Geochemistry of metabasalts and hydrothermal alteration zones ...

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Warrawoona Groups, Pilbara Craton are compared with altered metabasalts from immediately beneath bedded cherts of these groups to provide evidence for ...
Geochemistry of metabasalts and hydrothermal alteration zones associated with c. 3.45 Ga chert and barite deposits: implications for the geological setting of the Warrawoona Group, Pilbara Craton, Australia Martin J. Van Kranendonk & Franco Pirajno Geological Survey of Western Australia, 100 Plain St, East Perth, Western Australia 6004, Australia (e-mail: [email protected]; [email protected]) ABSTRACT: Relatively unaltered metabasalts of the Archaean Coonterunah and Warrawoona Groups, Pilbara Craton are compared with altered metabasalts from immediately beneath bedded cherts of these groups to provide evidence for the depositional environment and hydrothermal alteration processes of crust formation. The geochemistry of relatively unaltered basalt, stratigraphy, and inherited zircon data indicate that the lower Warrawoona Group (3.53–3.43 Ga) formed as an oceanic plateau complex built on a sialic basement to 3.724 Ga, following an analogue with the Phanerozoic Kerguelen oceanic plateau, and not as a mid-ocean ridge or convergent volcanic-arc complex as previously proposed. Advanced argillic, argillic, phyllic, and propylitic alteration zones in footwall basalts of this succession are products of repeated episodes of seafloor hydrothermal circulation, syngenetic with bedded chert deposition, in the distal parts of high-sulphidation epithermal systems. The upper part of the Warrawoona Group (3.350–3.315 Ga Euro Basalt) represents a continental flood basalt event, up to 8 km thick, that erupted onto the older succession across a regional unconformity on which the Strelley Pool Chert was previously deposited. Widespread silica–alunite alteration of dolomitic chert protoliths and phyllic and propylitic alteration of footwall basalts are interpreted as products of fluid circulation driven by heat from the overlying, newly erupted lavas. KEYWORDS: Archaean, Pilbara Craton, basalt geochemisty, hydrothermal alteration, mineralization

INTRODUCTION The Archaean Warrawoona Group in the Pilbara Craton, Western Australia, contains some of Earth’s best-preserved, oldest rocks. The group consists of a thick section of dominantly basaltic rocks, intercalated with subordinate felsic volcanic rocks and cherty metasedimentary rocks. The tectonic setting of the former and the origins of the latter are controversial. For example, Kitajima et al. (2001), and Terabayashi et al. (2003) have suggested that basaltic rocks of the group in the North Pole Dome represent several slices of obducted oceanic crust and that hydrothermal chert–barite deposits within the basalts were formed at a mid-ocean ridge. Alternatively, Green et al. (2000) and Van Kranendonk et al. (2002) presented data on the lowermost, dominantly basaltic part of the stratigraphy and suggested it had been contaminated by continental crust during eruption and therefore did not represent true oceanic crust. Arndt et al. (2001) suggested that the lower parts of the stratigraphy (Coonterunah Group and Talga Talga Subgroup of the Warrawoona Group) formed from a depleted mantle source in a largely oceanic setting, as an oceanic plateau that was accreted rapidly onto the Pilbara Craton. These authors suggested that the younger part of the Warrawoona Group (Euro Basalt) formed in a convergent margin setting at the margins of the older oceanic plateau. Others have suggested that the Warrawoona Group initially formed in an oceanic plateau or Geochemistry: Exploration, Environment, Analysis, Vol. 4 2004, pp. 253–278

marginal (back-arc) basin that evolved into a convergent volcanic-arc setting (Barley 1993, 1997; Krapez 1993; Barley & Pickard 1999). Similar controversy exists in regard to the origin of the bedded chert  barite horizons that host stromatolites and (possible) microfossils. These rocks are considered by many to be evaporitic and clastic sedimentary deposits that were silicified by later fluid circulation (Groves et al. 1981; Lowe 1983; Buick & Barnes 1984; Buick & Dunlop 1990; Schopf 1993; Shen et al. 2001) during either shallow seafloor alteration (e.g. Barley 1984), structural doming at c. 3.3–2.9 Ga (Hickman 1984), faulting during flood basalt volcanism at c. 2.7 Ga (Buick 1988), and/or Caenozoic weathering (cf. Hocking & Cockbain 1990). However, recent mapping has shown that many of the bedded cherts are intimately related to the emplacement of swarms of hydrothermal silica  barite veins (Nijman et al. 1999; Van Kranendonk 2000; 2004; Van Kranendonk et al. 2001; Brasier et al. 2002). In this paper, we present geological, petrographic and geochemical data from metabasaltic rocks of the Warrawoona Group in the well-preserved North Pole Dome, and from three bedded chert horizons within the group and their feeder silica vein swarms, in order to assess the nature of the alteration associated with the silica-rich vein swarms and bedded chert deposition and the tectonic setting in which they formed. Our rationale is that if silica replacement of bedded rocks were due 1467-7873/04/$15.00  2004 AEG/ Geological Society of London

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to later events (i.e. doming-related faults, flood basalt-related faulting, or weathering), rocks on either side of the moderately to steeply dipping cherts would be equally affected by alteration, whereas in the syngenetic hydrothermal model, hydrothermal alteration associated with the deposition of individual chert horizons would be restricted to the footwall of those chert horizons. Furthermore, it is well established (e.g. Pirajno 1992, and references therein) that the style of hydrothermal alteration developed in different Phanerozoic tectonic environments is distinct: for example, mid-ocean ridges are characterized by white and black smokers (depending on the redox conditions of the fluids) and extensive chlorite development, whereas backarc Kuroko-style environments contain high-sulphur and low-sulphur systems characterized by extensive silica–sericite alteration, and common sulphate and banded-iron-formation deposits. Thus the style of hydrothermal alteration in the Pilbara may allow us to settle the debate on the tectonic setting of the Archaean Warrawoona Group. GEOLOGICAL SETTING The Archaean Warrawoona Group of the eastern Pilbara Craton is a thick succession of dominantly low-grade (prehnite–pumpellyite to greenschist facies) metabasalts with interbedded felsic volcanic horizons (dominantly tuffs, some flows) and numerous beds of chert (Hickman 1983; Van Kranendonk et al. 2002). (In this paper, the term ‘chert’ is used to refer to bedded fine-grained silica-rich rocks, and the term ‘silica’ is used to refer to massive, fine-grained, amorphous silica, usually dark blue–black to light grey in colour, that occur as veins; both rock types have been recrystallized under low-grade metamorphic conditions.) This succession, with a cumulative lateral stratigraphic thickness of c. 18 km, is well dated at between 3490 Ma and 3310 Ma and has consistent stratigraphic way up indicators and progressively upwardyounging age dates in individual greenstone belts, and evidence for autochthonous deposition of all groups (see Van Kranendonk et al. (2002) and references to age data therein; Van Kranendonk 2004). (All quoted ages, except this age of 3490 Ma, are U–Pb zircon dates and have errors of less than 5 Ma. This date of 3490 Ma is a Pb–Pb age on syngenetic galena in barite from the Dresser Formation, which does not have an associated error, but is within 10 Ma; Thorpe et al. 1992b.) The Warrawoona Group has been divided into three subgroups, each of which comprises a lower succession of mafic–ultramafic volcanic rocks overlain by intermediate to felsic volcanic rocks and capped by a major chert unit. Whereas the lowermost two (Talga Talga and Salgash) subgroups were deposited as a broadly conformable package from 3490 to 3426 Ma, they are separated from the younger Kelly Subgroup by a major regional unconformity marked by the Strelley Pool Chert (Buick et al. 1995; Van Kranendonk et al. 2002; 2003b). The Kelly Subgroup is a minimum of 5–8 km thick, the uppermost parts having been eroded by a regional unconformity at the base of the overlying Sulphur Springs and Gorge Creek Groups. The Euro Basalt at the base of the subgroup consists of a basal unit of olivine spinifex-textured komatiite to pyroxene spinifex-textured komatiitic basalt up to 3 km thick, and this passes conformably up into, and is interbedded with, up to 5 km of tholeiitic basalt. The basalts are in conformable contact with up into 1 km of crustally derived rhyolite (Wyman Formation) at the top of the subgroup (Hickman 1983; Van Kranendonk et al. 2002). Major unconformities also mark the base of the younger volcanic Sulphur Springs Group (3255–3235 Ma; Van Kranendonk et al. 2002; Buick et al. 2002), the dominantly sedimentary Gorge Creek Group

(c. 3235–3050 Ma), and the coarse clastic sedimentary De Grey Group (2970–2930 Ma: Van Kranendonk 2000; Smithies et al. 2001; Van Kranendonk et al. 2002). Previously, the major chert horizons of the Warrawoona Group in the East Pilbara were correlated across greenstone belts and ascribed to the Towers Formation, after the type area near the town of Marble Bar (Hickman 1983; 1990). However, more recent detailed mapping and geochronology has shown that the major chert horizons of the greenstone belts represent three distinct stratigraphic units (Fig. 1; Van Kranendonk 2000; Van Kranendonk et al. 2002). These include from base to top: (1) up to five chert  barite  carbonate  jasper beds interbedded with pillowed metabasalt in what is now known as the Dresser Formation (Van Kranendonk et al. 2002), which has been dated by Pb–Pb on barite as c. 3490 Ma (Thorpe et al. 1992b) and which contains both stromatolites and possible microfossils (Walter et al. 1980; Nijman et al. 1999; Ueno et al. 2001; Van Kranendonk 2004); (2) the jaspilitic Marble Bar Chert Member of the Towers Formation at the top of the 3467 Ma Duffer Formation; (3) the silicified quartzite– carbonate rocks of the Strelley Pool Chert, which contain conical stromatolites (Lowe 1983; Hofmann et al. 1999; Van Kranendonk et al. 2001; 2003a) and lie on a variety of older rock types across an erosional surface that is locally a paraconformity on c. 3458–3426 Ma felsic volcanic rocks of the Panorama Formation, a disconformity on the c. 3467 Ma Duffer Formation, or a high angle unconformity on the 3515 Ma Coonterunah Group (Buick et al. 1995; Van Kranendonk et al. 2002). The basalt stratigraphy contains numerous other, thinner chert horizons that, at least in part, are horizons of originally varied protolith that have been replaced by silica (Buick & Barnes 1984; Van Kranendonk 2000). Some of these chert horizons contain putative microfossils (Awramik et al. 1983; Schopf 1993), although those described from the Apex chert have been the subject of recent intense debate (Brasier et al. 2002; Schopf et al. 2002). Brasier et al. (2002) and Van Kranendonk (2004) have shown that, like the other major cherts in the Warrawoona Group, the Apex chert was deposited from a swarm of hydrothermal feeder veins (now recrystallized microcrystalline quartz) that extend up to 750 m below the palaeosurface, and it is from within one of these veins, 25 m below the palaeosurface, that the microfossiliferous sample of Schopf (1993) was collected (see Apex chert section below). In the following sections we describe the geology, petrography and geochemistry of altered and unaltered rocks associated with chert horizons of the Dresser Formation, the Apex chert and Strelley Pool Chert of the Warrawoona Group. Dresser Formation Geology. The Dresser Formation crops out in the Panorama greenstone belt in the North Pole Dome and consists of up to five bedded cherts interbedded with pillow basalts (Fig. 2; Nijman et al. 1999; Van Kranendonk 2000). Chert horizons contain predominantly grey and white layered chert, thick sections of chert interlayered with stratiform barite, and lesser amounts of jaspilitic chert, carbonate–chert laminates, and clastic sedimentary rocks (Buick & Dunlop 1990; Nijman et al. 1999; Van Kranendonk 2004). Buick & Dunlop (1990) presented evidence for very shallow water deposition of at least some of the sediments, and desiccation cracks have locally been observed. Sandstone and conglomerate that contain clasts of silicified komatiite and barite, amongst other lithologies, disconformably overlie the main chert–barite horizon(s) of the formation and were deposited towards the end of chert–barite

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Fig. 1. Schematic geological map of part of the study area in the East Pilbara Granite-Greenstone Terrane. Numbered rectangles refer to subsequent figures. Circled numbers refer to Localities 1 and 2 of the Strelley Pool Chert discussed in the text.

vein emplacement during period(s) of normal growth faulting, possibly during caldera collapse (Van Kranendonk 2004). Buick & Dunlop (1990) proposed that many of the bedded cherts were originally carbonates or gypsum evaporites that have been replaced by silica. Alternatively, the recognition of a syndepositional relationship of bedded chert–barite with hydrothermal silica  barite feeder veins (see below) supports an interpretation that much of the bedded chert was deposited as primary chemical sediment and that barite was also a primary hydrothermal deposit (Nijman et al. 1999; Runnegar et al. 2001; Van Kranendonk 2004). The chert horizons are deposited above, and fed by, a dense boxwork swarm of weakly radiating silica  barite veins (Fig. 3A) which were synchronous with chert–barite deposition (Figs 3B, 4), synsedimentary slumping, and growth faulting, and possible growth of stromatolites (Nijman et al. 1999; Van Kranendonk et al. 2001; Van Kranendonk 2004). The veins show multiple episodes of silica  barite veining through a single conduit via the crack–seal mechanism of vein growth. Veins may be zoned, with an outer margin of silica and an inner zone of barite, or vice versa, with both types showing clear evidence for contemporaneous crystallization of silica and barite from a single fluid phase (Fig. 3C). The cores of some zoned silica  barite veins show good textural evidence of late epithermal (shallow depth, low temperatures c.150–300 C) quartz in the cores of the veins (Fig. 3D: cf. Panteleyev 1986). Kojima et al. (1998) estimated a temperature of 150 C for the hydrothermal fluids responsible for deposition of the cherts from North Pole, based on sulphide mineral compositions in the chert. Ueno et al. (2001) described carbonaceous

filamentous microstructures from a silica vein that have 13C values of from 42 to 32‰. The interrelationship between veining and sediment deposition is clearly observed at a variety of scales. At the regional scale, silica  barite veins intrude basaltic country rocks up to the base of individual bedded chert–barite horizons, but they do not cross-cut them (Figs 2, 3B). At the outcrop scale, feeder veins contain the same components as the bedded horizons, including chert, barite and ferruginous gossan (Fig. 4), and all of these components occur as clasts in immediately overlying sandstones (Van Kranendonk 2004). The complexity of chert– barite veining and its relationship to sedimentation is shown in Figure 5. At this location, layered primary grey and white chemical sedimentary chert with evidence of deposition under the influence of weak currents overlies wrinkle-laminated, baritic stromatolite and is cut by two types of hydrothermal chert breccias, and later barite veins (Fig. 5). Alteration assemblages. The silica  barite feeder veins under the Dresser Formation are associated with an extensive area, up to 1.5 km stratigraphically beneath the cherts, of highly altered footwall metabasalt (Fig. 2). The zone of alteration is marked by a stratigraphically highest zone, at the base of the chert, of soft, white-weathering rocks composed entirely of clay but in which primary volcanic textures are perfectly preserved (Fig. 6A–D), characteristic of advanced argillic alteration (e.g. White & Hedenquist 1995). The mineral assemblage identified in slightly less altered rocks immediately beneath this zone is quartz–white mica–rutile  chlorite  epidote (Terabayashi et al. 2003), characteristic of phyllic alteration.

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M.J. Van Kranendonk & F. Pirajno metabasalts that are characterized by extensive carbonate (calcite, dolomite and Fe-dolomite: Kitajima et al. 2001) alteration and a mineral assemblage of epidote–chlorite–Ca-Na plagioclase and Ca-amphibole (e.g. Kitajima et al. 2001), characteristic of propylitic alteration (White & Hedenquist 1995). Petrography. Hydrothermal silica  barite veins and silica breccias of the Dresser Formation are characterized by several pulses of silica  barite, including complex stockworks and late open space fillings (Figs 3D, 5). Direct evidence for a hydrothermal origin for the silica  barite veins occurs in the discordant vuggy microcrystalline quartz breccia from the area of Figure 5 in which ragged, white microcrystalline quartz– kaolinite–barite veins are observed to cut through silicified sediment, creating a “pseudo-conglomerate” (Fig. 7A). The presence of kaolinite within the microcrystalline quartz matrix of the breccia (Fig. 7B) suggests that the kaolinite is of hydrothermal origin and not due to weathering. This interpretation is supported by X-ray diffraction (XRD) analyses of the vein kaolinite, which has a markedly different diffraction profile from kaolinite in recent soils (R. Clarke, Chemistry Centre of Western Australia, written communication 2002: comparative XRD diffractograms available on request from corresponding author).

Fig. 2. Simplified geological map of the Dresser Formation in the Barite Range of the North Pole Dome (see Fig. 1 for location). Direction of view for Figures 3A and 3B indicated.

These types of alteration assemblages are in sharp contrast to those of metabasalts immediately overlying the bedded chert– barite horizons and to metabasalts >1.5 km stratigraphically beneath the cherts and outside of the zone of intense hydrothermal alteration. These relatively unaltered metabasalts are characteristically mid- to dark brown weathering, pale to medium green on fresh surfaces, well indurated, and have characteristic low-pressure metamorphic mineral assemblages of epidote–chlorite–actinolite  albite  quartz  hornblende that correspond to lower greenschist and epidote– amphibolite facies temperatures of c. 350C at low pressures (Kitajima et al. 2001). Between the zone of white, clay-altered pillow basalts immediately beneath the cherts and the relatively unaltered metabasalts >1.5 km stratigraphically beneath the cherts are

Apex chert Geology. The Apex chert is one of several chert horizons that occur between thick basalt flows of the Apex Basalt in the Marble Bar greenstone belt (Figs 8, 9). The bedded chert unit consists of up to 10 m of white and black layered chert that is interbedded with felsic tuff that contains sills of massive black silica. The bedded deposits overlie a swarm of weakly radiating black silica veins that extend for up to 750 m stratigraphically down into the footwall metabasalts, but which do not penetrate above the bedded chert horizon (Brasier et al. 2002; Van Kranendonk 2004). The bedded chert–tuff horizon forms laterally discontinuous lenses situated stratigraphically above individual silica vein swarms in the footwall (Fig. 8). The veins are comprised dominantly of massive dark blue–black silica, but can include multiple generations of dark grey to black silica to white quartz, core zones of felsic tuffisite breccia, and phreatomagmatic breccias with jigsaw puzzle fit of exploded fragments at their tops (Van Kranendonk 2004). Brasier et al. (2002) described hydrothermal minerals from the veins and altered basaltic host rocks beneath the bedded chert–tuff horizon, including barite, alunite–jarosite, Al- and K-rich phyllosilicates from hydrothermally altered feldspars, and trace amounts of native metals (Fe, Ni, Cu, Zn, Sn), from which they deduced relatively high hydrothermal temperatures (c. 250–350˚C). A dolerite sill emplaced into the underlying Marble Bar Chert Member matches the distribution of one of the hydrothermal vein swarms in the overlying basalts (Fig. 8), suggesting that the heat from such sub-volcanic sills was the driving force for the hydrothermal circulation. Petrography. Samples of metabasalt from below, at, and above the Apex chert all display some degree of non-pervasive to locally pervasive hydrothermal alteration. This alteration is dominated by microcrystalline quartz, euhedral open-space filling quartz in hydrofractures, carbonate (dolomite and Fe-dolomite), chlorite, sericite and leucoxene. The original basaltic textures are clearly recognizable in all samples in this low strain area and rocks stratigraphically above the chert are little altered by minor chlorite. In contrast, metabasalts below the chert exhibit pervasive carbonate (dolomite and Fe-dolomite) and chlorite alteration and extensive quartz-filled hydraulic fracturing.

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Fig. 3. Features of hydrothermal amorphous silica–barite veins of the Dresser Formation, North Pole Dome. (A) Aerial view to WSW of boxwork amorphous silica–barite veins (high-standing ridges) in altered metabasalt (recessive-weathering material) stratigraphically underneath bedded chert–barite horizons of the Dresser Formation; small gum trees in midground are 3 to 4 m high. (B) View to east of hydrothermal amorphous silica veins in altered metabasalt feeding bedded chert–barite horizon running along the top of the ridge, which dips 40 away from the viewer. (C) Plan view of a vertical, zoned barite (grey, outer areas with euhedral, prismatic crystals) – amorphous silica (white central area), showing evidence of crystallisation from a single hydrothermal event. (D) Zoned epithermal quartz cavity fill in zoned amorphous silica–barite vein. (see Fig. 2 for locations of A and B).

Strelley Pool Chert Geology. The Strelley Pool Chert occurs in all greenstone belts over a distance of 220 km across the east Pilbara. The chert consists of a silicified unit of basal clastic rocks, including quartzite and conglomerate, a main unit of variably silicified evaporitic rocks, including laminated dolomite and silicified dolomite with discontinuous horizons of conical stromatolites, beds of planar to wavy laminites of undetermined origin (Lowe 1983; 1994; anhydrite?, Van Kranendonk 2000), and an overlying unit of conglomerate containing clasts of dolomite, black chert and volcanic rocks (Lowe 1980; 1983, 1994; Buick et al. 1995; Hofmann et al. 1999; Van Kranendonk et al. 2001; 2003a). As with the other two chert units described above, recent mapping has shown that the Strelley Pool Chert is transected by numerous blue–black silica veins that cross-cut the footwall basalts and pass upwards to various stratigraphic levels within the chert, but do not extend into metabasalt above the bedded chert (Van Kranendonk et al. 2001; Van Kranendonk 2004). These veins are commonly zoned with black silica outer margins and white quartz internal zones (Fig. 10A), the latter including local vughs partly filled by coarse quartz crystals (Fig. 10B). The characteristic signature of hydrothermal processes identified in the previously described cherts is also evident in this unit, including a lower zone of carbonate– chlorite alteration (Fig. 10C) and quartz–sericite and pyrophyllite–quartz–sericite alteration of footwall volcanic rocks

immediately stratigraphically beneath the chert (Fig. 10D); these contrast with the relatively unaltered, weakly metamorphosed basalts above the chert (Fig. 10E). Suites of unaltered and altered metabasalt were collected from two localities of the Strelley Pool Chert (Fig. 1). At locality 1 in the North Pole Dome, the chert lies unconformably on pyrophyllite–quartz–sericite schist derived from altered mafic volcanic rocks, whereas in locality 2, the chert lies with high angular unconformity on tilted, folded, and quartz–sericite altered pillowed metabasalts of the Coonterunah Group (Fig. 11). Petrography. Metabasalts well below the Strelley Pool Chert exhibit well-preserved volcanic textures but are pervasively carbonate–chlorite altered (Figs 10C, 12A). Metabasalts immediately beneath the chert are affected by pervasive pyrophyllite–quartz–sericite–leucoxene and quartz–sericite– leucoxene alteration that has rendered the original basaltic texture barely recognizable (Fig. 12B). In contrast, basaltic rocks above the chert are extremely well preserved and affected only by lower greenschist facies metamorphism, with mineral assemblages of chlorite–calcite–epidote (Fig. 12C). The finely layered rocks of the Strelley Pool Chert are characterized by several generations of microcrystalline and coarse-grained quartz, containing euhedral and zoned alunite (Fig. 12D).

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Fig. 4. Geological sketch map of a vertical cross-section through the Dresser Formation (see Fig. 2 for location). (A) Overview of the relationship between a zoned hydrothermal feeder vein of amorphous silica–barite–Fe–gossan that penetrates partway up into, and terminates within, a bedded unit of chert–barite–Fe–gossan sediment. Locality indicated on Figure 2. (B) Closeup of the vein termination (location shown in A). Note the growth of a stromatolite mound adjacent to the mouth of the vein, and of the development of stromatolitic laminations near the top of the vein, contrasting with the massive nature of the barite in the core of the vein. (from Van Kranendonk 2004).

GEOCHEMISTRY Two suites of samples were analysed for this study. The first is a complete section through the c.14 km thick, dominantly basalt stratigraphy of the Warrawoona Group across the southern limb of the North Pole Dome, including seven samples from below the Dresser Formation (d3490 Ma: North Star Basalt or Coonterunah Group), two from within the Dresser Formation (c. 3490 Ma), ten from above the Dresser Formation to the base of the Panorama Formation (c. 3480–3460 Ma: Mt Ada and Apex Basalts), and three from the Euro Basalt (3350–3325 Ma) above the felsic volcanic, c. 3458–3426 Ma Panorama Formation (Table 1: for sources of age data see references in Van Kranendonk et al. 2002). Samples collected for analysis were of the least altered rocks at c. 100–500 m spacing through this transect. Visible carbonate alteration was noted in some outcrops as indicated in Table 1, but this varied across the transect as also indicted by Terabayashi et al. (2003). A good measure of the degree of rock weathering is the Chemical Index of Alteration (CIA=Al2O3/sAl2O3+Na2O+K2O+CaOd), in which high CIA values (>0.80) commonly result from recent, intense weathering and from feldspar-destructive, chloritic or advanced argillic alteration (e.g. Nesbitt & Young 1982). Although Nesbitt & Young (1982) suggested that this calculation incorporates only the amount of CaO in the silicate fraction of the rock and that a correction factor for apatite and carbonate

minerals should be applied, our analyses did not include CO2 thus precluding our ability to make such a correction. However, for our samples, such a correction would only increase CIA values up by c. 0.1, as both strongly carbonate-altered samples and relatively unaltered samples from the North Pole Dome fall in the same general range from 0.44 to 0.61, which is slightly above the values of 0.30–0.45 for unaltered basalts indicated by Nesbitt & Young (1982) (Fig. 13A; data from Table 1). In addition to this transect, samples of least altered volcanic rocks below the studied chert horizons, of highly altered volcanic rocks immediately beneath the chert horizons, and of unaltered volcanic rocks above the three studied chert horizons were collected for geochemical analysis in order to determine the loss and gain of elements during hydrothermal alteration (Table 1). Analytical procedures and precision Geochemical analyses were performed at Geoscience Australia. Rock samples were first crushed using a Nugget jaw crusher and then ground in a tungsten carbide ring mill. Major elements (Si, Ti, Al, Fe, Mn, Mg, Ca, Na, K, P and S) were determined by wavelength-dispersive X-ray fluorescence (XRF) on fused discs using methods similar to those of Norrish & Hutton (1969). Precision for these elements is better than 1% of the reported

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Fig. 5. Geological sketch of a vertical cross-section through the lower part of the main chert horizon of the Dresser Formation. Locality shown in Figure 2. Note the two types of hydrothermal chert breccias that cut obliquely across bedded grey and white chert.

values. Loss on ignition (LOI) was determined by gravimetry after combustion at 1100 C. FeO abundances were determined by digestion and electrochemical titration using a modified methodology based on Shapiro & Brannock (1962). Duplicate analyses indicate a half relative deviation of 0.2%. Barium, Cr, Cu, Ni, Sc, V, Zn and Zr were determined by wavelengthdispersive XRF on pressed powder pellets using methods similar to those described by Norrish & Chappell (1977) and are precise to better than 1% of the reported values. Cesium, Ga, Nb, Pb, Rb, Sr, Ta, Th, U, Y and the rare earth elements (REE) were analysed by inductively coupled mass spectrometer (ICPMS: a Perkin Elmer ELAN 6000) using methods similar to those of Eggins et al. (1997) as described by Pike (2000). These elements have a precision better than 3% of the reported values. Results Basalt geochemistry. The relatively unaltered metabasalt samples from the North Pole Dome transect vary compositionally from tholeiitic basalts to basaltic komatiites with SiO2 between 43 and 52 wt%, MgO between 4 and 14 wt% and Al2O3 between 4 and 15 wt% (Table 1). Many samples have high loss on ignition values, indicating a high degree of carbonate alteration. The magnesium number (Mg#={fMgO/40.32g/ sfMgO/40.32g+fFeOstd/71.85gd}100) shows systematic, progressive changes upsection, from generally high values in

rocks below and within the Dresser Formation, and lowest values above the Dresser Formation (Table 1). The highest MgO abundances are from a group of metabasalts stratigraphically lowest down in the succession, below the Dresser Formation (samples 159632–159636), with low Al2O3 (4.6–9.6 wt%), high MgO (12–14 wt%), high Cr (1470–1934 ppm) and high Ni (289–557 ppm) indicative of olivine accumulation. These rocks contrast with the rest of the samples from this stratigraphic level that have relatively high Al2O3 (12–15 wt%) and lower MgO (13.5–8.9 wt%). However, these major element differences are not reflected by the rare earth element (REE) patterns of the two groups of rocks, which are characterized by flat to slightly light rare earth element (LREE)-enriched profiles without negative Eu anomalies at c. 10 chondrite (Fig. 13B; Sun & McDonough 1989). Metabasalt samples from below the Dresser Formation all the way up to the base of the Panorama Formation are similar, with slightly LREE-enriched patterns (Fig. 13B) and no negative Eu anomalies. Metabasalts from above the Dresser Formation to below the Panorama Formation have generally higher REE abundances, but similar overall LREE enrichments (La/Yb = 2.00–3.87, average 3.25) relative to those from below the Dresser Formation (La/Yb=1.31–4.33, average 3.26). Samples of high-Mg basalt from the base of the Euro Basalt, high in the Warrawoona Group stratigraphy of the North Pole Dome, differ from their stratigraphically lower counterparts in having flat REE patterns and small but consistently negative Eu

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Fig. 6. Features of white, hydrothermally altered pillow basalts from the steam-heated clay alteration zone immediately beneath the main bedded chert horizon of the Dresser Formation at the Dresser Mine (see Fig. 2 for location). (A) View looking east of zoned chert–barite vein cutting hydrothermally altered pillow basalts. Locations of Figure 7B–D indicated by rectangle. Locations of samples collected for geochemical analysis indicated by numbers. (B) Perfectly preserved pillow structures in shallow-dipping, white, clay-altered metabasalts (location indicated in A). (C) Detail of pillow margin in B, showing small pipe vesicles. (D) Detail of quartz-filled amygdales. Lens cap in C and D is 3.1 cm in diameter.

anomalies (Fig. 13B). These rocks are similar to stratigraphic correlatives to the west (Green et al. 2000; Van Kranendonk 2000) and to the southeast (Arndt et al. 2001; Bagas & Van Kranendonk 2002), the latter of which (from higher up in the formation) show good evidence for crustal contamination in the form of slight LREE enrichment, small negative Nb anomalies, and lower Nd values (Arndt et al. 2001). Almost all samples of North Pole metabasalts fall within the normal-type mid-ocean ridge basalt (N-MORB) field on a plot of Nb vs. La, with only one sample in the enriched MORB (E-MORB) field (Fig. 13C), although the REE patterns and low K/Rb, K/Ba and Sr/Rb ratios are more characteristic of E-MORB (Sun & McDonough 1989). Extended trace element spider diagrams of North Pole basalts show remarkable similarity with oceanic plateau basalts (Fig. 13D). Some samples higher up in the succession that have low MgO contents (e.g. samples 159641–159644, 159647) may have suffered hydrothermal alteration, the characteristic signature of which is identified by elevated Na2O and LOI and lower CaO contents (see Alteration Geochemistry section below). Alteration geochemistry. Samples of altered metabasalts from stratigraphically well below (1–2 km), from just below, and

from just above each of the studied chert horizons were analysed in order to characterize the geochemical characteristics of different alteration types and element mobility (see Table 1). Generally, metabasalts from stratigraphically above the cherts and those from well below the cherts have relatively unaltered geochemistry similar to the bulk of analysed samples from the North Pole Dome. Basalts sampled from immediately beneath the chert horizons (samples 177897, 177901, 177889, and 177893), however, show marked changes in geochemical composition from relatively unaltered metabasalts above and below the cherts. These changes include large positive increases in SiO2 values (from c. 47 to c. 75 wt%) and the almost complete removal of FeO (from c. 9 wt% to 0), MgO (from 5.5 to 0.2 wt%) and CaO (from between 7 to 11 wt % to 0.1 wt %) from all of the samples. The white clay-altered pillow basalt from the Dresser Formation (177897) is dominated by SiO2, Al2O3, and TiO2. The sample of Apex Basalt from adjacent to the silica vein (sample 177901; Fig. 9) has not been affected by the same type of intense alteration as the other samples collected from immediately beneath the bedded cherts, but has high SiO2 values and very low CaO due to mild quartz–chlorite alteration, resulting in a high CIA value of 0.92. The metabasalt sample collected from above the Apex chert (177900; Fig. 9)

Geochemistry of metabasalts and hydrothermal alteration zones

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Na2O in all of the samples. Sample 177897 (Dresser Formation) shows a large gain of Ba (almost 1000%) and sample 177893 is unique in showing gains of K2O and P2O5. The relative mass loss or gain of the altered samples, assuming immobile Al2O3 during alteration, is calculated by:

f1sCAl2O3o/CAl2O3Adgx100

Fig. 7. (A) Outcrop view of hydrothermal breccia (sample 155988) from the Dresser Formation, consisting of rounded, silicified sedimentary fragments (now microcrystalline quartz) with Fe-rich alteration rinds in white quartz matrix. White tip of pen at top of photo is 8 mm across. Location shown in Figure 5. (B) Cross-polarized light thin section view of A), showing microcrystalline quartz (chert) cut by a quartz vein, in turn cut by irregular kaolinite veinlets. Width of field = 4 mm.

also shows evidence of hydrothermal alteration in having depleted CaO contents relative to unaltered basalts, and a high CIA value of 0.91. Comparison of major elements and selected trace elements from altered rocks versus relatively unaltered rocks from well beneath the cherts for each of the studied localities (Table 1) follow the procedures outlined by Gresens (1967). The data columns labelled Gresens analysis in Table 1 are plots of the change in concentration of a component (Ci) relative to its concentration prior to alteration (Ci), assuming Al2O3 is immobile during alteration, following equation 6 of Grant (1986): Ci/Ci=sCAl2O3o/CAl2O3AdsCio/CiAd1 where CAl2O3o is the concentration of Al2O3 in the unaltered rock (o indicates original rock), CAl2O3A is the concentration of Al2O3 in the altered rock (A indicates altered), Cio is the concentration of the component in the unaltered rock and CiA is the concentration of the component in the altered rock. The results are expressed in decimal numbers such that 1=100% gain of an element and 1=100% loss of an element. They show an overall gain in SiO2 and loss of MnO, MgO, CaO and

The results are listed in Table 1 in the row marked “Total” under columns headed “Gresens analysis”. They show a mass loss for the altered Dresser Formation (24.1%) and Strelley Pool Chert locality 1 samples (13.5%), almost no loss or gain for the Apex chert sample, and mass gain for the Strelley Pool Chert locality 2 sample (+19%). The assumption of immobile Al2O3 during alteration is confirmed for the altered Dresser Formation metabasalt by a very low CZr/CZr of 0.073. In all of the other samples, however, CZr/CZr varies quite widely, indicating at least some mobility of what are generally accepted as “immobile elements” during alteration in these samples. The highly altered nature of the rocks from immediately beneath the chert horizons is evident from ternary A–CN–K and A–CNK–FM plots (Fig. 14A, B). Although these plots have not been corrected for carbonate and apatite as suggested by Nesbitt & Young (1982), the carbonate-altered and relatively unaltered metabasalts from the North Pole Dome (see Table 1) mostly plot within a narrow range on the A–CN tieline in Figure 14A, indicating that this omission accounts for only a minor shift in the plot of the data towards the Al2O3 apex. On this basis, we conclude that the North Pole metabasalts have not been subjected to subaerial weathering (cf. Nesbitt & Markovics 1997), consistent with the CIA data. The three samples that plot towards the CN apex of Figure 14A relative to the majority of North Pole samples, are from the distinctive group of komatiitic basalts with low Al2O3 and high CaO and MgO (samples 159632–159636: Table 1). This shift reflects the high-Mg character of these rocks and the primary mineralogical control of olivine and/or clinopyroxene (and little to no plagioclase) on bulk composition. Figure 14A and B show that rocks from immediately beneath the chert horizons plot near the A apex, indicating the near-complete removal of alkalis and Fe and Mg during alteration. Clay-altered pillow lava from the Dresser Formation (sample 177897; Fig. 6B), altered basalts from both above and below the Apex chert (samples 177900 and 177901), and the quartz–pyrophyllite schist from below the Strelley Pool 1 locality (sample 177889) all plot close to the A apex of the A–CN–K diagram, despite widely differing mineralogies (Fig. 14A) and have CIA values of 0.88 to 0.91 indicating extreme alteration by hot seawater or by later, non-seawaterderived fluids that leached all CaO and Na2O and caused slight addition of K2O and Al2O3 (Table 1). The K2O-enriched sample of altered pillow basalt from beneath the Strelley Pool Chert at locality 2 (177893) is the only sample that lies along the trend of weathering on the A–CN–K plot. Although this weathering may have occurred in recent times, rocks a similar distance stratigraphically above the subvertically dipping chert (sample 177895) do not show this K2O alteration and suggest that the weathering may have occurred during subaerial exposure of these rocks in the Archaean (Buick et al. 1995) as is characteristic of Archaean weathering profiles (Button & Tyler 1979). Figure 14B shows that the metabasalts have been affected by two distinct alteration trends. The first is a trend away from the CNK apex, reflecting the effects of stripping of alkalis by hydrothermal propylitic alteration (chlorite–carbonate–epidote).

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Fig. 8. Simplified geological map of part of the Marble Bar greenstone belt, showing the location of Schopf’s (1993) putative microfossil locality in the Apex chert. Note the swarms of weakly radiating hydrothermal amorphous silica veins that feed the Apex chert and the lensoid nature of associated bedded chert deposits. Note the control on vein swarm arrays by normal growth faults that progressively offset lithology down stratigraphic section. There are several chert beds with associated vein swarms in the Apex Basalt through c. 4 km of stratigraphic section of dominantly pillow basalts.

This is particularly evident for the altered basalts beneath and above the Apex chert (samples 177900 and 177901), which plot close to the A–FM tieline. The second trend has affected the sample of clay-altered metabasalt (saprock) in the Dresser Formation and samples from both localities of the Strelley Pool Chert. This trend, related to phyllic (quartz–sericite–pyrite), argillic (kaolinite) and/or advanced argillic (kaolinite– pyrophyllite–alunite–barite) alteration and/or surface weathering, causes the rocks to plot neat the A apex of the A–CNK–FM diagram, reflecting crystallization of illite and/or kaolinite. (The mineralogy of hydrothermal alteration styles is described in detail by Pirajno 1992; Lentz 1994). The REE systematics of highly altered rocks from immediately beneath the studied chert horizons are variable relative to those for overlying and underlying metabasalts (Fig. 15A–D). For the Dresser Formation and Apex chert localities, highly altered metabasalts immediately beneath the chert horizons have preferentially lost LREEs (Fig. 15A, B). In contrast, the

metabasalt from above the Apex chert (sample 177900) shows significant LREE enrichment and has high Ba and K2O, which may be the result of high fluid/rock ratios in a low-temperature hydrothermal alteration environment (cf. Ludden & Thompson 1978). At both of the Strelley Pool Chert localities, the underlying and overlying metabasalts have almost identical, flat REE patterns (Fig. 15C, D), whereas the immediately underlying altered rocks show markedly different REE patterns. At locality 1, the pyrophyllite–quartz–sericite schist has generally lower overall REE abundances (Fig. 15C), reflecting dilution by silica during low-temperature hydrothermal alteration that transformed this basaltic protolith into pyrophyllite–quartz–sericite schist. Unlike metabasalts of the Dresser and Apex Basalt Formations, this sample has preferentially lost heavy rare earth elements (HREE). That the schist was derived from a basaltic protolith is confirmed by its REE profile, which is different from that of altered felsic volcanic rocks (quartz–sericite) of the

Geochemistry of metabasalts and hydrothermal alteration zones

Fig. 9. Simplified geological map of the Apex chert microfossil locality, showing relationship between hydrothermal chert vein and bedded chert. Location of geochemical samples 177900–177902 indicated.

Panorama Formation (Fig. 15C). At Strelley Pool Chert locality 2, the pattern of REE alteration is different from all other localities, and is characterized by LREE enrichment and depletion of HREEs and some middle REEs (Fig. 15D), a pattern similar to that produced by advanced argillic alteration in the Rodalquilar gold alunite deposit in Spain (Arribas Jr. et al. 1995) and the Western Tharsis Cu–Au deposit, Mt Lyell, Tasmania (Huston 2001). Vein and bedded chert geochemistry. Two dark blue–black silica veins, a sample of a bedded chert and one of ferruginous gossan were analysed in order to assess what elements were added during hydrothermal alteration and to determine the mineralization potential of the hydrothermal systems (Table 1). Analyses of the silica margin and barite-rich core of a zoned vein of the Dresser Formation are compositionally distinct. The margin of the vein (sample 177898) is dominated by SiO2, has moderately high Ba and As, and depleted trace metal contents (Cu, Ni, Zn) relative to the host rocks. The barite-rich core of the vein (sample 177899) contains pyrite in addition to chert and barite and has high SO3, but low base metal contents. A sample of ferruginous gossan from bedded chert of the formation contains 67.07 wt% Fe2O3, 264 ppm Cu, 340 ppm Ni, 2318 ppm Ba, 13.8 ppm Pb and 1430 ppm Zn. At Strelley Pool Chert locality 1, a sill of black silica in layered sedimentary dolomite contains almost 100% SiO2 with slightly elevated As and Ba contents relative to stratigraphically underlying and overlying metabasalts (Table 1). Bedded chert from Strelley Pool Chert locality 2 (sample 177894) also has elevated As and high Pb (596 ppm) contents relative to stratigraphically underlying and overlying metabasalts, but low contents of other base metals, and Ni, Cr, V (Table 1). DISCUSSION Physico-chemical processes of hydrothermal alteration Dresser Formation and Apex chert. The development of alteration mineral assemblages (quartz, jasper, barite, kaolinite, alunite–

263

jarosite, sericite, chlorite, carbonate) in footwall basalts of chert horizons in the lower Warrawoona Group associated with generations of silica  barite veins provides compelling evidence for repeated episodes of seafloor hydrothermal alteration during hiatuses in mafic volcanism, involving syngenetic exhalite deposition (bedded chert horizons) from hot springs and replacement of host rocks along fluid pathways (silica veins and breccias: e.g. Barley 1984; Nijman et al. 1999; Van Kranendonk 2004). Vein emplacement into active growth faults, combined with stratigraphic evidence for chert deposition and volcanism during growth faulting, suggests that repeated episodes of hydrothermal circulation occurred during episodic development of one or more volcanic calderas (Van Kranendonk et al. 2001). The evidence for multiple phases of alternating silica- and/or barite-rich fluids in feeder veins to bedded chert–barite units in the Dresser Formation and the presence of syndepositional silica breccias with hydrothermal kaolinite support the data from alteration geochemistry studies that the clay-altered pillow basalts beneath the major chert horizons originated through steam-heated acid–sulphate alteration in a high-sulphur epithermal system. An acid–sulphate environment of alteration is supported by the fact that the altered rocks have preferentially lost LREE to solution, as occurs preferentially in low temperature acid waters (Michard 1989). The quartz–kaolinite–barite alteration is indicative of advanced argillic alteration (White & Hedenquist 1995), and these rocks may also have been affected by Archaean supergene alteration as is common in more recent epithermal systems (Pirajno 1992). The very shallow crustal level indicated by this style of alteration is supported by the evidence for very shallow water sedimentation at the top of the formation (Buick & Dunlop 1990), including desiccation cracks. Caenozoic weathering of these rocks may be a contributing, but probably minor, factor in the alteration history of these rocks, as hanging wall basalts are relatively unaltered. Passing stratigraphically downward from the Dresser Formation cherts, the alteration assemblage quartz–white mica–rutile  chlorite  epidote is characteristic of phyllic alteration (White & Hedenquist 1995). The carbonate–chlorite alteration assemblage (up to 18% LOI; e.g. samples 159635– 159639 in Table 1) further beneath the Dresser Formation and under the Apex chert is more characteristic of the propylitic alteration zone of epithermal systems (Pirajno 1992, p. 114; Lentz 1994, and references therein) than regional greenschistfacies metamorphism (e.g. Spear 1993). Considering the close association of this alteration assemblage with major hydrothermal chert horizons in the Dresser Formation, we infer that the lower part of the metabasalts beneath the Dresser Formation and Apex chert have been propylitically altered. The presence of barite, alunite–jarosite and high trace metal contents in the immediate footwall of the Apex chert (Brasier et al. 2002), together with extensive chlorite–carbonate alteration stratigraphically further down (this study), indicate a similar acid–sulphate epithermal environment of alteration to that in the Dresser Formation. The observed spatial link between bedded chert–tuff lenses, radiating swarms of underlying feeder veins, growth faults, and subvolcanic dolerite sills suggests that heat from the subvolcanic sills was at least partly the driving force for the hydrothermal circulation (Fig. 8; Van Kranendonk 2004). The lack of a steam-heated, advanced argillic alteration zone and extensive barite in the Apex chert suggests a possibly deeper level, and/or higher temperature of alteration, the former possibility supported by a lack of shallow water indicators in the sediments and presence of overlying pillow basalts. Modern analogues for the physico-chemical processes involved in the formation of the lower Warrawoona Group

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Fig. 10. Features of the Strelley Pool Chert in the North Pole Dome (parts A–D from locality 1, part E from locality 2; see Fig. 1 for location). (A) Zoned amorphous chert–quartz veins cutting bedded dolomite from locality 1. Veins are zoned from black amorphous silica outer margins to white quartz with epithermal growth zones in the cores. (B) Quartz crystal-filled dissolution cavity in silicified dolomite from locality 1. (C) Hydrothermal Fe-dolomite-altered pillow basalt from >100 m beneath the chert (sample 177892). Arrow points to Fe-dolomite infill of porous interpillow hyaloclastite breccia. (D) Lineated pyrophyllite–quartz–sericite schist immediately beneath the chert, interpreted to be hydrothermally altered mafic volcanic rocks. (E) View west of steeply dipping, well-preserved pillow basalt with devitrification varioles and chilled margin rinds from the Euro Basalt immediately above the Strelley Pool Chert (sample 177895). Stratigraphic top is to the left of the photo (south).

Geochemistry of metabasalts and hydrothermal alteration zones

265

Fig. 11. Simplified geological map of the Strelley Pool Chert at locality 2 (see Fig. 1 for location), showing zone of hydrothermal alteration beneath the chert, which faces to the south. Samples 177893–177895 come from the area where bedding in the underlying Coonterunah Group dips 72 SE. ‘A’ denotes area where chert unconformably overlies 3467 Ma granite and contains stromatolites in dolomite.

chert horizons are the DESMOS submarine caldera in the Manus back-arc basin in Papua New Guinea (Gamo et al. 1997; Gemmell et al. 1999), and the White Island volcanic complex in the Bay of Plenty, New Zealand, a submarine extension of the back-arc Taupo Volcanic Zone (Cole 1990; Brathwaite & Pirajno 1993; Hedenquist et al. 1993; Stoffers et al. 1999). Venting of low-temperature (88–120C), highly acidic hydrothermal fluids were discovered in the DESMOS submarine caldera (Gamo et al. 1997). These fluids are vented at depths of about 1900 m, through white smokers, and are characterized by being unusually sulphate-rich. Gamo et al. (1997) suggested, on the basis of isotopic data, that the high sulphate concentration might be sourced from a degassing magma beneath the caldera. This SO2-rich magmatic fluid would disproportionate to form sulphuric acid via the reaction: 4H2O+4SO2=H2S+3H2SO4 This acid is the primary agent of the advanced argillic alteration seen in these systems and drives precipitation of sulphate.

Gamo et al. (1997) concluded that the DESMOS submarine caldera is a submarine analogue of acidic hot springs of subaerial volcanism. Importantly in the context of the present work, Gamo et al. (1997) also suggested that the DESMOS hydrothermal system acts as a source of sulphate to seawater. Significantly, if this model applies to the Dresser Formation, it implies that the sulphates should have magmatic 34S signatures, as is indeed the case (Lambert et al. 1978). The hydrothermal system of White Island has alteration features of high-sulphidation epithermal systems and is controlled by acidic solutions that are responsible for alteration assemblages that include alunite, cristobalite, anhydrite, chlorite, Al-rich clays and pyrite. The hydrothermal fluids have temperatures ranging from 200 to 400C that locally debouch as hot springs and fumaroles within the caldera system (Hedenquist et al. 1993). There are numerous submarine hot springs at water depths of about 200 m, characterized by silica deposits, silicified muds, ashes and live bacterial mats (Stoffers et al. 1999). The fluids emanate from a degassing magma chamber at about

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Fig. 12. Petrographic features of hydrothermal alteration beneath the Strelley Pool Chert. (A) Plane polarized light thin section view of carbonate (white) – chlorite (light grey) vein in chlorite–epidote altered pillow basalt of the Coonterunah Group from well beneath the chert at locality 1 (sample 177892). Width of field = 4 mm. (B) Cross-polarized light thin section view of silicified pillow basalt of the Coonterunah Group from immediately underneath the chert at locality 2, showing pervasive quartz–sericite alteration, but note that original skeletal crystals (pyroxene) are still recognisable in centre of photomicrograph; sample 177893. Width of field = 4 mm. (C) Plane polarized light thin section view of well-preserved igneous texture in unaltered Euro pillow basalt from immediately above the chert at locality 1 (see Fig. 10E); sample 177890. Width of field = 4 mm. (D) Plane polarized light thin section view of euhedral zoned alunite crystals in the Strelley Pool Chert at locality 2; sample 177894. Width of field = 0.4 mm.

1.2 km below sea level, mixing and/or recharged with seawater at various stages of their evolution. A possible Archaean analogue of the lower Warrawoona Group hydrothermal systems is the geological setting of volcanogenic massive sulphide (VMS) deposits in the 3.24 Ga Sulphur Springs Group in the east Pilbara, located immediately west of the North Pole Dome (Fig. 1; Brauhart et al. 1998; Vearncombe et al. 1998; Van Kranendonk 2000). A model that schematically illustrates the environment of formation of the chert  barite horizons of the lower Warrawoona Group during hydrothermal veining and normal growth faulting in a steam-heated acid–sulphate volcanic caldera-type geothermal system is shown in Figure 16. Strelley Pool Chert. Silica–alunite alteration of the Strelley Pool Chert, and pyrophyllite–quartz–sericite and sericite–quartz mineral assemblages in alteration haloes beneath the Strelley Pool Chert indicate that silicic, advanced argillic and phyllic alteration occurred in a high-sulphidation epithermal system. The occurrence of advanced argillic alteration in this system is

supported by the HREE-depleted geochemical patterns of altered basalts beneath the chert (samples 177889 and 177893: Fig. 15C, D) that is characteristic of this style of alteration (Arribas et al. 1995). The distinctive, high K2O signature of altered footwall basalts suggests that Archaean weathering processes may have affected these rocks during the period of subaerial exposure that preceded deposition of the Strelley Pool Chert (Buick et al. 1995) as is characteristic of Archaean weathering profiles (Button & Tyler 1979). These effects were possibly enhanced by the effects of advanced argillic to phyllic hydrothermal alteration (Taylor & Fryer 1983). Field evidence shows that silicic alteration of the Strelley Pool Chert post-dated deposition of the siliciclastic and dolomitic protoliths of the formation (Lowe 1983; Van Kranendonk et al. 2003a), in contrast to the syngenetic style of hydrothermal alteration of the lower Warrawoona Group cherts. Thus, we suggest that the hydrothermal alteration associated with the Strelley Pool Chert was caused by fluid circulation beneath, and driven by the heat from eruption of, the overlying Euro Basalt as shown schematically in Figure 17.

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Geochemistry of metabasalts and hydrothermal alteration zones Table 1. Geochemical analyses of Warrawoona Group metabasalts and cherts from the east Pilbara

Sample No.

#BIR-1 average

#BIR-1 Rec

47.96 0.98 15.64 11.51

47.77 0.96 15.35 11.26

0.18 9.70 13.43 1.83 0.04 0.04

0.17 9.68 13.24 1.75 0.03 0.046

#BIR-1 average 3.64 6.88 1.87 409.07 0.03 127.45 2.54 1.74 0.53

#BIR-1 st. dev. 4.08 0.44 0.12 18.77 0.02 9.59 0.10 0.07 0.03

16.25 1.93 1.85 0.58 0.54

0.56 0.06 0.28 0.03 0.22

0.26 0.53 2.39 181.93 2.48 0.36 2.95

0.02 0.17 0.08 8.16 1.11 0.02 3.34

0.57

0.16

1.12 108.92 0.04 0.43 0.04 0.01 349.13 14.81 1.64 77.17 14.58

0.09 6.82 0.04 0.08 0.06 0.01 15.64 1.98 0.05 7.38 1.95

a

Lithology SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOIb Total FeO(t) Mg number As Ba Ce Cr Cs Cu Dy Er Eu F Ga Gd Ge Ho La Li Lu Nb Nd Ni Pb Pr Rb S Sb Sc Sm Sr Ta Tb Th U V Y Yb Zn Zr

Below Dresser Fm. North Pole Dome 159633 159632 159635

147944

159634

1 48.77 1.41 15.36 2.19 11.39 0.25 5.02 9.48 2.42 1.13 0.12 2.34 99.88 13.36 40

2 48.11 0.81 5.57 1.88 10.9 0.27 12.01 13.83 1.11 0.13 0.07 4.93 99.62 12.59 63

3 47.78 0.89 4.61 2 12.33 0.29 12.28 12.79 0.6 0.15 0.06 5.89 99.67 14.13 61

4 45.88 0.99 5.95 2.78 12.2 0.31 13.98 11.48 0.74 0.08 0.08 5.1 99.57 14.70 63

1.7 210 16.4 150 2.7 0 4.5 2.9 1.1 621 19.3 4.1 2.0 1.0 6.5 19.0 0.4 4.6 11.21 78 4.0 2.5 33 60 0.0 33 3.4 170 0.0 0.74 0.9 0.2 296 26.4 2.8 137 96

0.9 177 13.6 1476 1.1 145 2.5 1.4 1.1 445 8.7 2.9 2.0 0.5 5.2 8.5 0.2 3.3 9.98 289 2.0 2.1 6 290 0.4 39 2.7 65 0.0 0.46 0.5 0.1 193 13.6 1.2 91 49

0.9 221 10.6 1667 0.2 127 2.4 1.3 0.8 496 6.6 2.7 1.5 0.5 3.6 9.0 0.2 4.4 8.64 545 1.5 1.8 3 80 0.0 29 2.5 51 0.0 0.46 0.6 0.1 181 12.8 1.1 114 58

1.0 459 15.9 1650 0.3 232 3.0 1.6 1.1 438 9.8 3.4 1.7 0.6 5.8 14.5 0.2 5.1 11.69 557 1.5 2.5 2 170 0.2 27 3.2 53 0.0 0.55 0.7 0.2 168 15.5 1.4 124 65

Styles of mineralization Lower Warrawoona Group. Geothermal systems related to caldera-forming processes are responsible for a wide range of ore deposits and associated alteration envelopes, including polymetallic veins, epithermal gold deposits and volcanogenic massive sulphides. The chert  barite horizons of the lower Warrawoona Group are associated with an alteration mineral assemblage that is related either to venting of low-temperature fluids with high concentrations of Si, Ba, SO42 and H2S, or to distal sulphate-rich facies of hightemperature hydrothermal systems. The presence of cherts

159636

159637

4 49.07 0.59 9.6 1.03 8.86 0.21 14.18 6.53 1.28 0.04 0.06 8.19 99.64 9.79 72

4 43.23 0.47 6.86 0.55 8.29 0.25 13.06 8.89 0 0 0.05 18.01 99.66 8.78 73

5 43.93 0.65 12.57 0.49 9.4 0.16 5.59 11.49 0.49 0.28 0.05 14.79 99.89 9.84 50

1.0 504 8.7 1614 0.0 54 2.4 1.6 0.6 243 9.9 2.3 1.5 0.5 3.6 16.0 0.2 2.2 5.80 359 1.5 1.3 2 180 1.0 32 1.8 51 0.0 0.37 0.5 0.1 178 14.9 1.6 81 43

1.7 6 6.8 1934 0.2 32 2.0 1.4 0.5 352 7.4 1.8 0.8 0.4 2.8 3.0 0.2 1.8 4.66 365 4.0 1.0 2 200 0.8 35 1.4 55 0.0 0.37 0.3 0.0 172 12.3 1.3 68 36

1.6 109 4.9 305 0.5 69 1.9 1.3 0.5 648 12.7 1.7 1.3 0.4 1.7 53.5 0.2 1.5 3.80 84 1.0 0.8 11 70 1.0 48 1.3 61 0.0 0.28 0.2 0.0 250 11.9 1.3 90 36

associated with minerals such as barite, jasper and Zn, Pb and Ag (Ferguson 1999, p. 231) in the Dresser Formation is similar to the distally precipitated exhalites or chemical sediments of the Phanerozoic Kuroko-style VMS deposits (Large 1992; Sillitoe et al. 1996; Spry et al. 2000; Peter 2002). Kuroko-style deposits are well represented in the Phanerozoic (e.g. Japan; Tatsumi 1970) as well as in the Archaean (e.g. Noranda district in Canada; Gibson et al. 1993), the main difference between the deposits of different age being essentially their metal association, with the former being Pb-rich and the latter Zn-rich (Franklin et al. 1981; see also Pirajno

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Table 1. continued

Sample No. a

Lithology

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOIb Total FeO(t) Mg number As Ba Ce Cr Cs Cu Dy Er Eu F Ga Gd Ge Ho La Li Lu Nb Nd Ni Pb Pr Rb S Sb Sc Sm Sr Ta Tb Th U V Y Yb Zn Zr

159638

Within Dresser Fm. North Pole Dome 159639 177897 177898

5

4

46.22 0.56 13.26 0.51 7.51 0.23 8.02 7.43 2.95 0 0.06 13.11 99.86 7.97 64

49.59 0.9 13.96 1.75 8.59 0.19 6.98 6.11 2.83 0.01 0.08 8.93 99.92 10.1647 55

1.4 0 5.7 772 0.1 17 2.1 1.5 0.5 483 13.5 1.7 1.3 0.5 2.3 42.0 0.2 1.5 3.71 171 1.0 0.8 1 80 0.8 46 1.2 70 0.0 0.28 0.3 0.0 269 13.8 1.5 65 36

2.1 15 9.5 164 0.0 89 3.2 2.0 0.8 519 15.5 2.8 1.5 0.7 3.7 15.5 0.3 2.7 6.75 74 2.0 1.4 1 70 0.6 38 2.2 136 0.0 0.55 0.5 0.1 254 18.8 1.9 96 59

177899

White, clay-altered pillow basalt 70 1.88 19.54 0.89 0.06 0 0.15 0.08 1.18 1.35 0.06 4.41 99.63

Edge of zoned silica-barite vein 97.96 0.01 0.16 1.1 0.05 0 0.04 0.04 0.12 0.01 0.03

2.9 2926 17.18 338 2.06 33 6.25 3.85 1.29 322 24.1 4.94 0.9 1.43 5.86

29.8 3856 1.02 40 0.06 59 0.15 0.05 0 0 0.3 0.10 2 0 0.33

3 343900 0.44 0 0.05 49 1.06 0 0 0 0 0.11 0.6 0 2.76

0.65 5.1 13.56 24 1.1 2.42 41.3

0 0 0.54 15 5.3 0.1 0

0 0 0.90 102 0 0.01 0

0.5 55 4.03 81 0.3 0.91 0.7 0.21 484 27.4 3.8 19 131

1.7 4 0.1 16.4 0 0.00 0.1 0 11 0 0.05 23 2

0.1 0 1.04 2278.9 0 0.01 0 0 387 1.4 0.01 40 0

1992; Misra 2000 and references therein). In these environments, massive sulphides are associated with chemically precipitated sedimentary rocks, which may either cap the sulphide bodies or occur distally and laterally away from the sulphides (Spry et al. 2000; Peter 2002). The full range, from massive sulphides + chemical sediments to chemical sediments with no sulphides, form in a subaqueous caldera setting (Ohmoto & Takahashi 1983; Ohmoto & Skinner 1983; Halley & Roberts 1997; Callaghan 2001; Peter 2002), but the low total sulphide content of the Dresser Formation

99.68

Core of zoned silicabarite vein 47.23 0.01 0.24 0.42 0.04 0.03 0.09 0.08 0.1 0.01 0.02 0.96 69.23

Gresensc analysis 177897 vs 147944 0.128 0.048 0.68 0.996 1 0.977 0.993 0.617 0.061 0.607 21.40%

9.953 0.18 0.77

0.31 0.3

0.073

may be the result of high-level, low-temperature venting of fluids based on the evidence for very shallow crustal levels presented above (see also Barley 1984; Peter 2002). The large scale of the hydrothermal system at the Dresser Formation (c.10–15 km diameter: Van Kranendonk 2000) suggests that this horizon, or associated alteration haloes in the footwall basalts, may be prospective for low-temperature disseminated Au mineralization (e.g. Romberger 1986), such as those that form near volcanic fumaroles (Valette 1973). A potential Phanerozoic analogue is in the Jurassic Karoo

269

Geochemistry of metabasalts and hydrothermal alteration zones Table 1. continued

Sample No. a

Lithology SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOIb Total FeO(t) Mg number As Ba Ce Cr Cs Cu Dy Er Eu F Ga Gd Ge Ho La Li Lu Nb Nd Ni Pb Pr Rb S Sb Sc Sm Sr Ta Tb Th U V Y Yb Zn Zr

Above Dresser Fm. North Pole Dome 159645 159643 159647 159648

159640

159641

159644

5 43.6 1.03 8.92 1.68 10.44 0.23 8.68 9.05 0.88 0.04 0.09 15.1 99.74 11.95 56

4 49.28 1.59 12.59 4.13 8.76 0.21 4.89 5.27 3.92 0.04 0.16 8.96 99.8 12.48 41

5 46.67 2.35 12.1 3.08 10.98 0.22 4.39 7.09 3.06 0.04 0.23 9.28 99.49 13.75 36

4 43.19 1.44 13.09 0.83 10.74 0.34 3.57 10.61 1.24 1.02 0.14 13.53 99.74 11.49 36

3 47.41 1.42 13.53 1.32 10.05 0.2 5.39 6.34 3.64 0.02 0.19 10.1 99.61 11.24 46

3 51.77 1.69 13.75 1.63 8.14 0.18 4.03 5.71 4.38 0.08 0.17 8.12 99.65 9.61 43

2.3 20 13.5 189 0.2 119 3.3 1.9 1.0 627 13.2 3.4 1.8 0.7 5.4 13.0 0.2 3.4 9.31 111 3.0 2.0 2 510 2.2 31 2.8 171 0.0 0.55 0.7 0.2 256 17.9 1.8 106 71

1.7 289 22.1 16 0.1 176 4.6 2.8 1.4 674 19.4 4.8 1.1 0.9 9.0 13.5 0.4 6.3 14.44 38 4.0 3.2 2 340 0.6 21 4.3 188 0.5 0.83 1.4 0.3 269 26.1 2.7 119 135

3.4 12 24.8 0 0.3 35 7.0 4.2 1.9 923 23.5 6.7 1.6 1.5 9.9 11.0 0.6 7.3 17.29 23 1.5 3.7 2 1370 0.2 30 5.3 70 0.5 1.20 1.5 0.4 377 41.2 3.8 119 164

4.1 18 14.7 219 0.9 38 4.6 2.8 1.2 600 19.8 4.2 1.2 1.0 6.0 24.0 0.4 3.9 10.26 62 1.5 2.2 28 610 1.0 43 3.2 47 0.0 0.74 0.7 0.2 375 27.0 2.6 118 94

7.3 8 28.3 70 0.3 97 6.3 3.8 1.7 589 21.9 6.0 1.5 1.3 11.9 18.5 0.5 7.0 17.77 60 4.0 4.1 1 900 1.2 29 5.1 75 0.5 1.01 1.6 0.4 280 36.2 3.5 132 172

4.6 3 31.5 120 0.1 137 5.8 3.4 1.9 453 20.1 5.9 1.2 1.2 12.0 13.5 0.5 8.4 19.67 59 2.5 4.6 4 730 0.4 26 5.4 73 0.5 1.01 2.1 0.5 302 31.9 3.1 109 164

continental flood basalt province, where Au–barite mineralization at Springfontein (near the town of East London; Mountain 1974) formed as an epithermal deposit (possibly of the high-sulphidation type) characterized by abundant barite and sulphides, chalcedonic sinters, quartz and calcite associated with argillic and advanced argillic alteration envelopes (Pirajno, unpublished data). A similar association of disseminated Au with black laminated chalcedonic quartz in silicic altered rocks with underlying advanced to intermediate argillic, phyllic, and propylitic alteration zones, occurs in the Rodalquilhar Mine in

159649

159650

159651

2 51.36 1.05 13.77 2.43 7 0.19 5.32 9.77 3.52 0.22 0.11 4.87 99.61 9.19 51

3 51.93 0.94 13.35 1.62 7.31 0.2 6.92 8.54 3.67 0.28 0.1 4.95 99.81 8.77 58

5 48.03 0.65 12.93 0.97 7.99 0.18 8.14 7.37 2.64 0.04 0.06 10.68 99.68 8.86 62

5 47.53 0.62 12.76 0.34 7.65 0.22 6.21 7.65 0.28 1.49 0.06 15.06 99.87 7.96 58

5.5 117 22.7 186 0.2 98 4.5 2.8 1.2 332 16.2 4.3 1.5 0.9 9.7 9.0 0.4 5.9 13.49 66 2.0 3.2 7 880 0.2 33 3.7 174 0.0 0.74 1.7 0.4 235 26.6 2.6 81 118

11.6 154 20.3 366 0.6 84 4.1 2.6 1.0 336 14.9 3.9 1.7 0.9 8.9 13.0 0.4 5.5 11.88 78 2.0 2.8 8 200 0.8 32 3.2 130 0.0 0.64 1.6 0.4 218 24.6 2.5 76 110

14.1 0 8.6 945 0.3 50 2.7 1.8 0.6 313 12.8 2.3 2.0 0.6 3.6 21.0 0.3 2.3 5.32 211 2.0 1.2 3 460 7.8 42 1.7 81 0.0 0.46 0.6 0.2 252 16.8 1.8 69 54

3.5 157 11.6 428 1.9 43 2.7 1.8 0.6 337 12.9 2.4 1.5 0.6 4.9 34.5 0.3 3.3 6.84 96 3.5 1.6 51 120 2.8 42 2.0 34 0.0 0.46 0.9 0.2 251 16.2 1.9 68 73

Spain, in a caldera setting above an andesitic intrusion (Arribas et al. 1995). Strelley Pool Chert. The high-sulphidation hydrothermal system associated with the Strelley Pool Chert has a limited prospectivity for VMS deposits as it occurs in a continental setting. However, it may be prospective for disseminated Au (Romberger 1986), following the analogue with the Karoo example referred to above. Geochemical assays from the chert and altered footwall basalts are required to

270

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Table 1. continued

Sample No. Lithologya SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOIb Total FeO(t) Mg number As Ba Ce Cr Cs Cu Dy Er Eu F Ga Gd Ge Ho La Li Lu Nb Nd Ni Pb Pr Rb S Sb Sc Sm Sr Ta Tb Th U V Y Yb Zn Zr

159652

Euro Basalt North Pole Dome 159653 159654

2

2

3

50.39 0.59 13.57 1.67 7.94 0.23 8.28 7.69 4.25 0.11 0.05 5.07 99.96 9.44 61

48.75 0.56 13.08 2.02 7.27 0.22 7.33 12.83 1.8 0.03 0.05 5.9 99.97 9.09 59

50.02 0.65 14.04 0.66 7.5 0.14 8.87 5.25 0.03 0.83 0.07 11.83 100.07 8.09 66

1.6 33 5.5 333 0.3 87 2.1 1.4 0.4 453 12.1 1.8 1.7 0.5 2.1 11.5 0.2 1.4 3.99 106 0.0 0.8 3 250 1.4 37 1.3 30 0.0 0.37 0.3 0.0 225 13.0 1.4 72 33

2.0 7 5.4 346 0.1 83 2.1 1.4 0.5 413 11.5 1.8 1.6 0.5 2.1 15.0 0.2 1.3 3.80 103 0.5 0.8 1 200 0.2 42 1.3 82 0.0 0.28 0.3 0.0 224 12.9 1.4 71 29

1.7 247 9.6 329 0.7 103 2.1 1.4 0.5 240 12.6 1.9 1.3 0.5 4.1 43.0 0.2 2.1 5.99 86 1.0 1.4 31 170 0.8 51 1.6 18 0.0 0.37 0.6 0.1 275 12.4 1.5 65 56

Apex Basalt (Marble Bar greenstone belt) 177902 177901 177900

177890 3

Altered basalt at chert

45.1 0.94 13.03 0.65 9.73 0.15 4.51 10.32 0.6 0.45 0.09 13.21 98.82

64.14 0.85 13.05 2.3 8.8 0.07 3.95 0.13 0.75 0.32 0.07 4.4 98.89

2.5 161 7.92 298 0.83 81 3.49 2.19 0.82 396 14.2 2.79 1.2 0.81 2.76

9.2 390 4.65 270 0.46 90 3.04 1.84 0.8 737 13.5 2.38 1.6 0.68 1.49

5.4 1485 34.67 165 0.44 128 4.83 2.93 1.58 946 22.1 4.81 1.1 1.07 17.58

0.3 2.0 5.75 104 0.0 1.1 32.3

0.35 1.9 6.68 129 1.3 1.16 15

0.29 1.8 4.04 158 1 0.68 7.6

0.43 6.7 19.28 270 21.4 4.15 31

3.5 44 1.8 52 0.0 0.35 0.7 0.1 285 13.6 1.5 72 46

5.9 49 2.18 55.9 0 0.52 0.3 0 404 18.4 2.17 97 56

13.7 43 1.71 57.5 0 0.42 0.2 0 324 15.9 1.81 177 50

5.1 36 4.96 33 0.4 0.75 3.4 0.79 299 26.4 2.7 236 180

45.27 0.6 12.19 1.16 7.47 0.16 5.19 11.44 0.18 1.32 0.08 98.87 8.51 52 3.1 2077 8.5 408 0.4 68 2.4 1.6 0.5 354 11.2 2.1 1.1 0.6 3.3

determine if there is any potential for disseminated Au mineralization. Tectonic setting of the Warrawoona Group Regional geological data, supported by the geochemical and alteration data presented herein, show that the lower part of the east Pilbara stratigraphy (Coonterunah and Warrawoona

Overlying basalt Gresensc analysis 177901 vs 177902 49.57 0.42 1.27 0.107 17.04 2.42 2.53 15.67 0.097 0.1 0.534 4.62 0.126 0.32 0.987 0.24 0.248 1.11 0.29 0.26 0.22 5.33 98 0.15%

Altered pillow under basalt

1.419 0.41 0.095

0.46 0.2

0.109

Groups) is composed of two main successions. The lower stratigraphic succession includes the Coonterunah Group and lower part of the Warrawoona Group (Talga Talga and Salgash Subgroups) that were deposited continuously over 100 Ma from c. 3530 to 3430 Ma. The upper stratigraphic succession comprises the Kelly Subgroup of the Warrawoona Group, which is dominated by the 5–8 km thick Euro Basalt deposited from 3350 to 3325 Ma, followed by eruption of the felsic volcanic

271

Geochemistry of metabasalts and hydrothermal alteration zones Table 1. continued

Sample No. a

Lithology

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOIb Total FeO(t) Mg number As Ba Ce Cr Cs Cu Dy Er Eu F Ga Gd Ge Ho La Li Lu Nb Nd Ni Pb Pr Rb S Sb Sc Sm Sr Ta Tb Th U V Y Yb Zn Zr

177892

Strelley Pool Chert – locality 1 177889 177891

gc190697d

177890

Carbonatealtered pillow basalt

Altered schist under chert

Black silica sill

Overlying pillow basalt

47.5 0.48 13.26 0.8 7.5 0.12 7.91 7.97 0.34 0.59 0.05 12.51 99.06

79.32 0.58 15.33 0.15 0 0 0.13 0.04 0.16 0.23 0.03 3.94 99.96

99.42 0.06 0.13 0.06 0.02 0 0.05 0.07 0.1 0.01 0.02 0 99.9

45.27 0.6 12.19 1.16 7.47 0.16 5.19 11.44 0.18 1.32 0.08 13.66 98.87

3.5 146 6.64 613 1.22 7 2.45 1.7 0.44 393 11.4 1.81 1.1 0.59 2.60

5 79 3.84 249 0.16 7 0.8 0.5 0.21 153 13.1 0.73 1.5 0.17 1.27

14 100 0.00 4 0.07 6 0 0 0 0 0.3 0.00 0.9 0 0.00

3.1 2077 8.49 408 0.41 68 2.42 1.57 0.54 354 11.2 2.09 1.1 0.56 3.31

0.28 1.7 4.28 129 0.7 0.84 18.1

0.07 2.3 2.43 31 0 0.5 5

0 0 0.02 18 0 0 0

0.25 2 5.75 104 0 1.09 32.3

2.3 52 1.46 34.9 0 0.34 0.5 0.11 291 13.6 1.71 64 43

2 39 0.63 20.3 0.1 0.11 0.7 0.18 290 3.5 0.51 12 64

6.2 0 0.03 15.1 0 0.00 0 0.32 0 0 0 32 1

3.5 44 1.81 52 0 0.35 0.7 0.14 285 13.6 1.53 72 46

c

Gresens analysis 177889 vs 177892 0.44 0.045 0.838 1 1 0.986 0.996 0.593 0.663 0.481

Unaltered metabasalt below chert 49.78 0.95 13.7

0.23 6.83 10.12 2.22 0.44 0.05

13.50%

0.532 0.50 0.649

40.46 7.61

2.9 1.82 0.72

2.44 0.58

3.03

0.784

0.274 2.285 5.69

13.6

1.81 91.2 0.14 0.37 0.32 0.07 18.44 1.82 0.287

46.47

Strelley Pool Chert – locality 2 177893 177894

177895

Altered pillow basalt just under chert 80.62 0.89 11.48 0.49 0.24 0.01 0.29 0.25 0.18 3.37 0.18 1.82 99.93

Bedded chert

Overlying pillow basalt

99.44 0.06 0.1 0.08 0.03 0 0.05 0.07 0.1 0.01 0.02 0 100.03

48.65 0.6 12.06 1.13 7.58 0.17 6.59 7.75 2.21 0.02 0.08 11.43 98.6

45.9 240 21.24 36 1.19 11 1.56 0.95 0.75 300 13 1.98 0.6 0.33 9.69

195 0 0.22 0 0.03 4 0 0 0 0 0.4 0.00 0.6 0 0.00

5.6 4474 7.59 548 0.21 77 2.76 1.71 0.54 257 11.7 2.25 1.3 0.62 2.82

0.19 4.3 11.71 26 1.7 2.5 69.8

0 0 0.10 12 595.6 0.01 0

0.27 1.9 5.72 129 3.2 1.03 0

2.8 20 2.55 28.1 0.3 0.24 1.6 0.34 140 7.4 1.07 35 95

8 0 0.06 15.9 0.1 0.00 0 0 0 0 0 11 1

6.3 44 1.91 135.6 0 0.39 0.4 0 302 14.2 1.61 76 44

Gresensc analysis 177893 vs gc190697 0.93 0.118

0.948 0.949 0.971 0.903 8.14 3.296 19%

6.079 2.33

2.82 0.172

1.44

# BIR-1 is an international standard basalt; shown are average values vs. recorded values for major elements analysed by X-ray fluorescence, and average values vs. standard deviation (st. dev.) for the trace elements analysed by ICP-MS Mg number={fMgO/40.32g/sfMgO/40.32g+fFeOstd/71.85gd}100 Zero values are below detection limits a 1, Amphibolite; 2, massive flow; 3, pillow basalt; 4, dolerite/basalt; 5, carbonate-altered pillow basalt b LOI=loss on ignition c See text for discussion d Data from Green et al. (2000)

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Fig. 14. Ternary A–CN–K (A) and A–CNK–FM (B) plots of metabasalts from the study area. In A, metabasalts from below (circles), within (triangles), and above (squares) the Dresser Formation plot on the A–CN tieline and indicate no effects from recent weathering. Samples of altered metabasalt from beneath the Dresser Formation chert (177897), altered metabasalts from beneath and above the Apex chert (samples 177901 and 177900), and quartz–sericite altered metabasalt at Locality 1 of the Strelley Pool Chert (177889) plot near the Al2O3 apex and indicate extensive hydrothermal alteration. The arrow indicates modern chemical weathering trend (after Nesbitt & Marcovics 1997), at the apex of which lies the sample of altered pillow basalt from immediately beneath the Strelley Pool Chert at locality 2 (sample 177893). In B, all metabasalts from the North Pole Dome are plotted as dots and altered samples indicated by designated symbols. Metabasalts beneath the Dresser Formation are distinctly high-Mg and have not been affected by alteration. Metabasalts in and above the Dresser Formation have been affected by propylitic alteration (left to right arrow) and either intense phyllic alteration (samples plotting at kaolinite) and/or weathering (arrow pointing towards illite).

Wyman Formation (3325–3315 Ma) and intrusion of voluminous granites (3315–3307 Ma) derived from melting of c. 3460 Ma TTG (e.g. Collins 1993). The two intervals are separated by a regional deformation event followed by a hiatus in volcanic activity of 75 Ma during which the shelf-type quartzite–carbonate rocks of the Strelley Pool Chert were deposited. Lower succession. The suggestion that the lower volcanic succession formed in a mid-ocean ridge (MOR) setting (Kitajima et al. 2001; Kato & Nakamura 2003) is precluded by the fact that the c. 12 km thick basaltic pile was deposited over c. 70 Ma Fig. 13. (A) Variation in chemical index of alteration (CIA; Al2O3/Al2O3+Na2O+K2O+CaO) of metabasalts from the North Pole Dome with approximate stratigraphic height. B) Chondritenormalized (Sun & McDonough 1989) REE plot of metabasalt samples from the North Pole Dome. Fields for Ontong–Java and Caribbean oceanic plateau basalts for comparison from Mahoney et al. (1993); Révillon et al. (2000 and references therein). (C) La vs Nb plot (Gill 1981) of metabasalts from the North Pole Dome, showing N-MORB characteristics. (D) Extended chondritenormalized (Sun & McDonough 1989) trace element spidergram of North Pole metabasalts (data from Table 1) compared with modern oceanic plateau basalts (inverted solid triangle = Malaita, filled circle = ODP leg 13; data from Mahoney et al. 1993).

Geochemistry of metabasalts and hydrothermal alteration zones

273

Fig. 15. Chondrite-normalized (Sun & McDonough 1989) REE plots of unaltered and altered metabasalts from the studied chert horizons: (A) Dresser Formation; (B) Apex chert; (C) Strelley Pool Chert locality 1; (D) Strelley Pool Chert locality 2. Data from Table 1, except Panorama Formation data in C, which are from Cullers et al. (1993).

(Van Kranendonk et al. 2002; 2004; Hickman & Van Kranendonk 2004), is interbedded with felsic volcanic horizons (Hickman 1983; Van Kranendonk et al. 2002), contains sedimentary evidence for (at least occasional) shallow water deposition (e.g. Buick & Dunlop 1990), and lacks the critical elements of ocean crust stratigraphy (e.g. sheeted dykes) without any evidence for their having been removed by thrusting (Van Kranendonk 2000; Van Kranendonk et al. 2002). A setting other than MOR is also supported by the acid– sulphate white smoker style of hydrothermal system identified here, which is atypical of MOR. The low regional strain state, low metamorphic grade, and lack of significant massive sulphide deposits in the lower succession of the east Pilbara also precludes formation in a subduction/accretion arc-back-arc tectonic setting as pointed out by Green et al. (2000). Rather, the metabasalts of the lower succession contain all of the characteristic elements of deposition in an oceanic plateau setting (Kerr et al. 2000), including thick sequences of predominantly homogeneous basalt erupted under subaqueous conditions, the presence of komatiites (Glikson & Hickman 1981; Hickman 1983), and similar REE pattern and extended trace element spidergram plots compared to Phanerozoic mantle plume-related oceanic plateaux (Fig. 13B, D). Geochemical and Nd isotope evidence for contamination of the lower succession basalts by continental crust (Hamilton

et al. 1981; Gruau et al. 1987; Green et al. 2000), combined with widespread evidence of older inherited and detrital zircons (Van Kranendonk et al. 2002), indicate that the Warrawoona Group was contaminated by older, at least partly sialic crust to 3724 Ma (e.g. McNaughton et al. 1988; Thorpe et al. 1992a). The presence of several million tonnes of synvolcanic barite in the Dresser Formation (Hickman 1983), combined with the fact that syngenetic Pb in the barite has anomalously high µ values (238U/204Pb) of 9.45, indicates derivation of the Pb from highly evolved (felsic) crust, “{possibly a basement to the Warrawoona Group, that evolved with a high-µ character from at least 3700 Ma ago, and probably from much earlier than that.” (Thorpe et al. 1992b, p. 404). The presence of a sialic basement to the lower succession means that plume magmatism formed either an oceanic plateau based on a continental substrate following analogy with the Phanerozoic Kerguelen oceanic plateau (e.g. Operto & Charvis 1995; Frey et al. 2000), or a volcanic rifted marginal basin such as the Phanerozoic marginal basins flanking the North Atlantic (e.g. Fitton et al. 1995). The fact that the lower succession of the east Pilbara is unconformably overlain by a shelf-type quartzite–carbonate succession (Strelley Pool Chert), and not by oceanic crust, precludes a volcanic rifted margin setting and supports the analogy with the Kerguelen oceanic plateau. The c. 70 Ma of continuous

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Fig. 16. Schematic illustration showing a model of submarine caldera setting, the inferred tectonic environment for the hydrothermal systems of the Dresser Formation and Apex chert. Inset shows idealized alteration patterns associated with a single hydrothermal circulation event and deposition of bedded chert. Note background propylitic alteration in the basaltic footwall rocks.

Fig. 17. Schematic illustration showing a model of formation for the hydrothermal alteration of footwall basalts under the Strelley Pool Chert as a function of circulating fluids driven by heat from the overlying, newly erupted Euro Basalt.

volcanism in the lower succession is not an uncommon length of time for hotspot magmatism, as exemplified by the Iceland hotspot or Hawaiian–Emperor island chain. Phanerozoic oceanic plateaux, such as Kerguelen and Ontong Java, are poorly known in terms of their associated alteration–mineralization systems, although Révillon et al. (2000) report seafloor metamorphism to amphibolite facies in the Caribbean oceanic plateau, which accords well with observations in the Warrawoona Group (cf. Kitajima et al. 2001). However, it is unlikely that c. 20–30  106 km3 of igneous material was emplaced into the ocean without considerable regional and local convective hydrothermal flow and associated alteration and seafloor-type metamorphism. Submarine oceanic crust at mid-ocean ridges and the submarine parts of volcanic arcs and back-arcs are endowed with numerous caldera-like

systems that are the sites of hydrothermal venting and ore deposition, such as DESMOS. Under-explored oceanic plateaux would be no exception, although with a lack of associated extension, the potential for mineralization in oceanic plateaux may not be as great. There is evidence that some VMS deposits, thought to be formed in back-arc basins, may in fact be related to bimodal magmatism linked to oceanic plateaux and ultimately to mantle plumes. The Early Devonian Ashele and Koktal VMS deposits in the Altay Orogen, along the northeastern margin of the Junggar block, have been linked to hot spot volcanism (Wang 1999; 2003). More specifically, Wang (1999) suggested that these VMS deposits formed at a continental margin that was rifted during hot spot magmatism. Independent support for this idea comes from seismic tomography evidence and

Geochemistry of metabasalts and hydrothermal alteration zones shear-wave velocities (Davis 1996; Zhang 1998), which indicates that the Junggar block, now covered by a thick Mesozoic–Cenozoic sedimentary cover, may perhaps represent a buried oceanic plateau. Felsic volcanic intervals in the lower succession are restricted to thin tuffaceous horizons a few tens of metres thick and associated subvolcanic intrusions, except in the vicinity of Marble Bar where the andesitic through rhyolitic (dominantly dacite) Duffer Formation reaches a few kilometres of stratigraphic thickness around an eruptive centre (Hickman 1983; Di Marco & Lowe 1989). These felsic volcanic intervals can generally be explained in terms of fractionation from a basaltic parent (Barley et al. 1984; Di Marco & Lowe 1989), and outline a c. 15 Ma volcanic cyclicity (Van Kranendonk et al. 2001). However, the presence of voluminous tonalite–trondhjemite–granodiorite intrusions over the period 3470–3450 Ma (e.g. Bickle et al. 1993; age data summarized in Van Kranendonk et al. 2002), contemporaneous with Duffer volcanism, suggests that plume magmatism was probably accompanied by shallow subduction around the margins of the plateau (Smithies et al. 2003). Volcanism in the lower succession concluded with up to 32 Ma of rhyolitic volcanism in the Panorama Formation, from c. 3458 to 3428 Ma, the earliest part coeval with intrusion of the North Pole Monzogranite laccolith in the core of the North Pole Dome (Fig. 1: data in Van Kranendonk et al. 2002). The Panorama Formation is lithostratigraphically and geochemically distinct from the Duffer Formation, with evidence for derivation from melting of eclogite (Di Marco & Lowe 1989). Except for a small polymetallic Zn–Pb–Cu–As–Ag–Sb–Au deposit in the North Pole Dome associated with a high-level stock (Goellnicht et al. 1988), there are no mineral deposits associated with the Panorama Formation. We interpret the Panorama Formation to be the product of melting of the base of a thick (eclogite grade; Cullers et al. 1993) plateau as a result of waning heat from the cooling plume head (e.g. Campbell & Hill 1988). In the oceanic plateau model, the four panels of basalt-chert stratigraphy with downward-increasing metamorphic temperatures identified by Terabayashi et al. (2003) in the southern part of the North Pole Dome and interpreted by them as obducted slices of mid-ocean ridge basalt, are re-interpreted as successive major volcanic eruptive events capped by successive bedded cherts deposited from repeated episodes of seafloor hydrothermal circulation. Each hydrothermal circulation episode is interpreted to be driven by heat from subvolcanic intrusions associated with each of the volcanic eruptive events. The downward increase in temperature in each of the panels reflects the increasing proximity to the subvolcanic intrusions, and the restricted nature of hydrothermal alteration within the panels is the result of the preceding cherts acting as aquacludes to downward hydrothermal circulation. Similarly, the change in the composition of cherts up through the lower succession of the Marble Bar greenstone belt documented by Kato & Nakamura (2003) just to the east of the North Pole Dome, was interpreted by these authors as a tectonically stacked set of slices of a once laterally continuous oceanic crust that changed along strike from MOR, via a hotspot, to a convergent plate boundary as a result of horizontal plate motions. We interpret these changes as due to the effects of varying intensities of hydrothermal fluid circulation, distance of the deposits from the source of the hydrothermal fluids, and varying chemistries of the hydrothermal systems associated with melts derived from different parts of a mantle plume during progressive oceanic plateau formation (i.e. plume head: tholeiitic basalts and jasper; plume tail; high-Mg basalts and high Cr, Ni cherts; siliceous mudstones of the Panorama Formation; derivation from differentiated

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silicic melts resulting from melting of the base of the plateau: cf. Campbell & Hill 1988; Campbell et al. 1989). Upper succession. The upper succession conformably overlies the Strelley Pool Chert and consists of a c1.5 km thick basal unit of olivine spinifex-textured komatiite and basaltic komatiite and up to 5 km of overlying interlayered tholeiitic and high-Mg basalts. This succession was erupted in c. 25 Ma from 3350 to 3325 Ma (data in Van Kranendonk et al. 2002; D. R. Nelson, Geological Survey of Western Australia, written communication 2003), followed by voluminous, high-K2O felsic magmatism (Wyman Formation and monzogranitic plutonism) at c. 3325–3310 Ma that was derived from melting of older crust (Barley & Pickard 1999; Collins 1993; Van Kranendonk et al. 2002). The Euro Basalt is characterized by flat REE patterns with small negative Eu anomalies and has been contaminated by continental crust (Arndt et al. 2001). The komatiite–basalt association and later felsic crustal melts are the predicted consequences of an Archaean mantle plume erupted into pre-existing continental crust (Campbell et al. 1989), and we therefore interpret the upper succession as a continental flood basalt province (cf. Arndt et al. 2001).

CONCLUSIONS Geological mapping, petrographic studies, and geochemical analyses of metabasalts and alteration zones associated with several stratiform chert  barite horizons in the Warrawoona Group of the Pilbara Craton have revealed that cherts were deposited during, or affected by, repeated episodes of hydrothermal circulation in two distinct environments in different parts of the stratigraphic sequence and are separated by a regional unconformity. The lower part of the succession contains chert horizons with all the classic elements of hydrothermal silica deposits, including grey and white layered chemical sedimentary chert, locally thick sections of chert interlayered with stratiform barite, jaspilitic chert, and carbonate–chert laminates. The chert horizons are deposited above, and fed by, weakly radiating swarms of silica  barite veins. These horizons are associated with extensive hydrothermal alteration of footwall metabasalts, in contrast to overlying metabasalts that are relatively unaltered and indicate the syngenetic nature of chert deposition and alteration. Styles of footwall alteration include advanced argillic, argillic, sericitic, and propylitic alteration, similar to high-sulphidation epithermal systems in a subaqueous, caldera setting. This depositional setting and alteration style is similar to present day DESMOS submarine caldera in the Manus back-arc basin (Papua New Guinea) and the off-shore extension of the Taupo Volcanic Zone (White Island). However, based on geological and geochemical features, the geotectonic setting of the lower succession is an oceanic plateau built on, and contaminated by, a basement of older continental crust, analogous to the Phanerozoic Kerguelen oceanic plateau. The lower succession is separated from the upper part of the Warrawoona Group by a 75 Ma hiatus in volcanism and deposition of the Strelley Pool Chert. The upper succession (3350–3310 Ma Kelly Subgroup of the Warrawoona Group) represents a continental flood basalt event, the heat from which is interpreted to have caused the silicic, advanced argillic and phyllic alteration of the Strelley Pool Chert and footwall basalts. We would like to thank Dave Huston and Andrew Saunders, and the editor, Jan Peter, for helpful reviews that have made this a better contribution. Paul Morris provided constructive comments on an early version of this paper and helped with the presentation and

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understanding of the geochemical data. Published with permission of the director of the Geological Survey of Western Australia.

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