Survey
* Your assessment is very important for improving the workof artificial intelligence, which forms the content of this project
* Your assessment is very important for improving the workof artificial intelligence, which forms the content of this project
J. metamorphic Geol., 2001, 19, 583±599 Sea¯oor hydrothermal alteration at an Archaean mid-ocean ridge K . K I T A J I M A , 1 S . M A R U Y A M A , 1 S . U T S U N O MI Y A 2 , * AND J . G. LIOU 3 Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Meguro, Tokyo 152±8551, Japan ([email protected]) 2 Mineralogical Institute, Graduate School of Science, University of Tokyo, Bunkyo, Tokyo 113±0033, Japan 3 Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305, USA 1 ABSTRACT A hydrothermally metamorphosed/altered greenstone complex capped by bedded cherts exposed in the North Pole, Pilbara Carton, Western Australia, is interpreted as an accretionary complex. It is distinctive in being characterised by both duplex structure and an oceanic crust stratigraphy. This complex is shown to represent an Archaean upper oceanic crust with a mid-ocean ridge hydrothermal metamorphism that increases in grade stratigraphically downward. Three mineral zones have been de®ned; Zone A of the zeolite facies, the prehnite-pumpellyite facies or the lower-greenschist facies at high-XCO2 condition, Zone B of the greenschist facies, and Zone C of the greenschist/amphibolite transition facies. In Zone A metabasites, Ca-Al silicates including Ca-zeolites, prehnite and pumpellyite are absent and epidote/clinozoisite is extremely rare. Instead, abundant carbonates are present with chlorite suggesting high-XCO2 composition in the ¯uid. On the other hand, in Zones B and C metabasites, where Ca-amphibole+epidote/clinozoisite+chlorite+Ca-Na plagioclase are the dominant assemblages, carbonate is not identi®ed. The metamorphic conditions boundary of Zones B/C were estimated to be about 350 uC at a pressure of <0.5 kbar. Fluid compositions coexisting with Archaean greenstones at the transition between Zones B and C were estimated by thermodynamic calculation in the CaFMASCH system (T=350±370 uC, P=150±1000 bar) at XCO2 of 0.012±0.140, such values are higher than present-day vent ¯uids collected near mid-ocean ridges with low-XCO2 values, up to 0.005. The Archaean seawater depth at the mid-ocean ridge was estimated to be 1600 m at XCO2=0.06 using a depth-to-boiling point curve for a ¯uid. The carbonation due to high-XCO2 hydrothermal ¯uids occurred near the ridge-axis before or was coincident with ridge metamorphism. Key words: Archaean; greenstone; high-XCO2 ¯uid; hydrothermal metamorphism/ alteration; Pilbara Craton. INTRODUCTION Since the discovery of sea¯oor hydrothermal vents at the Galapagos Spreading Centre in 1977 (Corliss et al., 1979), a number of workers have investigated the vent ¯uid chemistry and hydrothermal alteration of oceanic crust at mid-ocean ridges (Campbell et al., 1988; Butter®eld et al., 1994; Mottl & Wheat, 1994; Alt, 1995; Von Damm et al., 1995; Charlou et al., 1996). However, the physico-chemical environment around Archaean mid-ocean ridges and the features of hydrothermally recrystallized Archaean oceanic crust have only been known from the Barberton greenstone belt (3.5±3.1 Ga) of South Africa (e.g. de Wit et al., 1982; Cloete, 1994; de Ronde et al., 1994). The geochemical features of hydrothermal systems at the sea¯oor are well-preserved in exposed oceanic crust in both Phanerozoic (e.g. Schiffman et al., 1987) and Archaean (e.g. Ohta et al., 1996) orogenic belts. Therefore, the study of *Present address: Nuclear Engineering and Radiological Sciences, College of Engineering, University of Michigan, MI 48109±2104, USA. # Blackwell Science Inc., 0263-4929/01/$15.00 Journal of Metamorphic Geology, Volume 19, Number 5, 2001 parageneses and compositions of secondary minerals in Archaean oceanic crust is essential to understanding hydrothermal alteration process at Archaean mid-oceanic ridges. The goals for the present petrological study are (1) to determine the compositions and parageneses of secondary minerals in Archaean oceanic crust; (2) to estimate P-T-¯uid compositions of hydrothermal alteration responsible for the observed parageneses; and (3) to speculate on the characteristics of Archaean seawater and its interaction with oceanic crust. The composition of modern seawater is controlled by two large ¯uxes: river discharges and interaction between seawater and oceanic basalt taking place mostly within hydrothermal circulation cells at mid-ocean ridges. The chemistry of seawater is also controlled through water-vapour circulation between hydrosphere and atmosphere. In post-Archaean times, the Earth has had large continental landmasses and the dominant ¯ux has been continental river discharges (Windley, 1995). However, Archaean seawater was predominantly buffered by water/rock interaction at the sea¯oor (Veizer, 1988). The Archaean atmosphere 583 584 K. KITAJIMA ET AL. has been considered to have had higher partial CO2 pressure than the modern atmosphere, so Archaean seawater was enriched in CO2 due to equilibration with this CO2-rich atmosphere (Owen & Cess, 1979; Walker et al., 1983; Holland, 1984; Kasting, 1987; Nisbet, 1995). One of the best examples for understanding the Archaean hydrothermal alteration processes and the physico-chemical environment around mid-ocean ridge is exposed in the Pilbara Craton, Western Australia. Awramik et al. (1983) reported wellpreserved ®lamentous-shaped microfossils in chert as the Earth's oldest fossil from the Pilbara Craton, Western Australia. They considered that this microfossil-bearing chert formed in a shallow water environment, and is representative of a photosynthetic bacteria, such as cyanobacteria. Maruyama et al. (1991), however, suggested that the Pilbara Craton, where these microfossils occur, is an Archaean accretionary complex, and the cherts may have been formed in a deep-sea environment. Furthermore, Isozaki et al. (1998) and Ueno et al. (2001) found some additional localities of fossil bacterium and suggested that the cherts were precipitated in an environment where hydrothermal ¯uid discharged, similar to the present-day mid-ocean ridges (Isozaki et al., 1998). We have selected the North Pole area of the Pilbara Craton, Western Australia for detailed petrological study because it has well-preserved, very low-grade metamorphic assemblages and continuous exposures of a sequence that is similar to modern ophiolite stratigraphies. The investigated regions are outside the contact metamorphic in¯uence from tonalite-trondjemite-granodiorite (TTG) plutons within the North Pole area. We mapped the regions at 1 : 5000 scale and numerous fresh and altered greenstones, bedded cherts, bedded and veinlet barites, silica dykes, clastic rocks and felsic lava samples were collected. GEOLO GI CAL OU TLIN E Geology of the North Pole area The North Pole area of the East Pilbara greenstone belt is located about 160 km south of Port Headland and 50 km west of Marble Bar (Fig. 1). The area is underlain by a granite-greenstone belt of the Warrawoona Group which records a subgreenschist facies metamorphism (Hickman, 1983). Bedded chert and greenstone are dominant, with minor silica dykes and veinlets or layers of barite (Fig. 2). The chert contains abundant barite and is the lowest sedimentary unit in the North Pole area (Hickman, 1973). The North Pole fault-bounded accretionary complex is composed of a series of imbricated piles of pillowed basalts (>500 m) with minor sheeted dykes and overlying bedded cherts (>30 m) that locally contain volcanoclastic or terrigenous sedimentary rocks. These Archaean sedimentary rocks and greenstones preserve an oceanic plate stratigraphy similar to those welldocumented sequences of many young circum-Paci®c accretionary complexes (e.g. the Mino-Tanba belt in south-west Japan; Matsuda & Isozaki, 1991). We divided the greenstone in this area into midocean ridge basalt (MORB) and oceanic-island basalt (OIB) types based on their mode of occurrence and associated rocks. The MORB type greenstone occurs with numerous silica dykes and one or two layers of thick-bedded (several tens of m) cherts on the top of sequence, and has been subjected to intense carbonation. The OIB type greenstone is intercalated with thin layers of (=5 m) bedded cherts, and has a few silica dykes, lacks carbonates, and is composed mainly of komatiitic basalt. The greenstone terrane at North Pole consists of three thrust-fault-bounded greenstone-chert units. The lowest unit is highly altered with numerous silica dykes and is described in the next section. The lowest and middle greenstone units belong to the MORB type, whereas the uppermost unit is the OIB type. In the selected area in the lowest unit, the deformation caused by the adamellite intrusion and its thermal overprint, which extends only hundreds of metres from the adamellite (M. Terabayashi, personal communication) are negligible (Fig. 1). Geology of the study area The study area is located to the SE of the North Pole adamellite which has been dated at 3.46 Ga (Thorpe et al., 1992) (Fig. 1). The bedded cherts have a total thickness of up to 70 m, and dip between 30 and 68 u SSE to S. (Fig. 2). The greenstones were divided into basaltic and doleritic types; most basaltic greenstones are more altered than the doleritic greenstones. Depending on the degree of alteration, greenstone varies in colour from pale green to white or light brown, as secondary minerals such as carbonate, mica and clay minerals are dominant. The bedded cherts conformably overlie the basaltic greenstones and contain no terrigenous debris. Silica dykes are interpreted to be the fossil pathway of hydrothermal ¯uids, which have circulated along normal faults in the upper oceanic crust in the vicinity of a mid-ocean ridge. The younging direction of this sequence is indicated by the shape of basaltic pillows as their tops have convex smooth surface and are always oriented orthogonal to the strike of bedding in the cherts. Hence, the chert horizons represent the original sea¯oor. The Archaean oceanic crust exposed in this area has an estimated minimum thickness of up to 1435 m. Greenstones Some basaltic greenstones preserve pillow and igneous intersertal and trachytic textures. Igneous minerals, however, are completely replaced by secondary H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR 5 85 32 32 35 57 38 20 35 76 30 60 80 20 36 30 55 35 45 32 38 30 45 30 58 North Pole Adamellite (3.46 Ga) 32 60 34 40 35 26 32 bedded chert 38 no outcrops silica dyke 59 42 42 59 25 greenstone 32 basic dyke 41 Fortescue Group < 2.8 Ga 52 36 62 12 40 low-angle reverse fault (thrust) 35 62 55 high-angle normal fault 40 40 40 representative dip & strike of bedding 52 STUDY AREA 36 29 46 30 32 38 45 32 30 40 65 48 30 48 North Pole 68 65 N 42 44 60 44 47 48 52 45 Western Australia 46 65 50 Pilbara Craton 63 40 40 47 52 45 56 80 5km 60 Fig. 1. Location map of the North Pole area in the Archaean Pilbara Craton (modi®ed after Hickman, 1981) and geological map of the North Pole area. Detailed ®eld mapping and collection of samples was completed using 1 : 5000 scale topographic maps. The study area is located in an area 5 km SE of the North Pole Adamellite intrusion. 586 K. KITAJIMA ET AL. 25 middle unit 33 30 47 55 40 42 55 56 40 42 45 64 75 75 30 68 83 42 53 50 56 4 50 80 72 38 70 61 Zone A Zone A 88 65 42 65 38 36 25 30 Zone B 80 Zone A (35 60 0˚C 65 ) Zone C bedded chert silica dyke barite basaltic greenstone N doleritic greenstone Zone C secondary minerals reverse fault (thrust) high-angle normal fault chlorite zone boundary carbonate white mica 65 dip & strike of bedding plagioclase 70 dip & strike of foliation epidote/ clinozoisite 0 500m amphibole Fig. 2. Geological map showing the distribution of secondary mineral assemblages in the study area. Samples were collected where rocks are exposed and structural data were determined. The poorly exposed areas are shown in the map without structural elements or secondary minerals. H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR minerals, i.e. plagioclase by albite, chlorite and white mica, pyroxene by chlorite and white mica, and glass by quartz, calcite and rutile. Highly altered greenstones with very ®ne-grained mixtures of quartz and sericite have no preserved textures. In particular, samples collected near the silica dykes are extensively silici®ed. Many basaltic greenstones contain various amounts of carbonate minerals including calcite, dolomite and Fe-rich dolomite which often coexist in the same sample. Carbonates and quartz occur as veins, up to 5 mm in width. Hyaloclastite is observed in only one locality and is completely replaced by chlorite and quartz. The doleritic greenstones are massive and preserve igneous holocrystalline and intergranular textures. All igneous minerals are replaced by secondary phases, except augite, which is partially replaced by Ca-amphibole along grain margins and cores of crystals. Plagioclase is replaced by albite or oligoclase, chlorite, white mica and minor epidote/clinozoisite, and Fe-Ti oxide by titanite, rutile and chlorite. Only the upper sections of doleritic greenstones contain minor thin silica dykes. Chert In the study area, bedded cherts of various colours range in thickness from 0.1 to 70 m, and are intercalated with barite layers 0.1±0.5 m thick. Similar lithologies have been reported from the Barberton greenstone belt (Paris et al., 1985). Most chert beds are discontinuous and are terminated by many latestage high-angle normal faults. These faults were formed due to the adamellite intrusion and have also cut the thrust planes separating the imbricated piles of accreted oceanic crust; these faults cut the top of bedded cherts. 5 87 sulphides, Fe-oxides and ®ne-grained zircon (<3 mm). Most dyke exposures are black and grey in colour, as they contain sulphides and oxides; some contain organic carbon (Ueno et al., 2001). MINERAL ZONES Some 300 samples were collected from the area (see sample locality in Fig. 2); all were thin-sectioned and some were analysed by EPMA, SEM-EDS, laser Raman spectroscopy, transmission electron microscopy (TEM) and analytical electron microscopy (AEM). The greenstones are divided into three mineral zones (A to C) in an ascending order of increasing metamorphic grade based on characteristic secondary minerals. The distribution of carbonate, epidote, chlorite and amphibole is shown in Fig. 2. Most Zone A greenstones contain carbonate-bearing and carbonate-free assemblages; only a few samples near the lower part of this zone contain Ca-Al silicates. Zone B is de®ned by the occurrence of actinolite and lacks carbonate-bearing assemblages. Zone C is the highest metamorphic grade, characterised by the occurrence of hornblende together with Ca-plagioclase. The mineral paragenesis of each zone is illustrated in Fig. 3. Zone A This zone is divided into two assemblages: carbonatefree and carbonate-bearing assemblages. Carbonatefree assemblages include: (A-1) quartz+white mica (A-2) chlorite+quartz+white mica (A-3) chlorite+albite+quartz+white mica (A-4) chlorite+epidote/clinozoisite+quartz+ white mica Silica dykes Silica dykes have been termed as `T-chert' (Hickman, 1973); we use the simple term of `silica dyke' rather than `quartz vein' as some are over 10 m in width and about 1000 m long and were probably formed by precipitation of hydrothermal silica along large fractures. More than 1500 silica dykes were recognized in the North Pole area; 70% are concentrated in the lowermost unit. The silica dykes cut the basaltic greenstone and are con®ned to the upper 1000 m of the greenstone sequences. Silica dykes are capped by, but do not pass through, the chert. The host greenstones are altered; the degree of alteration is positively correlated with the number of silica dykes. The silica dykes decrease in number and in width with increasing distance from the bedded cherts. They strike SE to E±W in the upper part of the sequence, and gradually change to N±S in the lower part with more gentle dips. The silica dykes are predominantly composed of ®ne-grained quartz and minor Fe, Zn, Pb and Ni Zone A Zone B Zone C albite~oligoclase albite~andesine Epidote/clinozoisite Actinolite Hornblende Plagioclase albite Chlorite Mica Carbonate calcite~dolomite Titanite Rutile Quartz Fig. 3. Schematic mineral parageneses for metabasites from the study area. Heavy lines show major phases and dashed lines show minor, or trace phases. The dashed line for rutile in Zones B and C indicates that it is completely rimmed by titanite and is not stable. 588 K. KITAJIMA ET AL. Assemblages (A-1) and (A-2) predominantly occur in the uppermost part of the greenstones at the contacts with overlying bedded cherts and silica dykes. These ®ne-grained mineral assemblages were identi®ed either as replacement of basaltic glass or primary plagioclase and pyroxene, or as mineral aggregates in veins and amygdules of altered greenstones. Assemblage (A-4) predominantly occurs in the lowermost section of Zone A near the boundary between Zones A and B (Fig. 2). Carbonate-free rocks, except (A-4), are strongly altered and contain chlorite as the dominant phase. Carbonate-bearing assemblages were also divided into four assemblages: (A-5) carbonate+chlorite+quartz+white mica (A-6) carbonate+chlorite+albite+quartz (A-7) carbonate+chlorite+albite+quartz+ white mica (A-8) carbonate+chlorite+epidote/clinozoisite+ quartz+white mica All the above assemblages with the exception of (A-4) also contain minor amounts of very ®ne-grained rutile. Carbonate-bearing assemblages have similar modes of occurrence as carbonate-free assemblages. In Zone A metabasites, carbonate-bearing assemblages are more abundant than carbonate-free assemblages; both appear in same domains and are dif®cult to differentiate. The predominant carbonates include anhedral calcite and euhedral dolomite rhombs. Some dolomite rhombs were partially to pervasively altered to hematite. Zone A metabasites contain no index Ca-Al silicate minerals for determination of the metamorphic facies and are characterised by quartz+chlorite+white mica¡carbonate¡albite¡hematite. Trace amounts of epidote/clinozoisite occur at the boundary between Zones A and B. Based on the occurrence of albitic plagioclase, we place Zone A in the lower greenschist facies or even subgreenschist facies. Zone B This zone is characterised by the occurrence of actinolite and the absence of carbonate. Primary igneous textures are preserved; except for minor relict augite rimmed by actinolite+chlorite, all primary phases are replaced by secondary minerals. Primary plagioclase is dusted with albite+quartz ¡ white mica. This zone consists of three mineral assemblages: (B-1) epidote/clinozoisite+chlorite+actinolite (B-2) epidote/clinozoisite+chlorite+actinolite+ albite +quartz (B-3) epidote/clinozoisite+chlorite+actinolite+ albite +quartz+white mica Albite in this zone contains <10 mol% anorthite content. The ubiquitous distribution of actinolite, chlorite, epidote/clinozoisite and albite indicates that Zone B belongs to the greenschist facies. Zone C This zone occupies the middle to stratigraphically lower part of the greenstones (mainly doleritic greenstone) and is characterised by two coexisting Ca-amphiboles (actinolite and hornblende). Primary igneous textures and minerals are obliterated by recrystallization. Mineral assemblages of this zone include: (C-1) chlorite+albite (C-2) epidote/clinozoisite+chlorite+albite (C-3) epidote/clinozoisite+actinolite+albite+chlorite (C-4) epidote/clinozoisite+actinolite+albite+ oligoclase+chlorite (C-5) epidote/clinozoisite+actinolite+hornblende+ albite+chlorite (C-6) epidote/clinozoisite+actinolite+hornblende+ albite+oligoclase+chlorite All Zone C assemblages contain white mica, quartz and minor titanite. The lower part of this zone is occupied by assemblages (C-5) and (C-6). These assemblages indicate that Zone C belongs to the greenschist/amphibolite transition facies. MINERAL CHEMISTRY Carbonates Carbonates in Zone A include calcite, dolomite and Fe-dolomite. Calcite is ubiquitous in Zone A carbonate-bearing samples. Dolomite becomes predominant at about 500 m from the chert/greenstone contact and crystals range in size from 0.5 to 2.5 mm. Some calcite and dolomite occur together as veinlet minerals that range in size from 0.1 to 0.2 mm across. Dolomite ranges in Fe content from 0.00 to 1.39 atoms and has <0.14 Mn atoms per formula unit (p.f.u.), based on six oxygen (Fig. 4). The Fe/(Fe+Mg) ratio (XFe) ranges from 0.00 to 0.47, with an average of 0.24. Fe-rich dolomite occurs dominantly in the upper part of Zone A. No apparent trend is shown for compositional variations of the carbonates with distance from the chert/greenstone contact, apart from the Mg content of dolomite, which increases slightly away from the chert. Calcite at 560 m (96NP322) contains relatively high Mg with an average value of 0.31 atoms. The Mn content of calcite is locally higher at 316 m (96NP573) and 776 m (96NP514) from the chert. Chlorite Chlorite is ubiquitous in all zones and replaces pyroxene, plagioclase and interstitial glass. Chlorite compositions are plotted in a Hey's diagram (Hey, 1954) assuming no ferric iron (Fig. 5). The compositional H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR 5 89 Fig. 4. Compositional variation of calcite and dolomite in Zone A greenstones as function of distance from the chert/greenstone contact. variation in terms of Al-Fe*-Mg against metamorphic grade is illustrated in Fig. 6. Zone A chlorite has a larger compositional variation and is enriched in Al2O3 compared with that in other zones; the XFe ranges from 0.35 to 0.80, and the Si content ranges from 4.96 to 5.97. The XFe and Si content of Zone B chlorite are in the range 0.23±0.33, and 5.35±6.43, respectively. The XFe of Zone C chlorite ranges from 0.23 to 0.71, while the Si content ranges from 5.09 to 6.89. Chlorite in the lowest grade, Zone A, has the highest XFe value and lowest Si. Chlorite in the North Pole basaltic greenstones shows highXFe compositions, because bulk rock composition of the basaltic greenstone has Fe-rich and Mg-poor composition compared to the doleritic greenstone. Chlorite of Zone B doleritic greenstones has Mg-rich composition. 0.9 Zone A Zone B Zone C 0.8 0.7 XFe Hole 504B Leg 83 0.6 Horokanai ophiolite 0.5 Hole 504B Leg137/140 0.4 0.3 0.2 4.5 5 5.5 6 6.5 7 Si (atoms) Fig. 5. Chemical compositions of chlorite from Zones A, B and C plotted on a Hey diagram (Hey, 1954) compared with those from the DSDP/ODP Hole 504B (Alt et al., 1982; Laverne et al., 1995) and Horokanai ophiolite (Ishizuka, 1985). The rock types of Hole 504B Legs 83 are from the pillow/dyke transition and upper dyke section, and those of Leg 137/140 are from lower dyke section. Zone A chlorite shows a wider range and higher values in XFe than those in the modern sea¯oor and Horokanai ophiolite. White mica Transparent to pale-yellow mica occurs in all three zones and replaces the plagioclase and the interstitial glass of all the observed greenstones. In basaltic greenstones, it occurs as very ®ne-grained aggregates (<1 mm). This mica ranges from muscovite or phengite to paragonite in composition. The margarite component is < 0.10, except for sample 99NP112 (j0.21). The Si content, based on 22 oxygen, ranges from 6.12 to 7.11 atoms p.f.u. and has a mean value of 6.34 in 590 K. KITAJIMA ET AL. Zone A, and it increases slightly in Zone C with a mean value of 6.59 (Fig. 7). With increasing distance from the chert in this area, white mica decreases in AlVI value at the rate of 0.60 per km, and has increased Fe and Mg contents indicative of Tschermak substitution [(Mg,Fe)Si=AlAl]. Zone A white mica has a large variation in Na/(Na+K) ranging from 0.00 to 1.00; however, Zone C white mica shows very low values (<0.30) except for one sample (97NP112), which is an end-member paragonite that occurs at 1221 m from the chert. composition. The variation of pistacite content [100Fe3+/(Fe3++Al)] is illustrated in Fig. 8, assuming no ferrous iron and a formula of 12.5 oxygen. The pistacite content for Zone A samples ranges from 26.1 to 31.0, from 6.5 to 29.8 for Zone B samples, and from 1.9 to 39.80 for Zone C samples. Although scattered, especially in Zone C, a trend toward an Al-end member with increasing metamorphic grade is apparent. Some zoned crystals in Zone C show a decrease in pistacite content from 27.0 in core to 14.7 in the rim (Fig. 8). Plagioclase Epidote/clinozoisite Epidote/clinozoisite is common in Zones B and C, but rare in Zone A and has heterogeneous Plagioclase is common in three zones of the study area. The primary calcic plagioclase in Zone A is completely replaced by albite, white mica, carbonates and rare Al Zone A Fig. 6. Chemical compositions of chlorite from Zones A, B and C plotted on Al-Fe*-Mg diagram. Fe* means total iron. Zone B Zone C Fe* Mg Zones 0 Distance from chert (m) A 500 B 1000 C Zone A Zone B Zone C 1500 6 6.2 6.4 6.6 6.8 Si 7 7.2 1.5 2 2.5 3 Al 3.5 (VI) 4 4.5 0 0.5 1 Ca 1.5 2 0 0.5 1 Na 1.5 2 0 0.5 1 K 1.5 2 0 0.2 0.4 0.6 0.8 1 Na/(Na+K) Fig. 7. Composition-distance variations of analyzed white mica from Zones A (open circles) and C (solid circles) greenstones. H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR 5 91 in Zones B and C. The rims of some brown amphibole are replaced by actinolite and green hornblende. Thus, these relict amphiboles are excluded from the chemical compositions of the amphiboles in Figs 9 and 10. In this study, hornblende is de®ned by Si content (Si<7.20) based on a formula of 23 oxygen. The Si content of metamorphic amphibole in Zone B greenstones ranges from 7.26 to 7.92, and 6.18±8.00 in Zone C. The Fe2+/(Fe2++Mg) ratio of amphibole from Zone B ranges from 0.18 to 0.38, with an average value 0.26. Zone C amphibole shows a much large Fe2+/(Fe2++Mg) ratio, ranging from 0.11 to 0.91. Based on the Ca-amphibole discrimination diagram of Leake et al. (1997), Zone B amphibole is either an actinolite or magnesiohornblende, whereas those in Zone C are actinolite, magnesiohornblende and ferrohornblende (Fig. 9). epidote/clinozoisite. The anorthite content (An%) of secondary Na-plagioclase reaches 5.6 (albite) in Zone A and ranges from 0.9 to 14.0 (albite to oligoclase) in Zone B and 0.2±18.3 (albite to oligoclase) in Zone C (Fig. 8). The most calcic plagioclase (An%=18.3) occurs furthest from the chert (1433 m). Plagioclase in these zones does not show a clear peristerite gap, but an increase in An% with distance from the chert at the rate of 4.7 An% per km is apparent (Fig. 8). Calcic amphibole Calcic amphibole is common in Zones B and C, and occurs as a replacement of clinopyroxene. Colourless to pale green actinolite occurs in Zones B and C. Pale green to green hornblende occurs only in Zone C. In addition, minor igneous Ti-rich amphibole is preserved Zones 0 a b Distance from chert (m) A 500 B core 1000 rim core C Zone A Zone B Zone C 1500 0 Fig. 8. Composition-distance variation of analyzed epidote/clinozoisite (a) and plagioclase in (b). 5 10 15 20 25 30 35 40 0 5 10 15 XMg 25 30 35 An% pistacite content (%) 1.0 0.9 20 tremolite actinolite magnesiohornblende 0.5 ferrohornblende Fig. 9. Chemical compositions of Ca-amphibole from Zones B and C greenstones plotted onto the amphibole classi®cation diagram of Leake et al. (1997). Zone B Zone C ferroactinolite 0 8 7.5 6.5 Si (atoms) 40 592 K. KITAJIMA ET AL. Zones 0 Zone B Zone C Distance from chert (m) A 500 B 1000 C 1500 6 6.5 7.0 Si 7.5 8 0 0.2 0.4 Al 0.6 (VI) 0.8 1 0 0.2 0.4 0.6 0.8 1 0 0.2 0.4 XMg 0.6 0.8 1 0 0.1 0.2 Na 0.3 Ti 0.4 0.5 0 0.2 0.4 0.6 0.8 1.0 1.2 3+ Fe VI Fig. 10. Composition-distance variation of Ca-amphibole from Zones B and C greenstones. Al , Na and Ti ranges of Ca-amphibole increase with distance whereas Si and Fe3+ contents decrease with distance. The Fe3+ content of amphibole was estimated using the method of Terabayashi (1993) for all examples. The maximum Na content of amphibole increases with increasing metamorphic grade due to the edenite substitution; it ranges from 0.03 to 0.49 in Zone B, and from 0.01 to 0.81 in Zone C. Amphibole in Zone C contains higher Ti than those in Zone B; the maximum Ti content increases with increasing metamorphic grade. The amphibole Si content shows no clear compositional gap, although Zone C amphibole shows lower Si content than Zone B amphibole (Fig. 10). Similar trends have been reported from other metamorphic belts such as the Mt. Menzies section and the Elk River section of the Karmutsen metabasites in Vancouver Island of Canada (Terabayashi, 1993). The maximum values of AlVI, Na and Ti increase with distance from the chert. The Na and Ti contents show maximum values at 1221 m and 1433 m, respectively. The Fe3+ content decreases with distance from the chert and shows a maximum value at 672 m in Zone B (Fig. 10). Rutile, identi®ed by EPMA, EDS and laser Raman, occurs in all zones with very different grain sizes. Very ®ne-grained TiO2 was analysed by AEM and selected area electron diffraction (SAED/ Fig. 12a), allowing it to be positively identi®ed as rutile. In Zone A greenstones, rutile occurs with ®ne-grained (several micron) calcite, quartz and minor chlorite (Fig. 12b). Zones B and C rutile is included by titanite and is not in direct contact with surrounding minerals such as Ca-amphibole and epidote/clinozoisite (Fig. 12c), hence rutile may not be in equilibrium with them. Small aggregates of rutile in altered greenstone have been reported from Onverwacht Group (3.5±3.4 Ga), Barberton greenstone belt, South Africa (Hanor & Duchac, 1990). Titanite DISCUSSION Titanite is a common mineral in Zones B and C, but it does not occur in Zone A. It occurs as ®ne-grained crystals in Zone C basaltic greenstones, and rims rutile in Zones B and C doleritic greenstones. Titanite compositions are illustrated in terms of Al-Ti/10-Fe in Fig. 11. Zone B titanite is depleted in Al and Fe and enriched in Ti, while in Zone C it shows a wide variation in Al. Rutile Timing of the carbonation Carbonation of greenstone is a common feature in Archaean greenstone belts. The nature and timing of carbonation are different in each greenstone belt and H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR provide information about the changes in the physicochemical environment during the alteration process. Carbonate-bearing assemblages of Archaean greenstones re¯ect the composition of the ancient hydrothermal ¯uid evolved from circulated palaeo-seawater. The North Pole greenstones have been subjected to intense carbonation; 40% of collected samples contain carbonate-bearing assemblages. The timing Al basaltic greenstone doleritic greenstone 5 93 of carbonation of the North Pole greenstones is estimated by ®eld relations and mineral parageneses. The carbonate-bearing assemblages are discontinuous in areal distribution and are bounded by thrust planes between units and high-angle normal faults due to adamellite intrusion. Such occurrences suggest that carbonation occurred on the sea¯oor prior to the accretion of the greenstone units to the continental margin. Furthermore, carbonation is recognized in greenstones adjacent to silica dykes that were formed during the precipitation of overlying chert near the mid-ocean ridge-axis (Isozaki et al., 1998). All these observations suggest that the carbonation of the North Pole greenstones was a result of alteration with highXCO2 hydrothermal ¯uid and occurred near the ridgeaxis before or coincident with the ridge metamorphism. Comparison with ophiolites and modern sea¯oor rocks Ti/10 Fe Al Zone B Zone C 4 1 3 2 Ti/10 Fe 1. prehnite-pumpellyite facies 2. pumpellyite-actinolite facies 3. greenschist to amphibolite facies 4. DSDP/ODP Hole 504B Leg 137/140 Fig. 11. Chemical compositions of titanite from Zones B and C greenstones compared with those from the prehnitepumpellyite, pumpellyite-actinolite (NystreoÈm, 1983), greenschist/amphibolite transition facies (Liou & Ernst, 1979) and 504B Leg 137/140 (Laverne et al., 1995). We compare the paragenetic sequence of secondary minerals for the Archaean hydrothermally altered upper oceanic crust described above with those from the modern sea¯oor, using examples from the DSDP/ ODP Hole 504B, and some on-land ophiolites. Petrological and metamorphic transitions with depth of the oceanic crust in these tectonic settings are shown in Fig. 13. The stratigraphic sequence is similar among these three settings. The topmost section of the oceanic crust is capped by pelagic sediments, which conformably overlie pillowed basalt; coarser-grained doleritic rocks occur in the deeper parts. In most ophiolite sequences and the present-day oceanic crust in the Carlsberg ridge and Mid-Atlantic ridge, gabbroic rock occurs at about 3 km depth. The overlying basaltic section grades to doleritic rocks at about 600, 700 and 800 m from the top of the oceanic crust in Hole 504B, Oman and the North Pole, respectively. The metamorphic facies of modern oceanic crust grades from the zeolite and prehnite-pumpellyite facies, through greenschist, greenschist/amphibolite transition facies to amphibolite facies with increasing depth. Similarly, the metamorphic facies for the ophiolite sequences shown in Fig. 14 change from the zeolite and prehnite-pumpellyite facies through the greenschist to amphibolite facies with depth while lower gabbroic rocks are largely unmetamorphosed. The metamorphic features observed in the North Pole greenstones are signi®cantly different from those in modern sea¯oor and typical ophiolite sections for the upper part of the sequence. In the North Pole, the lowest grade metamorphic zone has carbonate-chlorite assemblages instead of Ca-Al silicates including Ca-zeolites, prehnite and epidote/clinozoisite that are common in ophiolites and modern sea¯oor rocks (Fig. 13). There are also differences in compositions of secondary minerals between North Pole and modern sea¯oor rocks and ophiolites. The XFe and Si contents in Zones B and C chlorite overlap those of greenschist or greenschist/amphibolite transition facies chlorite in the 594 K. KITAJIMA ET AL. Leg 137/140, ODP/DSDP Hole 504B (Fig. 5). The differences in the chlorite Al2O3 content of the North Pole Zones A, B and C cannot be explained by progressive metamorphism with increasing temperature, as seen in other metamorphic zones (e.g. the Horokanai ophiolite in the Kamuikotan Zone, Hokkaido, Japan, Ishizuka, 1985). Zone A chlorite shows Fe-rich and Si-poor composition as compared with the Hole 504B Leg 83 chlorite which belongs to subgreenschist facies (Fig. 5). The large compositional variation of chlorite, particularly those in Zone A, in part is due to their occurrence in different domains of carbonate-free or carbonate-bearing assemblages, and in part is due to their modes of occurrence as replacement after plagioclase, pyroxene or glass. Furthermore, chlorite composition re¯ects the Fe/Mg ratio of hydrothermal ¯uid and temperature (Saccocia & Seyfried, 1994). In the study area, Na-rich mica occurs at 1221 m from the chert. Na-rich micas have been reported from the Trans-Atlantic Geotraverse (TAG) active hydrothermal mound and are considered to re¯ect high Na/K ratio of hydrothermal ¯uids (Honnorez et al., 1998). Epidote/clinozoisite, plagioclase and amphibole compositions show a systematic change with metamorphic grade. Epidote zoning and variation of pistacite content indicates that the progressive reaction Fe3+rich epidote+chlorite+actinolite assemblage is temperature dependent and the pistacite content of epidote/ clinozoisite decreases with increasing metamorphic grade (e.g. the Horokanai ophiolite, Ishizuka, 1985). The increase of An% in North Pole plagioclase is consistent with that reported for metamorphic plagioclase from the low-pressure metamorphic facies series (e.g. Maruyama et al., 1982). The North Pole titanite compositions overlap with those of East Taiwan ophiolite, which belongs to the greenschist to amphibolite facies (Liou & Ernst, 1979) and 504B Leg 137/140 (Laverne et al., 1995). Some Zone C titanite isdepleted in Al and Fe, and enriched in Ti compared to those from the prehnite-pumpellyite facies and pumpellyite-actinolite facies in Central Sweden (NystreoÈm, 1983). Conditions of the Archaean hydrothermal alteration Fig. 12. (a) A ®ne-grained TiO2 phase identi®ed as rutile by SAED from [1Å01]. (b) Back-scattered electron image of ®negrained minerals in Zone A greenstone. (c) Back-scattered electron image of rimmed titanite around rutile in Zone B. Light coloured needles in lower part of Fig. 12 (c) consist of titanite and rutile which have replaced exsolution lamella of ilmenite. The dark coloured part around the needles comprises quartz and chlorite replacing magnetite. The boundary between Zones B and C is de®ned by the occurrence of hornblende together with calcic plagioclase for metabasites, hence it corresponds to the transition from the greenschist to amphibolite facies. The metamorphic temperature of this boundary has been estimated at 350 uC for the Semail ophiolite of Oman using oxygen isotope systematics (Gregory & Taylor, 1981), whereas Suzuki (1975) estimated it at 370 uC in a contact metamorphic aureole in Kasuga, Japan using a calcite-dolomite geothermometer. These values suggest that the metamorphic temperature at the boundary of the North Pole greenschist faciesgreenschist/amphibolite transition facies (Zones B/C) to be about 350±370 uC. This boundary is mapped as H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR 5 95 Fig. 13. Simpli®ed lithological and metamorphic sections showing variation in metamorphic facies for representative sea¯oor section and several ophiolite sequences, and the North Pole, modi®ed after Ishizuka & Suzuki (1995). Abbreviations used are: Z: zeolite facies; PP: prehnite-pumpellyite facies; GS: greenschist facies; T: greenschist/ amphibolite transition facies; AM: amphibolite facies; UM: unmetamorphosed; A: Zone A; B: Zone B; and C: Zone C. 450 500 bar Zones B/C A SiO2 epidote/ clinozoisite 400 bulk 350 Am p+E calc ite+ p+CO 2+ chl H orit e+q 2O ua rt z T(˚C) C bulk chlorite amphibole F titanite (2) (1) CaO 22 e++CCOO +calcite tahneitne stip rtz +qua rutile GAP Zone A SiO2 quartz bulk calcite A 300 TiO2 CaO rutile TiO2 epidote/ clinozoisite bulk chlorite calcite C 250 0 0.1 F 0.2 0.3 0.4 0.5 XCO2 Fig. 14. T-XCO2 section for the CaFMASCH and CaTiSCH systems at 500 bar. Phase relations were calculated in the CaFASCH system, with quartz in excess for actinolite, clinozoisite, daphnite (Fe-chlorite), anorthite and calcite and for tremolite, clinozoisite, clinochlore (Mg-chlorite), anorthite and dolomite in the CaMASCH system. Solid line and dashed lines show equilibria for reactions (1) and (2), respectively. The shaded area shows Zone A condition at 350±370 uC, and oblique line area shows Zones B and C conditions. The ternary diagrams show mineral assemblage and average basalt composition in the CaTiSCH (SiO2-CaO-TiO2) and CaFMASCH (ACF) systems. subparallel to the chert bedding suggesting that the isothermal lines of sea¯oor metamorphism were subparallel to the sea¯oor and that metamorphic temperature increased with increasing distance from the chert. This feature is consistent with the modern sea¯oor metamorphism, hence this area is interpreted as having preserved primary metamorphism at the Archaean sea¯oor. The boundaries of Zones A/B and A/C are not parallel to the chert bedding, and highly oblique to the Zones B/C boundary. Zone B assemblages are restricted to the western part of the study area and are absent in the eastern part, where Zone A is widespread. It is to be expected that Zone B should be present in the eastern part of the area, if the geothermal gradient and ¯uid composition were similar for both areas. However, the disparity in the distribution of the zone boundaries between the western and eastern part can be explained by a difference in the XCO2 of the ¯uid that interacted with the greenstones in the two areas. The change in mineral assemblages from Zone A, with representative assemblage calcite+chlorite+quartz, to Zone B, with actinolite+epidote/clinozoisite+ chlorite+albite assemblage, may be related to a reaction in the CaO-FeO-MgO-Al2O3-SiO2-CO2-H2O (CaFMASCH) system for metabasites. In the presence of excess quartz, the reaction for the Zones A/B and A/C boundaries in the eastern part is written as: 3 amphibole+2 clinozoisite+10 CO2+8 H2O (1) =10 calcite+3 chlorite+21 quartz On the other hand, Zone A carbonate-bearing assemblages contain rutile whereas both Zones B and C assemblages have titanite as an equilibrium Ti-phase, in addition to other Ca-Al hydrosilicates described in the previous sections. The difference in Ti-bearing phase assemblages can be attributed to the following reaction in the CaO-TiO2-SiO2-CO2-H2O (CaTiSCH) system: titanite+CO2=rutile+calcite+quartz (2) Both reactions (1) and (2) are a function of temperature, pressure and ¯uid composition, and their temperatures can be estimated by intersection of these zone boundaries. The boundaries of Zones A/B and B/C intersect at about 1000 m from the chert. The hornblende±in boundary of Zones B/C occurs at a temperature of j350 uC, which can be considered as an isotherm. Phase relations for reactions (1) and (2) for the CaFMASCH and CaTiSCH systems were 596 K. KITAJIMA ET AL. calculated using the THERMOCALC program (Powell & Holland, 1988); the results are described below. CaFMASCH system Solid-solution minerals such as amphibole and chlorite are assumed to have no Tschermak substitution and no ferric iron; phase relations were calculated using Fe-Mg ideal mixing. Non-ideal mixing in CO2-H2O ¯uids was assumed following the approach of Powell & Holland (1985). An isobaric T±XCO2 section in the CaFMASCH system with excess quartz was calculated at 500 bar for temperatures ranging from 250 to 450 uC (Fig. 14). The section shows that the Zones B and C assemblage amphibole+epidote+chlorite is restricted to extremely low-XCO2 conditions; with increasing temperature, this assemblage extends towards a higher XCO2 condition. With increase in pressure, reaction (1) shifts toward more CO2-poor ¯uid compositions. and C assemblages should be stable in equilibrium with aqueous-dominated ¯uids. It appears that phase separation of circulating hydrothermal ¯uid may have occurred while it rose to the sea¯oor. Fluid inclusions in barite from same area show that various liquid/vapour ratios were developed due to boiling or phase separation (Rankin, 1978; Nijman et al., 1998). This would generate CO2-rich and H2O-rich ¯uids at shallower depths. In Zone A, carbonate-free assemblages and carbonate-bearing assemblages equilibrated with H2O-rich and CO2-rich ¯uid, respectively. To estimate the ¯uid XCO2 during hydrothermal alteration, it is necessary to use the pre-separation, homogenous ¯uid phase, which was estimated using reaction (1). Reaction (2) could not be used as it yields a minimum XCO2 of the CO2-rich ¯uid. Two isothermal P±XCO2 sections (350 & 370 uC) for reaction (1) over a pressure interval of 150±1000 bar are shown in Fig. 15. This pressure range equates with the hydrostatic CaTiSCH system 9000 800 7000 rtar au +q it e te rch lor cih t e+ c alc c te + e pi P (bar) i b ole d ote 3000 tz C e+ isit ozo /clin 400 5000 SEAWATER DEPTH at Archaean MOR (m) ca a m ph 600 O CO2 concentration of Archaean hydrothermal ¯uid 2+ H 2O At an estimated temperature of 350±370 uC, Zone A greenstones with the calcite+rutile+quartz assemblage are in equilibrium with high-XCO2 ¯uids, whereas Zones B and C with titanite and with the assemblage epidote/clinozoisite+actinolite are associated with low-XCO2 ¯uids. At 500 bar and 350 uC, there is a large gap from 0.030 to 0.193 in XCO2 of the ¯uid (Fig. 14) at the boundaries of Zones A/B and A/C. Thus the observed disposition of mineral assemblages in the three zones for the Archaean greenstones from the study area indicate the in¯uence of ¯uid XCO2 on the development of secondary mineral parageneses. A signi®cantly higher XCO2 (e.g. XCO2i0.193 at 350 uC, 500 bar) of hydrothermal ¯uid is necessary to stabilize calcite-quartz-chlorite assemblages in the lower part of Zone A greenstones. In contrast, Zones B 370ûC 350˚C In Zones B and C, titanite is ubiquitous and rims around relict rutile (Fig. 12c). This difference can be interpreted by reaction (2). The T±XCO2 relations for reaction (2) were calculated at P=500 bar (Fig. 14). At high-XCO2 condition, titanite is unstable with respect to the assemblage rutile+calcite+quartz at constant temperatures. This suggests that Zones A, B and C greenstones located at the same distance from the chert, at the region of intersection of Zones A/B or A/C and Zones B/C boundaries, were subjected to nearly isothermal metamorphism (c. 350 uC). In Zones B and C, relict rutile rimmed by titanite coexisting with greenschist and greenschist/amphibolite transition facies minerals (Fig. 12c). This occurrence may suggest that these rocks had initially equilibrated with higher XCO2 ¯uid in an earlier metamorphic/alteration stage. CaFMASCH system 1000 2 CO OH2 200 X 1000 0.012 ~ 0.140 0 0 0.05 0.1 0.15 XCO2 Fig. 15. P-XCO2 diagram shows two isothermal sections for reaction (1) at 350 and 370uC for the CaFMASCH system. Dotted line shows a depth-to-boiling curve for binary H2O-CO2 ¯uid at 350uC (Takenouchi & Kennedy, 1964). Intersection (X) of reaction (1) at 350uC and depth-to-boiling curve indicates pressure = 260 bar and XCO2=0.06. Arrow indicates that maximum XCO2 value of present-day sea¯oor hydrothermal vent ¯uid (XCO2=0.005). H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR 5 97 composition by using reaction (1) with a depth-toboiling point curve of H2O±CO2 ¯uid shown in Fig. 15 (Takenouchi & Kennedy, 1964). At 350 uC these two reactions intersect at P=260 bar and XCO2=0.06 suggesting that the North Pole seawater depth was about 1600 m and is similar to modern mid-ocean ridges. pressure (seawater depth) developed at modern midocean ridges. At the interpreted condition (Ptotal= P¯uid=150±1000 bar and T=350±370 uC), reaction (1) buffers the ¯uid XCO2 in the range 0.012±0.140. The maximum XCO2 value of modern mid-ocean ridge hydrothermal vent ¯uid is 0.005; subgreenschist facies assemblages are common in drilled and dredged metabasalt samples (Alt et al., 1986; Von Damm et al., 1985). This value cannot be compared with those deduced from the present study for the Archaean hydrothermal ¯uid directly, because they were estimated at different P-T conditions. However, the hydrothermal ¯uid for the North Pole greenstones is estimated to have 12 times higher CO2 concentration than present-day hydrothermal ¯uid. Hence, Ca-Al silicates such as Ca-zeolite, prehnite, epidote/clinozoisite and titanite are not stable in Archaean upper oceanic crust. If Archaean ¯uid ranges in XCO2 from 0.012 to 0.140 and has undergone subsea¯oor phase separation or boiling at 1 km depth from sea¯oor, the pressure range for this depth is estimated to be 180±370 bar (Fig. 15). This pressure range requires an overlying minimum water column of 1800 m, represented by a seawater depth of 800 m and a water column from the sea¯oor to the depth of phase separation/boiling at 1000 m. Furthermore, we constrain the pressure and ¯uid CONCLUSIONS A schematic model for the Archaean oceanic crust consisting of gabbro dolerite dyke sequence and pillow basalt capped by bedded chert is shown in Fig. 16. Such layered oceanic crust created at the spreading axis has been subjected to ridge-hydrothermal alteration similar to the modern ocean crust. The shallow transition depth for the greenschist to amphibolite facies suggests a higher geothermal gradient for Archaean ocean-ridges (see Fig. 13). Figure 16(b) shows a zonal distribution of hydrothermal metamorphism/alteration of the Archaean oceanic crust. The formation of a heterogeneous ¯uid as a result of phase separation of an ascending homogeneous ¯uid at depths of about 1000 m and 350 uC, plus the disposition of Zones A, B and C assemblages with a few representative silica dykes formed along fractures for hydrothermal vents are also illustrated. The 3.5 Ga greenstone sequence (b) precipitation 0 Archaean Mid-Ocean Ridge sea level RIDGE AXIS seawater depth 1600 m (b) bedded chert magma chamber id c la 500 1000 Zone A phase separation or boiling Zone A PHASE SEPARATION or BOILING Zone B Zone B 350˚C Low XCO2 stable assemblage and mineral BASALT DOLERITE calcite + rutile + quartz Zone C u fl Zone A H2O-rich fluid and CO2-rich fluid (XCO2 > 0.06) homogeneous fluid n u c ir CRUST High XCO2 stable assemblages calcite + chlorite + quartz o ti OCEANIC Distance from seafloor (m) (a) bedded chert hert NATURE OF FLUID heterogeneous fluid hydrothermal vents 1500 low-XCO2 fluid (XCO2 < 0.06) amphibole + epidote/clinozoisite Zone C titanite GABBRO Fig. 16. Schematic cross-section of the suggested Archaean mid-ocean ridge. (a) Hydrothermal ¯uid circulates in the upper oceanic crust which consists of bedded chert, basalt, dolerite and gabbro. (b) Enlarged view showing detailed cross section of Archaean uppermost oceanic crust with respect to the nature of hydrothermal ¯uid. Circulating ¯uid rises along normal faults and fractures, and is subjected to phase separation or boiling at 350 uC and roughly 1000 m distant from the chert. Phase separation or boiling resulted in H2O-rich and CO2-rich ¯uids in the upper part, where high-XCO2 assemblages are stable. Thick-bedded chert precipitated from vented hydrothermal ¯uid at the top of the greenstone sequence. 598 K. KITAJIMA ET AL. from the North Pole Pilbara Craton, Western Australia preserves Archaean mid-ocean ridge hydrothermal alteration. The upper 1000 m of the greenstone was subjected to carbonation by high-XCO2 hydrothermal ¯uid with an estimated XCO2 of 0.012±0.140. Ca-Al silicates, common in present-day oceanic crust and ophiolites, do not occur in the Archaean upper oceanic crust. During the ascending path, a high-XCO2 ¯uid may have separated into H2O- and CO2-rich phases at depths of about 1000 m below the sea¯oor. Interaction of upper oceanic crust with discharged H2O ¯uid would result in chlorite-white mica-quartz assemblages (carbonate-free assemblages of Zone A), whereas those with CO2-rich ¯uid develop carbonate-bearing assemblages. Because of the dynamic processes with continuous change in temperature as well as ¯uid composition during the migration of the oceanic crust away from the spreading ridge, the parageneses and compositions of secondary minerals in both basaltic and doleritic sections are extremely complicated. The schematic model for the Archaean mid-ocean ridge shown in Fig. 16 is provisional. Nevertheless, our observations for metamorphism/alteration of the Archaean oceanic crust indicate that the Archaean hydrothermal ¯uid was higher in XCO2 compared to its modern analogue. The North Pole carbonation occurred near the ridge-axis before or coincident with ridge metamorphism. ACKNOWLEDGEMENT S We thank Y. Isozaki, M. Terabayashi, Y. Kato, T. Kabashima, Y. Ueno and K. Nakamura for assistance during the ®eldwork. Collaboration in the ®eld with A. Thorne and A. H. Hickman was helpful. This research was supported by the Ministry of Culture and Education of Japan (International Scienti®c Research Program; Field research nos. 06041038 and 08041102 and Intensi®ed Study Area Program, no. 259 1995±97). Preparation of the manuscript was completed while Liou was taking sabbatical supported by the Tokyo Institute of Technology and the NSF Center of Global Partnership Fellowship, NSF INT-9820171. We thank Drs C. E. J. de Ronde, J. C. Alt and D. Robinson for critical reviews that materially improved the manuscript. R E FE R E N C E S Alt, J. C., 1995. Subsea¯oor Processes in Mid-Oceanic Ridge Hydrothermal Systems. In: Sea¯oor Hydrothermal Systems: Physical, Chemical, Biological Interactions Geophysical Monographs, 91 (eds Humphris, S. E., Zierenberg, R. A., Mullineaux, L. S. & Thomson, R. E.), pp. 85±114. American Geophysical Union, Washington DC. Alt, J. C., Honnorez, J., Laverne, C. & Emmermann, R., 1986. Hydrothermal Alteration of a 1km section through the upper oceanic crust, Deep Sea Drilling Project Hole 504B: mineralogy, chemistry, and seawater±basalt interactions. Journal of Geophysical Research, 91, 10 309±10 335. Alt, J. C., Laverne, C. & Muehlenbachs, K., 1982. Alteration of the upper oceanic crust: mineralogy and processes in Deep Sea Drilling Project Hole 504B, Leg 83. Initial Reports of the Deep Sea Drilling Project, 83, 217±247. Awramik, S. M., Schopf, J. W. & Walter, M. R., 1983. Filamentous fossil bacteria from the Archean of Western Australia. Precambrian Research, 20, 357±374. Butter®eld, D. A., McDuff, R. E., Mottl, M. J., Lilley, M. D., Lupton, J. E. & Massoth, G. J., 1994. Gradients in the composition of hydrothermal ¯uids from the Endeavour segment vent ¯uid: Phase separation and brine loss. Journal of Geophysical Research, 99, 9561±9583. Campbell, A. C., Bowers, T. S., Measures, C. I., Falkner, K. K., Khadem, M. & Edmond, J. M., 1988. A time series of vent ¯uid compositions from 21 u, East Paci®c Rise (1979, 1981, 1985) and the Guaymas Basin, Gulf of California (1982, 1985). Journal of Geophysical Research, 93, 4537±4549. Charlou, J. L., Fouquet, Y., Donval, J. P., Auzende, J. M., Jean-Baptiste, P., Stievenard, M. & Michel, S., 1996. Mineral and gas chemistry of hydrothermal ¯uids on an ultrafast spreading ridge: East Paci®c Rise, 17 u to 19 uS (Naudur cruise, 1993) phase separation processes controlled by volcanic and tectonic activity. Journal of Geophysical Research, 101, 15 899±15 919. Cloete, M., 1994. Aspects of volcanism and metamorphism of the Onverwacht group lavas in the south-western portion of the Barberton greenstone belt. PhD Thesis, University of Witwatersrand, Johannesburg. Corliss, J. B., et al. 1979. Submarine thermal springs on the Galapagos Rift. Science, 203, 1073±1083. Gregory, R. T. & Taylor, H. P., 1981. An Oxygen isotope pro®le in a section of Cretaceous oceanic crust, Semail ophiolite, Oman. Journal of Geophysical Research, 86, 2737±2755. Hanor, J. S. & Duchac, K. C., 1990. Isovolume tric silici®cation of Early Archean komatiites: geochemical mass balances and constraints on origin. Journal of Geology, 98, 863±877. Hey, M. H., 1954. A new review of the chlorites. Mineralogical Magazine, 30, 277±292. Hickman, A. H., 1973. The North Pole barite Deposits, Pilbara Gold®eld. Annual Reports of Geological Survey of Western Australia, 57±60. Hickman, A. H., 1981. Crustal evolution of the Pilbara Block, Western Australia. In: Archaean Geology (eds Glover, J. E. & Groves, D. I.), pp. 57±69. Geological Society of Australia, Perth. Hickman, A. H., 1983. Geology of the Pilbara Block and its environs. Western Australia Geological Survey Bulletin, 127, 286. Holland, H. D., 1984. The Chemical Evolution of the Atmosphere and Oceans. Princeton University Press, Princeton. Honnorez, J. J., Alt, J. C. & Humphris, S. E., 1998. Vivisection and autopsy of active and fossil hydrothermal alterations of basalt beneath and within the TAG hydrothermal mound. In: Proceedings of the Ocean Drilling Program, Scienti®c Results (eds Herzig, P. M., Humphris, S. E., Miller, D. J. & Zierenberg, R. A.) pp. 231±254. Ishizuka, H., 1985. Prograde metamorphism of the Horokanai ophiolite in the Kamuikotan Zone, Hokkaido, Japan. Journal of Petrology, 26, 391±417. Ishizuka, H. & Suzuki, S., 1995. Ophiolite Metamorphism and Ocean-Floor Metamorphism. Journal of Geography (in Japanese), 104, 350±360. Isozaki, Y., Ueno, Y., Kitajima, K., Kabashima, T. & Maruyama, S., 1998. Early Archean Mid-Oceanic Ridge sediments and the oldest bacteria from the Pilbara Craton, Western Australia. Geological Society of America (Abstract), 30, A98. Kasting, J. F., 1987. Theoretical constraints on oxygen and carbon dioxide concentrations in the Precambrian atmosphere. Precambrian Research, 34, 205±229. H Y D R OTHERM AL ALT E RATION AT A N A R CH AEA N M OR Laverne, C., Vanko, D. A., Tartarotti, P. & Alt, J. C., 1995. Chemistry and geothermometry of secondary minerals from the deep sheeted dike complex, Hole 504B. Proceedings of the Ocean Drilling Program, Scienti®c Results, 137/140, 167±189. Leake, B. E. et al., 1997. Nomenclature of amphiboles: Report of the subcommittee an amphiboles of the International Mineralogical Association, Commission on New Minerals and Mineral Names. American Mineralogist, 82, 1019±1037. Liou, J. G. & Ernst, W. G., 1979. Oceanic ridge metamorphism of the East Taiwan Ophiolite. Contributions to Mineralogy and Petrology, 68, 335±348. Maruyama, S., Isozaki, Y., Ohta, I., Endo, H., Suzuki, S., Kato, Y. & Kimura, G., 1992. Is the Mid-Archean Barite Formation from the Pilbara Craton, Australia, Under the Deep-Sea Environment? Eos Transactions, 80, 532. Maruyama, S., Liou, J. G. & Suzuki, K., 1982. The peristerite gap in low-grade metamorphic rocks. Contributions to Mineralogy and Petrology, 80, 268±276. Matsuda, T. & Isozaki, Y., 1991. Well-documented travel history of Mesozoic pelagic chert in Japan: From remote ocean to subduction zone. Tectonics, 10, 475±499. Mottl, M. & Wheat, G., 1994. Hydrothermal circulation through mid-ocean ridge ¯anks: Fluxes of heat and magnesium. Geochimica et Cosmochimica Acta, 58, 2225±2237. Nijman, W., Bruijne, K. C. H. & d. & Valkering, M. E., 1998. Growth fault control of Early Archaean cherts, barite mounds and chert-barite veins, North Pole Dome, Eastern Pilbara, Western Australia. Precambrian Research, 88, 25±52. Nisbet, E. G., 1995. Archean ecology: a review of evidence for the early development of bacterial biomes, and speculations on the development of a global-scale biosphere. In: Early Precambrian Processes (eds Coward, M. P. & Ries, A. C.), pp. 27±51. The Geological Society of London, London. NystreoÈm, J. O., 1983. Pumpellyite-bearing rocks in Central Sweden and extent of host rock alteration as a control of pumpellyite composition. Contributions to Mineralogy and Petrology, 83, 159±168. Ohta, H., Maruyama, S., Takahashi, E., Watanabe, Y. & Kato, Y., 1996. Field occurrence, geochemistry and petrogenesis of the Archean Mid-Oceanic Ridge Basalts (AMORBs) of the Cleaverville area, Pilbara Craton, Western Australia. Lithos, 37, 199±221. Owen, T. & Cess, R. T., 1979. Enhanced CO2 greenhouse to compensate for reduced solar luminosity on early Earth. Nature, 277, 640±641. Paris, I., Stanistreet, I. G. & Hughes, M. J., 1985. Cherts of the Barberton greenstone belt interpreted as products of submarine exhalative activity. Journal of Geology, 93, 111±129. Powell, R. & Holland, T. J. B., 1985. An internally consistent thermodynamic dataset with uncertainties and correlations: 1. Methods and a worked example. Journal of Metamorphic Geology, 3, 327±342. Powell, R. & Holland, T., 1988. An internally consistent dataset with uncertainties and correlations: 3. Applications to 5 99 geobarometry, worked examples and a computer program. Journal of Metamorphic Geology, 6, 173±204. Rankin, A. H., 1978. H2S-bearing ¯uid inclusions in Baryte from the North Pole deposit, Western Australia. Mineralogical Magazine, 42, 408±410. de Ronde, C. E. J., de Wit, M. J. & Spooner, E. T. C., 1994. Early Archean (>3.2 Ga) Fe-oxide-rich, hydrothermal discharge vents in the Barberton greenstone belt, South Africa. Geological Society of America Bulletin, 106, 86±104. Saccocia, P. J. & Seyfried, W. E., 1994. The solubility of chlorite solid solutions in 3.2 wt% NaCl ¯uids from 300 to 400 uC, 500 bar. Geochimica et Cosmochimica Acta, 58, 567±585. Schiffman, P., Smith, B. M., Varga, R. J. & Moores, E. M., 1987. Geometry, conditions and timing of off-axis hydrothermal metamorphism and ore-deposition in the Solea graben. Nature, 325, 423±425. Suzuki, K., 1975. On some unusual bands veins metasomatically developed in Kasuga, Gifu-ken. Journal of the Geological Society of Japan, 81, 487±504. Takenouchi, S. & Kennedy, G. C., 1964. The binary system H2O-CO2 at high temperatures and pressures. American Journal of Science, 262, 1055±1074. Terabayashi, M., 1993. Compositional evolution in Caamphibole in the Karmutsen metabasites, Vancouver Island, British Columbia, Canada. Journal of Metamorphic Geology, 11, 677±690. Thorpe, R. I., Hickman, A. H., Davis, D. W., Mortensen, J. K. & Trendall, A. F., 1992. U-Pb zircon geochronology of Archaean felsic units in the Marble Bar region, Pilbara Craton, Western Australia. Precambrian Research, 56, 169±189. Ueno, Y., Isozaki, Y., Yurimoto, H. & Maruyama, S., 2001. Carbon isotopic signatures of individual Archean probable microfossil from Western Australia. International Geology Review, (in press). Veizer, J., 1988. The evolving exogenic cycle. In: Chemical Cycles in the Evolution of the Earth (eds Gregory, B., Garrels, R. M., Mackenzie, F. T. & Maynard, J. B.), pp. 175±220. Wiley and Sons, New York. Von Damm, 1995. Controls on the Chemistry and Temporal Variability of Sea¯oor Hydrothermal Fluids. In: Sea¯oor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions (eds Humphris, S. E., Zierenberg, R. A., Mullineaux, L. S. & Thomson, R. E.), Geophigical Monograph 91, 222±247. American Geophysical Union, Washington DC. Walker, J. C. G., Klein, C., Schidlowski, M., Schopf, J. W., Stevenson, D. J. & Walter, M. R., 1983. Environmental evolution of the Archean-early Proterozoic Earth. In: Earth's Earliest Biosphere (ed. Schopf, J. W.), pp. 260±290. Princeton University Press, Princeton. Windley, B. F., 1995. The Evolving Continents, 3rd edn. Wiley & Sons, Chichester. de Wit, M. J., Hart, R., Martin, A. & Abbot, P., 1982. Archaean abiogenic and probable biogenic structures associated with hydrothermal vent systems and regional metasomatism, with implications for Greenstone Belt studies. Economic Geology, 77, 1783±1802. Received 26 July 2000; revision accepted 26 March 2001.