 Ga
An interesting and potentially important observation from the complex polyphase zircon crystals in SM/GR/97/8 is that key features of the c. 3·6 Ga growth phases, namely low Th/U ratios and depleted HREE, are shared by some of the 2·7 Ga growth phases (Fig. 10). Similar Th/U characteristics at both 3·6 and 2·7 Ga were considered by Mojzsis & Harrison (2002) to indicate that similar, specifically metamorphic, processes were responsible for these two episodes of zircon growth and/or reworking. It is clear, however, from a plot of Th against U concentrations from our study (Fig. 12) that Th/U ratios of the 3·6 Ga overgrowths (group 3a) and soccer-ball grains (group 2) are, on average, low because of relatively low Th-concentrations rather than U-enrichment. Increased U-contents, on the other hand, are the main reason for low Th/U ratios in 2·7 Ga overgrowths (groups 3a, 3b and 4) as well as late Archaean cores (group 1b). Low Th-concentrations are not unknown from magmatic or migmatitic zircons and, on the basis of Th/U ratios alone, it cannot be inferred that such zircon grew from a metamorphic fluid. The possibility remains, however, that at least some of the 2·7 Ga zircons grew from a U-rich but Th-poor metamorphic fluid.

View larger version (29K): [in this window] [in a new window]   Fig. 12. Plot of Th-concentration (ppm) vs U-concentration (ppm) contoured for Th/U ratios from 0·001 to 10. Group symbols follow other diagrams, except that group 3a is further subdivided into c. 3·6 and c. 2·7 Ga groups. Large symbols labelled with Roman numerals represent average compositions of: I, – >3·8 Ga cores (group 1a); II, – c. 2·7 Ga cores (group 1b); III, – c. 3·6 Ga overgrowths (group 3a); IV, c. 2·7 Ga overgrowths (groups 3a, 3b and 4). Vectors labelled Th-depletion and U-enrichment indicate that lowering of Th/U ratios may be resolved into a combination of these two components.

 
Our new REE data provide additional constraints on the growth environment for the 2·7 Ga zircon overgrowths. In common with the c. 3·6 Ga overgrowths, a wide range of HREE profiles are seen, ranging from steeply fractionated, with low (Dy/Yb)_N, to flat, with (Dy/Yb)_N 1 (Fig. 10b), whereas the LREE are similarly variable and, on average, more enriched (Fig. 10a). As discussed for Model 2, above, the HREE depletion may be consistent with melting of a tonalitic precursor under granulite-facies conditions in which garnet is residual, but, in this case, it would necessarily be remelting of a rock that had already been through such a process. In the earlier proposed melting episode at c. 3·6 Ga, only zircon with typically igneous, fractionated REE from the c. 3·84 Ga precursor was originally present, whereas at 2·7 Ga, previously generated HREE-depleted zircon in the form of group 2 soccer balls and group 3a overgrowths would be present. In this case, mechanisms other than melting may be invoked in order to explain the HREE-depleted rims of groups 3a, 3b and 4, such as dissolution–reprecipitation of existing zircon or Ostwald ripening. If one of these mechanisms is indeed responsible for the c. 2·7 Ga rims, some transport of the REE is required because, as shown in Fig. 11b, there are commonly significant differences between the REE pattern of the 2·7 Ga growth phase and the substrate upon which it is developed. For example, grains 1–12 and 2–23 have HREE-fractionated rims developed on zircon with flat HREE, whereas grains 2–22 and 2–5 have flat HREE rims developed on HREE-fractionated zircon.

Grains 1–4 and 1–15 are of particular significance in interpreting the nature of the 2·7 Ga event. These are grains in which the HREE-depleted phase actually constitutes the core rather than an overgrowth. Superficially, these grains resemble the c. 3·62 Ga group 2 soccer-ball grains (1–4 is assigned to that group on the basis of morphology, despite its 2·7 Ga age) and they appear to indicate that entirely new zircon grains might have been growing at 2·7 Ga—a process that would be consistent with their origin in a partial melt. An alternative explanation, in which the cores of these grains simply reflect total Pb-loss from early Archaean group 2 grains, seems unlikely, given that these latter grains with very similar U-concentration show limited Pb-loss, with ²⁰⁷Pb/²⁰⁶Pb ages slipping only to c. 3·57 Ga. A partial melt may also account better for the contemporaneous feldspar–leachate–whole-rock Pb–Pb date of 2746 ± 150 Ma (MSWD = 1·3) reported from this sample by Kamber & Moorbath (1998) that clearly points to profound disturbance and re-equilibration of the whole-rock U–Pb systematics.

The increasing complexity of zircon growth morphology and REE chemistry by the time of the late Archaean event serves to obscure its true nature. Although we may speculate that a second, granulite-facies partial melting event was responsible for the distinctive HREE-depleted nature of some of the overgrowths and whole grains that yield c. 2·7 Ga ages, we cannot entirely dismiss mechanisms such as dissolution–reprecipitation or Ostwald ripening that might have been responsible for redistributing Zr, Si and REE at this time.

Recognition of magmatic vs metamorphic zircon in quartzofeldspathic gneisses
The debate over the ‘age’ of the granodiorite gneiss sheet on Akilia serves to highlight a semantic but non-trivial issue of recognizing ‘metamorphic’ versus ‘igneous’ zircon. In the present context, proponents of a >3·8 Ga age for the mafic–ultramafic enclave on southwest Akilia (and any possible early terrestrial life relicts) require (1) an unambiguous cross-cutting field relationship with the granodiorite [recently discredited by Whitehouse & Fedo (2003)], and (2) evidence for a purely subsolidus process to explain all post-3·8 Ga zircon in the granodiorite, in order that original igneous field relationships remained undisrupted.

In accord with Hoskin & Black (2000), we consider ‘metamorphic’ zircon to be that generated strictly by subsolidus processes, namely (1) diffusion of Zr and Si released by metamorphic breakdown reactions leading to nucleation and crystallization of zircon (e.g. Fraser et al., 1997; Bingen et al., 2001) and (2) recrystallization of protolith zircon (e.g. Friend & Kinny, 2001). Zircon precipitated from a partial melt, although technically ‘magmatic’, is ultimately a product of metamorphism (e.g. Roberts & Finger, 1997). Under such circumstances, the distinction between metamorphic and magmatic zircon becomes blurred and herein lies the difficulty of defining the actual ‘emplacement’ age of a rock such as SM/GR/97/8. Specifically, in cases where the zircon is simply classified as ‘metamorphic’ on the basis of geochemistry, the possibility remains that it could have formed from sufficient partial melt for its host to be mobile and therefore disrupt and obliterate original field relationships. Using geochemical features of zircon, as well as textural relationships exposed by CL imaging, it is possible to distinguish between zircon crystallized from ‘primary melts’ (see Mojzsis & Harrison, 2002) and zircon that formed from a partial melt during a metamorphic event in which the solidus was reached and metamorphic zircon generated by subsolidus processes. In the concluding section of this paper, we discuss the importance of correctly applying criteria to discriminate between such zircon origin.

Interpretation of zircon Th/U ratios
In their depth-profiling study of a single zircon from an equivalent sample to our SM/GR/97/8 at this locality, Mojzsis & Harrison (2002) assigned a ‘metamorphic’ origin to c. 3·6 and c. 2·7 Ga outer regions of their studied zircon on the basis of low Th/U ratios. The only reasoning forwarded for this assignment was that the expected increase of Th/U ratios in orthogneisses during prograde metamorphism implied that associated metamorphic fluids should be characterized by correspondingly low Th/U. An ad hoc assumption was made by these workers that zircon precipitated from just such a metamorphic fluid, but no suggestion was made regarding the stage at which during prograde metamorphism this precipitation took place and, most importantly, whether Zr (and Si) would have been available to facilitate zircon crystallization.

It is not uncommon for low Th/U ratios to be used to discriminate zircon resulting from metamorphic processes from magmatic zircon, with competing effects of Th-rich minerals such as monazite and allanite typically being invoked (e.g. Williams et al., 1996; Rubatto et al., 2001). This discriminant cannot be applied universally, however, not least because ‘metamorphic’ zircon from high-metamorphic-grade orthogneisses can exhibit high Th/U ratios (e.g. Friend & Kinny, 2001; Möller et al., 2003), whereas other studies have documented low Th/U ratios in unambiguously magmatic zircon (e.g. Zeck & Whitehouse, 1999). In other cases, Th/U ratios appear to be robust through high-grade metamorphism, preserving protolith values (Möller et al., 2002).

Evolution of Th/U ratios during prograde metamorphism of orthogneisses is undoubtedly more complex than envisaged by Mojzsis & Harrison (2002) and depends largely on the location of each element and its ability to be released to a fluid at different times. Geochemical and Pb-isotope studies of metamorphism in the Lewisian Complex of NW Scotland (Fowler, 1986; Whitehouse, 1989) have shown that grain boundary hosted U, which can account for >50% of total U in amphibolite-facies rocks (Dostal & Capedri, 1978), is largely removed by the time hornblende–granulite facies is attained, resulting in higher Th/U ratios in the residual rock. Higher-grade pyroxene–granulite-facies rocks in the Lewisian, however, return to modest and rather homogeneous Th/U ratios, presumably as a result of the onset of breakdown of Th-bearing accessory phase(s) (e.g. allanite, apatite). The low Th/U fluid envisaged by Mojzsis & Harrison (2002) is therefore likely to be present only during the earliest stages of prograde metamorphism, as the labile grain boundary U is lost. The availability of Zr at this stage will depend on whether any early reacting phases release sufficient Zr to facilitate zircon crystallization. Fraser et al. (1997) suggested that Zr is most likely to be released during hornblende and/or garnet (the latter unlikely in the protolith to this particular lithology) breakdown, because these minerals host the highest Zr-concentration, but their breakdown will occur at higher grade than the loss of labile grain boundary U and possible release of a low Th/U fluid. Furthermore, our new REE data require that such a fluid would also have a highly specific (HREE-depleted) composition which is in fact the opposite of that expected to be released by breakdown of HREE-enriched hornblende and garnet. Hence, interpretation of Th/U systematics in zircon needs to consider the metamorphic conditions under which new zircon has grown and the likely mineralogy of the rock from which metamorphic fluids were released.

Mojzsis & Harrison (2002) also discussed the implications of Th/U ratios from their in situ analyses of 71 zircon grains. They attached particular relevance to the observed variation in Th/U ratio in c. 3·6 Ga zircon grains from values consistent with primary magmas to values that they considered to result from fluid precipitation. It is instructive to compare their fig. 7, which summarizes data from their own in situ study and two earlier studies of separated grains with our Fig. 10c, from which it can be seen that neither our data nor those of Nutman et al. (1997) and Mojzsis & Harrison (2000) show elevated Th/U in the 3·6 Ga population. Only the in situ data of Mojzsis & Harrison (2002) show a spread towards elevated Th/U ratios at c. 3·6 Ga. In our view, this spread is probably an artefact of their analytical procedure, which did not utilize any pre-analytical imaging to reveal internal morphology. Given the complexity of zircon internal morphology documented herein (Fig. 4) and elsewhere (Nutman et al., 2000), we consider it likely that many of the in situ analyses of Mojzsis & Harrison (2002) sample mixed zircon growth phases, resulting in both the aforementioned mixed ages as well as mixed Th/U signatures. Hence, although we agree with the interpretation that a partial melt was present in this rock at 3·6 Ga, in our view, as argued above, this was of strictly low Th/U character.

Textural observations
Minimization of analytical artefacts derived from in situ ion-probe data from mixed zircon growth phases as discussed above is not the only benefit of comprehensive mapping of zircon internal structure prior to analysis. Independent of geochemical observations, the internal morphology of complex zircon revealed by CL (as used in this paper) or BSE imaging provides critical information on whether purely subsolidus processes can be invoked to explain later phases of zircon growth or whether a (partial) melt is required. Several features of zircon morphology (Fig. 4) suggest that the >3·8 Ga zoned cores were at some stage in contact with a melt phase. Corrosive features, in particular embayments that cut across the earlier growth zoning, are particularly well illustrated by grains 2–4, 2–21, 2–32 and 2–36. In all cases, there is no evidence of ghost zoning as might be expected for solid-state recrystallization fronts (e.g. Hoskin & Black, 2000) and so we prefer to assign these embayments to the action of a corrosive melt. In addition, several of the zircon grains shown in Fig. 4 show clear evidence for fracturing. Most commonly, this is manifested as truncated zoning where an expected pyramidal termination that would preserve the symmetry of the original >3·8 Ga zircon, e.g. in grains 2–1, 2–10, 2–20 and 2–34, is missing. Grain 2–34 is particularly informative because, in addition to truncation, the grain has cracked down the centre and the internal zoning clearly reveals relative rotation of both parts prior to annealing of the crack with later zircon—a process which simply cannot be achieved in a subsolidus environment.

REE patterns from dated and texturally mapped zircon growth phases
Proponents of a 3·84 Ga intrusion age for the Akilia granodiorite would probably agree with us that the rock experienced partial melting at 3·6 Ga and that at least some of the concordant zircon of that age grew from a partial melt. In the absence of REE data, it could be argued that the extent of partial melting was relatively minor and remained insufficient to remobilize the rock per se, thus favouring preservation of original 3·84 Ga cross-cutting relationships. However, the REE patterns of type 2 zircon are very difficult to explain by ‘minor’ melting because the granodiorite gneiss itself is devoid of garnet and textures that could suggest the former presence of garnet. In our view, this implies that the granodiorite melt formed during prograde upper amphibolite- to peak granulite-facies metamorphism of a quartzofeldspathic, lower-crustal protolith from where the melt migrated to its present location. We envisage that once sufficient melt and strain had accumulated, possibly because of melting reactions involving biotite breakdown, the granodiorite (as we analyse it today) became sufficiently mobile to intrude along older lithological contacts. Judging from the textural and geochemical complexity of the studied zircon, it is plausible that zircon now present in the granodiorite sheet originated as: (1) entrained original igneous grains from the protolith possibly shielded from fluids and melts in silicate phases; (2) original igneous grains that experienced substantial Pb-loss in the 3·6 Ga melt; (3) original 3·84 Ga material that reprecipitated or recrystallized during high-grade metamorphism either during prograde fluid-dominated phases of growth or during melting; (4) freshly crystallized, new igneous 3·6 Ga zircons from Zr that was liberated during amphibole (± garnet) break-down. To what extent these processes were repeated during the high-grade 2·71 Ga event is difficult to estimate as a result of the generally thin zircon growth phases of that age, which make even ion-probe analyses difficult.

The complexity of relative and absolute age relationships of lithologies on Akilia Island demonstrates that new attempts to find evidence for early life on Earth must be based on a sound understanding of the field relationships and the protoliths of the studied lithologies, as well as robust absolute chronology. Elucidation of the geochronology of such target areas should ideally be carried out prior to publication of spectacular claims regarding the origin of life. In any case, establishment of a sound geochronological framework requires careful application of the full spectrum of tools available to the geologist.

	CONCLUSIONS

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Whole-rock major-element composition and U/Pb zircon age distribution leave no doubt that the granodiorite sample analysed in this study is equivalent to the samples analysed by Nutman et al. (1997) and Mojzsis & Harrison (2002). Data for the three samples can therefore be combined for a complete interpretation.

Three age clusters determined here at 3·83, 3·65–3·60 and 2·71 Ga correspond to distinct zircon growth phases identified with CL imaging. Our preferred interpretation is that the >3·8 Ga grains (and portions of grains) are from the melt source of the granodiorite, which we envisage to be a >3·8 Ga lower-crustal TTG.

The granodiorite in its present composition separated from its lower-crustal TTG source during a pervasive granulite-facies event in which crustal melts (such as the studied rock) intruded along zones of melt-enhanced structural weakness. The maximum age for the claimed discordance of this rock with mafic gneisses is thus 3·6–3·65 Ga, in agreement with the fact that it cuts a c. 3·65 Ga foliated gneiss (SM/GR/97/7) previously analysed by Whitehouse et al. (1999).

REE abundance patterns for the dated zircon grains are crucial for interpretation of the 3·60–3·65 Ga zircon growth phases, which were previously interpreted to have formed in equilibrium with a metamorphic fluid (Mojzsis & Harrison, 2002) rather than a melt. However, the HREE-depleted nature of many of these growth phases requires the presence of garnet—an expected residual mineral during granulite-facies melting of tonalites. We agree with previous workers that this growth phase has low Th/U but this feature is not uncommon for zircon crystallizing from partial crustal melts.

The youngest zircon growth phase formed at 2·71 Ga and manifests in a variety of textural contexts. In agreement with the textural observation, 2·71 Ga growth phases show several types of REE patterns, but they are generally richer in U than the older growth phases. It is likely that these growth phases formed from more than one process. Fluid precipitation is certainly a possibility but renewed (partial) melting cannot be dismissed. Hence, the claimed discordance of the rock as it is preserved today could even be as young as 2·71 Ga.

The combined data (field relationships, zircon U–Th–Pb, zircon internal structure, zircon REE, whole-rock major elements, whole-rock trace elements, mineral and whole-rock Pb-isotopes) make this one of the most thoroughly studied rocks ever. Because of the great complexity of the zircon data alone, it is not unexpected that various researchers will regard one or the other interpretation as more likely. This, however, is not the issue at stake. Rather, what matters is the realization that irrespective of the sophistication of analysis, there may never be broad agreement about whether this rock indeed cross-cuts potentially life-bearing lithologies, and, if so, how old the discordance truly is. Therefore, this rock cannot serve to date something as important as the oldest evidence for life on Earth. Attempts to date this milestone in planetary history need to work with samples from low-strain areas, in which the zircon systematics can be interpreted unequivocally.

In general, great caution must be applied to interpreting zircon U–Pb geochronological data from thin gneissic veins in polyphase high-grade metamorphic terranes. In particular, application of zircon solubility arguments require realistic estimates of melt temperature and should take into account factors that apply to melts that do not behave as ‘infinite reservoirs’ (Watson, 1996).

	ACKNOWLEDGEMENTS

 
We are grateful to Stephen Moorbath for providing the sample investigated here, as well as invaluable discussions, as well as Chris Fedo and John Hanchar for valuable discussions on aspects of Akilia geology and zircon geochemistry, respectively. We also thank Roland Maas and John Myers for their detailed and constructive reviews. The ion microprobe facility in Stockholm (Nordsims) is operated under an agreement between the joint Nordic research councils (NOS-N), the Geological Survey of Finland and the Swedish Museum of Natural History. This paper is Nordsims publication 106.

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^* Corresponding author. Telephone: +46 8 519 551 69. Fax: +46 8 519 540 31. E-mail: martin.whitehouse{at}nrm.se

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