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Stratigraphy-related, low-pressure metamorphism in the Hardey Syncline, Hamersley Province, Western Australia
Gondwana Research 18 (2010) 213–221
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Stratigraphy-related, low-pressure metamorphism in the Hardey Syncline, Hamersley Province, Western Australia Takazo Shibuya a,⁎, Kazumasa Aoki b, Tsuyoshi Komiya b, Shigenori Maruyama b a b
Precambrian Ecosystem Laboratory, Japan Agency for Marine-Earth Science and Technology (JAMSTEC), 2-15 Natsushima-cho, Yokosuka 237-0061, Japan Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 2-12-1 Ookayama, Meguro, Tokyo 152-8551, Japan
a r t i c l e
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Article history: Received 3 October 2009 Received in revised form 18 December 2009 Accepted 11 January 2010 Available online 20 January 2010 Keywords: Paleoproterozoic Low-pressure metamorphism Ophthalmian orogeny Continental rifting Tectonics Hamersley Province
a b s t r a c t The late Archean Fortescue Group and Paleoproterozoic Hamersley Group are exposed in the Hardey Syncline located at the southwestern edge of the Hamersley Basin, Pilbara Craton, Western Australia. The secondary mineral assemblages and compositions of the basaltic rocks of the Fortescue and Hamersley Groups reveal the metamorphic conditions of the study area. The estimated metamorphic grade ranges from prehnite–actinolite facies (Hamersley Group), through greenschist facies (Fortescue Group), to a transition between greenschist facies and actinolite–calcic plagioclase facies (Fortescue Group), indicating a lowpressure type metamorphic facies series. The metamorphic grade increases northward, which is opposite to the general southward increase of the regional metamorphic grade. Furthermore, the change of metamorphic grade strongly correlates with stratigraphy, and the metamorphic temperature increases with stratigraphic depth. These observations suggest that the metamorphism of the study area was caused by a thermal event before the folding due to the Ophthalmian orogeny that affected most of the Hamersley Province. Considering the presence of 2.2 Ga dolerite sills in the Hardey Syncline and the low-pressure metamorphic condition, it is suggested that the metamorphism of the study area was caused by the 2.2 Ga continental rifting, which is consistent with the reported metamorphic age. © 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
1. Introduction Well-preserved Archean–Proterozoic cratons carry important geologic records for understanding the early history of the Earth (e.g., Condie, 1981; Van Kranendonk et al., 2007; Eriksson et al., 2009) because the present day Earth has lost most of its old crust (e.g., Rino et al., 2008; Kawai et al., 2009). Metamorphic processes, including those ongoing in active subduction zones (e.g. Omori et al., 2009), provide critical clues in evaluating the tectonic history. The Pilbara Craton of northwestern Australia is known as one of the best preserved cratons which records accretion of oceanic crust, a large-scale continental rifting, voluminous depositions of banded iron formation and carbonate rock, and large igneous province (LIP) activity (Ohta et al., 1996; Barley et al., 1997; Müller et al., 2005; Van Kranendonk et al., 2007). Such processes have been intensely studied in various terranes to elucidate the early evolutions of solid earth, environment, and life (e.g., Ohta et al., 1996; Barley et al., 1997; Schopf, 2006; Van Kranendonk, 2006; Shibuya et al., 2007a; Van Kranendonk et al., 2007; Komiya et al., 2008; Rino et al., 2008; Santosh and Omori, 2008). Subsequent metamorphic processes on a local and/or regional scale harbor key clues for decoding the tectonic history of cratons.
⁎ Corresponding author. Tel.: + 81 46 867 9647; fax: + 81 46 867 9645. E-mail address: [email protected] (T. Shibuya).
It was previously suggested that greenstones in the Pilbara Craton regionally underwent post-depositional, low-grade burial metamorphism and the metamorphic grade generally increases toward the south (Fig. 1) (Smith et al., 1982). However, geochronological studies (e.g., isochron age) indicate that the greenstones were isotopically reset after the eruption and register thermal events with various ages (Nelson et al., 1992; Alibert and McCulloch, 1993). In-situ dating of metamorphic phosphates in sedimentary rocks was conducted to constrain the precise timing of the thermal events, which revealed that Pilbara Craton was affected by ca. 2.4 Ga and ca. 2.2 Ga thermotectonic events (Rasmussen et al., 2001, 2005). Furthermore, the precise U–Pb ages of the phosphates show a younging trend from south to north, leading to the model of a southward increasing metamorphic grade of the Pilbara Craton associated with the migration of metamorphic front from the collisional zone in the southern margin of the Pilbara Craton (Fig. 1) (Rasmussen et al., 2005). However, Thorne and Tyler (1996) pointed out that the metamorphic mineral isograd pattern reported by Smith et al. (1982) appears to reflect the fold pattern, with lower grade rocks in the synclines and higher grade rocks in anticlines. Thus, more detailed studies on a local scale are essential to reveal the cause of metamorphic events. Here we present a stratigraphy-related investigation of the metamorphism in the Hardy syncline, Hamersley Province, Pilbara Craton, the results from which indicate an opposite direction of increasing grade as against the general southward increasing grade of the regional metamorphism.
1342-937X/$ – see front matter © 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2010.01.002
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2. Geological setting The Pilbara Craton is subdivided into two tectonic components: 3.5 to 2.8 Ga granite–greenstone terrains and <2.8 Ga unconformably overlying volcano-sedimentary successions (Fig. 1). The latter is made up of three stratigraphic components, the Fortescue Group, Hamersley Group, and Turee Creek Group, in ascending order (Thorne, 1990; Thorne and Seymour, 1991). The Fortescue Group consists mainly of mafic volcanics, volcaniclastic rocks, and sedimentary rocks. The Hamersley Group unconformably overlies the Fortescue Group, and consists mainly of banded iron formation (BIF), shale, carbonate, and tuff, interlayered with mafic and felsic igneous rocks. After the deposition, the southern Hamersley Province underwent folding and deformation (mainly WNW–ESE trend) during the Ophthalmian orogeny associated with collision along the southern margin of the Pilbara Craton (Fig. 1) (Blake and Barley, 1992). The metamorphic
grade of the Fortescue Group decreases northward on a regional scale (Smith et al., 1982). The Hardey Syncline, southern Hamersley Basin is one of the main folds related to the Ophthalmian orogeny (Fig. 1). The study area is located in the northern limb of the Hardey Syncline and exposes the Fortescue and Hamersley Groups (Figs. 1 and 2). The lower part of the area is composed mainly of pillow basalts, massive sheet flows, hyaloclastites, reworked hyaloclastites, basaltic komatiite, and minor dikes, often intercalated with thin-bedded cherts and felsic tuffs, which correspond to the Fortescue Group (Figs. 2 and 3; Boongal Formation, Pyradie Formation, Bunjinah Formation, and Jerinah Formation in ascending order). Sub-spherical to ellipsoidal pillows are closed-packed and range from 50 to 200 cm across. They have a dark-colored chilled margin without any trace of vesicles. The way-up structure of the pillows points to the south (Komiya, 2004). Although the greenstones of the Fortescue Groups are part of the continental
Fig. 1. Simplified geological map of the Pilbara Craton (modified after Trendall, 1990). Broken lines indicate metamorphic boundaries estimated mainly from secondary mineral assemblages of greenstones in the Fortescue Group; ZI, prehnite–pumpellyite zone; ZII, prehnite–pumpellyite–epidote zone; ZIII, prehnite–pumpellyite–epidote–actinolite zone; ZIV, (prehnite)–epidote–actinolite zone (Smith et al., 1982). In-situ U–Pb ages of metamorphic phosphate minerals in sedimentary rocks are also plotted (Rasmussen et al., 2005).
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Fig. 2. Geological map of the study area (modified after Komiya, 2004). Secondary mineral assemblages of basaltic greenstones are also plotted. Grid number corresponds to the geological map of Thorne et al. (1995).
flood basalt, the greenstones in the study area erupted under the sea (e.g., Thorne and Trendall, 2001). Although the contact between the Fortescue and Hamersley Groups is generally unconformable, the Hamersley Group in the study area conformably overlies the Fortescue Group (Thorne and Tyler, 1996). The Hamersley Group in the study area consists of mainly sedimentary rocks of BIF, chert, tuff, and felsic, basaltic, and komatiitic volcanic rocks (Figs. 2 and 3). The BIF (Marra Mamba Iron Formation, Mount McRae Shale, and Brockman Iron Formation) is intercalated with minor tuffaceous layers, shales, cherty conglomerate and a thin greenstone sheet (Komiya, 2004). Repeated basaltic and komatiitic flows (Weeli Wolli Formation) conformably overlie the BIF and are intercalated with thin BIF layers, indicating the frequent hiatus and submarine eruption of basalt/basaltic komatiite volcanism. The overlying rhyolite (Woongarra Rhyolite) shows cooling joints, amygdules, flow banding, and breccias. The Weeli Wolli Formation and the Woongarra Ryholite are considered to have resulted from a LIP activity (Barley et al., 1997). The uppermost sedimentary rocks of the study area comprise alternating BIF and tuffaceous layers with monolithic chert layers (Boolgeeda Iron Formation). In the study area, the strike of sedimentary bedding is NW–SE, while the dips are mostly 50–60°S (Fig. 2). The depositional age of the Woongarra Rhyolite is constrained by the U–Pb age of zircons, which is dated at 2449 ± 3 Ma (Barley et al., 1997). The upper part of the Marra Mamba Formation (bottom of the Hamersley Group) and the Jeerinah Formation (top of the Fortescue) are dated at 2597 ± 5 Ma (Trendall et al., 1998) and
2629 ± 5 Ma (Nelson et al., 1999; Trendall et al., 2004), respectively. The depositional environment is interpreted as a sedimentary basin in an open rift valley (Thorne and Seymour, 1991). 3. Secondary mineral parageneses The volcanic rocks in the study area were weakly metamorphosed to yield secondary minerals such as clay minerals, quartz (Qz), calcite (Cc), chlorite (Chl), K-feldspar, sericite, and biotite in felsic rock, and chlorite, epidote (Ep), actinolite (Act), prehnite (Prh), albite (Ab), oligoclase (Olg), biotite, quartz, sericite, and calcite in basaltic and komatiitic rocks (Fig. 4A, B, and C). Under such low-grade metamorphism, the original igneous textures are ubiquitously preserved. Rhyolite has quartz and feldspar phenocrysts, and aggregates of quartz and feldspar with felsitic texture in the groundmass (Fig. 4A). Quartz phenocrysts include melt inclusions that were replaced by chlorite/ smectite and sericite. K-feldspar phenocrysts are replaced by sericite and secondary K-feldspar, chlorite/smectite, and quartz (Fig. 4A). Plagioclase phenocrysts are replaced by albite, chlorite/smectite, and calcite. Minor biotite phenocrysts are replaced by chlorite/smectite and/or secondary biotite. Among basaltic rocks, the fine-grained samples have intergranular, or intersertal textures, whereas coarse-grained samples display ophitic textures. The original igneous minerals in a groundmass of fine-grained volcanic rocks are partly or completely replaced by secondary minerals such as chlorite, prehnite, epidote, actinolite, quartz, sericite, biotite, and
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Fig. 3. Simplified stratigraphic column section with metamorphic zones A to C of the study area. The lithologic patterns are same as in Fig. 2.
calcite (Fig. 4B and C). Many igneous clinopyroxenes are well-preserved in some coarse-grained samples. Olivine phenocrysts in basaltic rocks are completely replaced by aggregates of quartz, chlorite, and/or epidote. Igneous plagioclase is replaced by epidote, albite, prehnite, sericite, and calcite. Basaltic komatiite has a spinifex texture with clinopyroxene dendrites that tends to preserve igneous clinopyroxene in the core and is rimmed by chlorite and/or actinolite. Interstitial plagioclase is replaced by chlorite, albite, epidote, sericite, and minor K-feldspar and biotite. We subdivided the study area into three zones from A to C on the basis of the secondary mineral assemblages of basaltic greenstones (Figs. 2 and 4); note that the secondary minerals of rhyolite were excluded from the following description because of the difference in effective bulk composition for secondary minerals. Zone A is characterized by the assemblage Chl ± Prh + Ep ± Act + Qz + Ab ± Cc (Fig. 5). The mineral assemblage of Zone B is Chl + Ep + Act + Qz + Ab ± Cc and is characterized by the absence of prehnite compared to Zone A (Fig. 5). Zone C is characterized by the first emergence of calcic plagioclase with oligoclase composition; the mineral assemblage is Chl + Ep + Act + Ab + Olg + Qz ± Cc (Fig. 5). The systematic change in the secondary assemblage from Zones A to C indicates that the assemblage changes stratigraphically downwards.
Fig. 4. Microscopic images of volcanic rocks in the study area, showing secondary minerals. (A) Rhyolite in the upper part of the volcanic unit, (B) Basaltic rock in the Hamersley Group, (C) Basaltic rock in the uppermost part of the Fortescue Group.
4. Mineral chemistry The chemical compositions of secondary minerals in basaltic greenstones were analysed with an electron probe analyzer (JEOLJXA-8800 M) at Tokyo Institute of Technology. All analyses were performed with an accelerating voltage of 15 kV, a 12 nA beam current, and a counting time of 10–40 s.
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samples; subhedral to irregular-shape oligoclase grain (<5 μm) is surrounded by host albite. For example, plagioclase in 95BR107 has a bimodal composition of 2.2–8.3 and 15.2–23.8% of An content (Fig. 6). In summary, the plagioclases in Zones A and B are one phase with albite composition, but two phases of plagioclase with a peristerite gap appears in Zone C (Fig. 6). 4.2. Prehnite
Fig. 5. Secondary mineral parageneses of basaltic greenstones in the study area.
Prehnite occurs only in Zone A, replacing the interstitial glass and/or plagioclase. Assuming that the total iron is ferric, the variation of the XFe3+ (=Fe3+/(Fe3+ + Al)) is illustrated in Fig. 7. The XFe3+ of prehnite tends to be lower than that of coexisting epidote.
4.1. Plagioclase
4.3. Epidote
Secondary plagioclase is ubiquitous in all zones, whereas the igneous calcic plagioclases were completely replaced by secondary sodic plagioclase, sericite, calcite, and/or epidote. The anorthite (An) content of the plagioclases from Zones A to C is shown in Fig. 6. In Zones A and B, all of the plagioclases are albite, and the An content ranges from 0.4 to 6.5%; plagioclases in Zones A and B do not show a clear compositional difference in An content. In Zone C, the An content of plagioclase is clearly higher than those in Zones A and B, and albite and oligoclase coexist with a peristerite gap in some
Epidote is one of the most common Ca–Al silicates in the greenstones of the study area. It occurs as granular aggregates in the groundmass with igneous textures and replaces plagioclase phenocrysts accompanied with chlorite in a few places. The variation of XFe3+ of epidote, assuming no ferrous iron and a formula of 12.5 oxygens, is illustrated in Fig. 7. Epidote in the study area shows a wide compositional variation in terms of XFe3+ but has relatively homogeneous composition within a single zone (Fig. 7). In Zone A, XFe3+ of epidote ranges from 0.27 to 0.31, and epidote coexisting with prehnite shows higher XFe3+ than
Fig. 6. Frequency diagrams for anorthite (An) content of secondary plagioclase from Zones A to C. A peristeritic gap between albite and oligoclase is also shown by broken lines.
Fig. 7. Frequency diagrams for XFe3+ in the prehnite and epidote from Zones A to C.
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that of prehnite. All the epidotes in Zones B and C have lower XFe3+ (0.10–0.19) as compared to those in Zone A. The epidotes in the Fortescue Group greenstones from the whole Pilbara Craton have a similar XFe3+ variation (Smith et al., 1982). 4.4. Calcic amphibole Colorless to pale-green Ca-amphiboles occur in Zones A, B, and C. The amphibole replaces the rim of relict clinopyroxene. Chemical compositions of Ca-amphiboles in Zones A to C are plotted on the amphibole classification diagram (Fig. 8A and B). The Fe3+ content of amphibole was estimated from the stoichiometry based on a formula of 23 oxygens (Terabayashi, 1993). The Si content of amphibole in Zone A ranges from 7.71 to 7.95, and the XMg (=Mg/(Mg + Fe)) ranges from 0.54 to 0.62 (Fig. 8A). The amphibole in Zone B has a Si content from 7.59 to 7.97 and an XMg from 0.51 to 0.77, while the composition of amphibole in Zone C ranges from 7.66 to 7.92 in Si content and from 0.50 to 0.64 in XMg. Based on the Ca-amphibole discrimination diagram of Leake et al. (1997), all amphiboles are classified as actinolite. Ca-amphibole plotted on the Hallimond diagram (Hallimond, 1943) does not show systematic change in the edenite, pargasite, and tschermakite components from Zones A to C (Fig. 8B). 4.5. Chlorite Chlorite replaces igneous olivine, clinopyroxene, plagioclase, and interstitial glass. The chemical compositions of chlorites from Zones A to C were plotted on Hey's diagram (Hey, 1954), assuming no ferric iron, on the basis of 28 oxygens in the formula (Fig. 9).
Fig. 9. Chemical composition of chlorite in the study area along with compositional variations of chlorites from the 3.2–3.0 Ga Cleaverville area, Pilbara Craton, Western Australia (Shibuya et al., 2007a).
The chlorites have large compositional variations even within a single zone, but are relatively homogeneous in terms of the XFe (=Fe/ (Fe + Mg)) in each sample. Chlorite in Zone A has smaller compositional variation than that in Zones B and C. The XFe of chlorite in Zone A ranges from 0.51 to 0.56, and the Si content from 5.38 to 5.99. The XFe and Si content of chlorite in Zone B ranges from 0.39 to 0.62 and 5.32 to 6.00, respectively. Chlorite in Zone C has XFe ranging from 0.39 to 0.63 and Si content ranging from 5.47 to 5.92. Chlorite in the study area mostly overlaps the composition of chlorites from hydrothermally metamorphosed greenstones in the 3.2–3.0 Ga Cleaverville area (Shibuya et al., 2007a).
5. Discussion 5.1. Low-pressure metamorphism
Fig. 8. Composition of Ca-amphiboles from Zones A to C, plotted on (A) Leake's diagram (Leake et al., 1997) and (B) Hallimond's diagram (Hallimond, 1943).
5.1.1. Mineral assemblage and composition Here we discuss the compositional changes of secondary minerals related to the metamorphic grade in the study area. In general, the An content of plagioclase increases with metamorphic grade, and two plagioclases have a peristerite gap at around 12% of An content under low-pressure conditions (e.g., Maruyama et al., 1982). Similar compositional changes of plagioclase were reported from the Horokanai ophiolite in Japan (Ishizuka, 1985) and Karmutsen metabasites of the contact metamorphic aureole in Vancouver Island, Canada (Terabayashi, 1993), indicating that the increase in the An content of plagioclase in the metabasite represents an increase in metamorphic grade (Ishizuka, 1985; Terabayashi, 1993). In the study area, the plagioclases in Zones A and B are in one phase with albite composition, but two phases of plagioclase with a peristerite gap appear in Zone C (Fig. 6). Therefore, the increase in the An content of plagioclase in the study area indicates that the metamorphic grade increases from Zone B to C under low-pressure conditions. Furthermore, XFe3+ of epidote in Zone A is clearly higher than that in Zones B and C (Fig. 7). This indicates the lower metamorphic grade of Zone A compared to that of Zones B and C because of the general decrease in XFe3+ of epidote with increasing metamorphic grade (e.g., Ishizuka, 1985). In addition, the XFe of chlorite reflects not only the metamorphic grade but also the XFe of effective bulk composition for secondary minerals (e.g., Hayashi et al., 2000). Therefore, the variation of the XFe within the single zone was probably generated by a difference in the effective bulk composition for secondary minerals.
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The transition from greenschist to amphibolite facies is subdivided into a higher-pressure albite–epidote–amphibolite subfacies and a lower-pressure actinolite–calcic plagioclase subfacies (e.g., Miyashiro, 1973; Maruyama et al., 1983). Therefore, the appearance of calcic plagioclase precedes that of hornblende with increasing temperature at low-pressure (<2 kbar; e.g., Maruyama et al., 1983). A calcic plagioclase-bearing assemblage without hornblende is also reported from the prehnite–actinolite facies of low-pressure contact metamorphism (Terabayashi, 1993) and on-land ophiolites (Liou and Ernst, 1979; Ishizuka, 1985; Shibuya et al., 2007b), which also indicates lowpressure conditions. In the study area, amphibole is the only actinolite (Fig. 8a and b), but calcic plagioclase is present in Zone C (Fig. 6). Therefore, the assemblage of Olg + Act + Ep + Chl in Zone C indicates low-pressure conditions (<2 kbar). 5.1.2. Metamorphic reactions and conditions Two metamorphic reactions that define the boundaries between Zones A to C are investigated based on the calcite-free assemblages. The pressure–temperature (P–T) conditions of the following reactions were calculated with THERMOCALC 3.25 (Holland and Powell, 1998 and its update). A model CaO–MgO–Al2O3–SiO2–H2O system, in the presence of excess quartz and H2O, is assumed. The diagnostic mineral assemblage changes from Chl+Ep+Prh+ Act+Ab+Qz in Zone A to Chl+Ep+Act+Ab+Qz in Zone B. The metamorphic reaction between Zone A and B is defined by the disappearance of prehnite from Zone A to B. This can be represented through the following reaction: 5prehnite þ chlorite þ 2quartz ¼ 4clinozoisiteinepidote þ actinolite þ 6H2 O ð1Þ This reaction has a steep Clapeyron slope on a P–T diagram (Fig. 10), which means that the metamorphic temperature of Zone A is less than approximately 300 °C. Furthermore, the prehnite-bearing assemblage in Zone A indicates prehnite–actinolite facies metamorphism, which strongly suggests low-pressure conditions (<2–3 kbar; Liou et al., 1985; Beiersdorfer and Day, 1995).
Fig. 10. Schematic P–T diagram for metamorphism of basaltic rocks in the study area. Lines 1 and 2 indicate P–T conditions for the metamorphic reactions between Zones A and B, and Zones B and C, respectively. The dark gray band indicates estimated P–T conditions (A–C for Zones A to C). For comparison, a petrological grid is also shown (Liou et al., 1985; Beiersdorfer and Day, 1995). ZEO, zeolite facies; PA, pumpellyite–actinolite facies; PP, prehnite–pumpellyite facies; PRA, prehnite–actinolite facies; BS, blueschist facies; GS, greenschist facies; EA, epidote–amphibolite facies; AP, actinolite–calcic plagioclase facies; AM, amphibolite facies.
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The reaction between Zones B and C, defined by the appearance of oligoclase in the metabasite, can be attributed to the following reaction: chlorite þ 6clinozoisite þ 7quartz ¼ actinolite þ 10anorthite in plagioclase þ 6H2 O
ð2Þ This reaction represents decomposition of epidote and formation of anorthite in secondary plagioclase. It is therefore suggested that the assemblage Chl + Ep+ Act+ Olg in Zone C was formed through reaction (2) and at temperatures >330 °C (Fig. 10). However, this assemblage occurs widely in Zone C, which is not consistent with the discontinuous reaction (2). This should be derived from the assumption of a FeO-free model system. The addition of FeO into the system makes reaction (2) a continuous reaction, and results in a wide temperature range depending on the variation of the FeO/MgO ratio in the system. In addition, the assemblage of Chl + Ep + Act + Olg is generally formed under transition from greenschist facies to actinolite–calcic plagioclase facies, which indicates low-pressure conditions (<2 kbar; Maruyama et al., 1983; Liou et al., 1985). 5.2. Metamorphism of the study area It is well known that the Pilbara Craton underwent post-depositional metamorphic and/or hydrothermal events in whole or part (Smith et al., 1982; Macfarlane and Holland, 1991; Nelson et al., 1992; Alibert and McCulloch, 1993; Erel et al., 1997; Rasmussen et al., 2001, 2005; Rasmussen, 2005). Based on the secondary mineral assemblages of the Fortescue Group greenstones, Smith et al. (1982) suggested that metamorphic grade of the Pilbara Craton increases southward, which was considered to reflect an increasing depth of the burial metamorphism (Fig. 1). However, the precise in-situ dating of metamorphic phosphates in sedimentary rocks revealed that the Pilbara Craton recorded two major thermal events at ca. 2.4 Ga and 2.2 Ga in regional scale (Fig. 1) (Rasmussen et al., 2005). The cause of the former event has been still uncertain but Rasmussen et al. (2005) suggested that the latter metamorphic age corresponds to the Ophthalmian orogeny based on the alignment of the dated phosphate fabrics. This event was considered to be associated with the collision along the southern margin of the Pilbara Craton (Blake and Barley, 1992). Thus, the general southward increasing metamorphic grade of the Pilbara Craton reported by Smith et al. (1982) was reinterpreted to be related to migration of metamorphic front from the collisional zone, spanning a 70 millionyear period from ca. 2215 Ma nearest the collisional margin in the south, to ca. 2145 Ma toward the craton interior in the north (Fig. 1) (Rasmussen et al., 2005). However, this interpretation may be not necessarily be consistent with the metamorphic zones reported by the Smith et al. (1982) because the metamorphic grade increases generally increases southward (from ZI to ZIII) but lower grade zone (ZIII) appears again in the south of the highest zone (ZIV) (Fig. 1). In the study area, the increase of the metamorphic grade from Zone A to C indicate that the metamorphic grade increases northward, which is opposite direction to the general southward increasing grade of the regional metamorphism but probably corresponds to southward decreasing metamorphic grade in the southernmost part of the Pilbara Craton (Fig. 1) (Smith et al., 1982). If the metamorphism had occurred during the Ophthalmian orogeny due to the collision from the south, metamorphic grade would decrease toward the north. It is therefore suggested that the metamorphism of the Hamersley basin does not correspond to a simple southward increasing grade. More importantly, the increasing metamorphic grade from Zone A to C also indicates an increase in metamorphic temperature with stratigraphic depth. The stratigraphy-related thermal structure could not be recorded if the metamorphism had occurred after folding due to the Ophthalmian orogeny. Therefore, the preservation of stratigraphy-related metamorphism strongly suggests that the metamorphism was caused by a thermal
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event before the Ophthalmian folding. Previously, Thorne and Tyler (1996) presumed that the isograd pattern in the Hamersley Basin reported by Smith et al. (1982) appears to reflect the fold pattern, with lower grade rocks in the synclines and higher grade rocks in anticlines. This indicates that the stratigraphically lower units show higher metamorphic grades and that the metamorphic pattern was folded by the Ophthalmian orogeny. Hence, the stratigraphy-related metamorphism reported here supports the prediction by Thorne and Tyler (1996). Although geochronological study on the metamorphic minerals indicates that thermal event prior to the Ophthalmian orogeny was alternatively the 2.4 Ga unknown thermal event (Rasmussen et al., 2005), there has been no geological evidence for the 2.4 Ga thermal event in the Hardey Syncline. However, pre-Ophthalmian magmatic activity was recorded in the Hardey syncline as dolerite sills that intruded into the Turee Creek Group, which is an evidence for ca. 2.2 Ga continental rifting before the beginning of the Ophthalmian orogeny (Müller et al., 2005). The dolerite sill has a baddeleyite age of 2208 ± 10 Ma (Müller et al., 2005), which is identical to the age of 2216 ± 13 Ma for metamorphic phosphate in sedimentary rocks near the study area (Fig. 1) (Rasmussen et al., 2005) within the errors. Furthermore, the mafic sill was folded by the Ophthalmian folding, indicating that the rifting event preceded the Ophthalmian folding (Müller et al., 2005). Therefore, the stratigraphy-related, low-pressure metamorphism of the study area was likely caused by the 2.2 Ga continental rifting. As pointed by Müller et al. (2005), the age of 2.2 Ga is a global peak in LIP activity (French et al., 2004) as observed in ca. 2210 Ma (2217.2 ± 4 and 2209.6 ± 3.5 Ma) Nippissing sills that intruded the Huronian Supergroup of the Superior Province, Canada (Noble and Lightfoot, 1992) and in 2222 ± 13 Ma volcanic rocks of the Ongeluk and Hekpoort formations in the Transvaal Supergroup of the Kaapvaal Craton, South Africa (Cornell et al., 1996). Such a vigorous magmatic activity should have generated a high geothermal gradient in the continental crust, which could cause the stratigraphy-related, low-pressure metamorphism observed in the study area.
6. Conclusions (1) The metamorphic grades of the Hamersley and the Fortescue Groups in the Hardey Syncline are newly defined in this paper. The secondary mineral assemblages and compositions of basaltic rocks from the study area indicate that the metamorphic grade ranges from prehnite–actinolite facies (Hamersley Group), through greenschist facies (Fortescue Group), to a transition between greenschist facies and actinolite–calcic plagioclase facies (Fortescue Group). (2) The metamorphic grade in the study area increases northward, namely stratigraphically downward, which is also not consistent with the model that the regional metamorphism was caused by the Ophthalmian orogeny (Rasmussen et al., 2005). However, this finding is consistent with a southward decreasing metamorphic grade in the southernmost part of the Pilbara Craton (Smith et al., 1982). Furthermore, the low-pressure type metamorphic facies series and the correlation between the metamorphic grade and stratigraphy suggest that the metamorphism of the study area was caused by a thermal event before the Ophthalmian folding. (3) The mafic sills with an age of 2208 ± 10 Ma in the Hardey Syncline are considered to be associated with a continental rifting event immediately prior to the Ophthalmian orogeny (Müller et al., 2005), which is interpreted as the most probable cause of metamorphism in the study area because the metamorphic age reported near the study area (2216 ± 13 Ma; Rasmussen et al., 2005) is identical to the age of intrusion. The low-pressure metamorphic condition and the increasing temperature with stratigraphic depth also support this model.
Acknowledgements We thank S. Johnson and an anonymous reviewer for improving the paper. This research was partially supported by the 21st Century COE Program “How to build habitable planets,” Tokyo Institute of Technology, sponsored by the Ministry of Education, Culture, Sports, Technology and Science, Japan. T.S. is grateful for a Research Fellowship from the Japan Society for the Promotion of Science for Young Scientists.
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Gondwana Research j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / g r
Stratigraphy-related, low-pressure metamorphism in the Hardey Syncline, Hamersley Province, Western Australia Takazo Shibuya a,⁎, Kazumasa Aoki b, Tsuyoshi Komiya b, Shigenori Maruyama b a b
Precambrian Ecosystem Laboratory, Japan Agency for Marine-Earth Science and Technology (JAMSTEC), 2-15 Natsushima-cho, Yokosuka 237-0061, Japan Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 2-12-1 Ookayama, Meguro, Tokyo 152-8551, Japan
a r t i c l e
i n f o
Article history: Received 3 October 2009 Received in revised form 18 December 2009 Accepted 11 January 2010 Available online 20 January 2010 Keywords: Paleoproterozoic Low-pressure metamorphism Ophthalmian orogeny Continental rifting Tectonics Hamersley Province
a b s t r a c t The late Archean Fortescue Group and Paleoproterozoic Hamersley Group are exposed in the Hardey Syncline located at the southwestern edge of the Hamersley Basin, Pilbara Craton, Western Australia. The secondary mineral assemblages and compositions of the basaltic rocks of the Fortescue and Hamersley Groups reveal the metamorphic conditions of the study area. The estimated metamorphic grade ranges from prehnite–actinolite facies (Hamersley Group), through greenschist facies (Fortescue Group), to a transition between greenschist facies and actinolite–calcic plagioclase facies (Fortescue Group), indicating a lowpressure type metamorphic facies series. The metamorphic grade increases northward, which is opposite to the general southward increase of the regional metamorphic grade. Furthermore, the change of metamorphic grade strongly correlates with stratigraphy, and the metamorphic temperature increases with stratigraphic depth. These observations suggest that the metamorphism of the study area was caused by a thermal event before the folding due to the Ophthalmian orogeny that affected most of the Hamersley Province. Considering the presence of 2.2 Ga dolerite sills in the Hardey Syncline and the low-pressure metamorphic condition, it is suggested that the metamorphism of the study area was caused by the 2.2 Ga continental rifting, which is consistent with the reported metamorphic age. © 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
1. Introduction Well-preserved Archean–Proterozoic cratons carry important geologic records for understanding the early history of the Earth (e.g., Condie, 1981; Van Kranendonk et al., 2007; Eriksson et al., 2009) because the present day Earth has lost most of its old crust (e.g., Rino et al., 2008; Kawai et al., 2009). Metamorphic processes, including those ongoing in active subduction zones (e.g. Omori et al., 2009), provide critical clues in evaluating the tectonic history. The Pilbara Craton of northwestern Australia is known as one of the best preserved cratons which records accretion of oceanic crust, a large-scale continental rifting, voluminous depositions of banded iron formation and carbonate rock, and large igneous province (LIP) activity (Ohta et al., 1996; Barley et al., 1997; Müller et al., 2005; Van Kranendonk et al., 2007). Such processes have been intensely studied in various terranes to elucidate the early evolutions of solid earth, environment, and life (e.g., Ohta et al., 1996; Barley et al., 1997; Schopf, 2006; Van Kranendonk, 2006; Shibuya et al., 2007a; Van Kranendonk et al., 2007; Komiya et al., 2008; Rino et al., 2008; Santosh and Omori, 2008). Subsequent metamorphic processes on a local and/or regional scale harbor key clues for decoding the tectonic history of cratons.
⁎ Corresponding author. Tel.: + 81 46 867 9647; fax: + 81 46 867 9645. E-mail address: [email protected] (T. Shibuya).
It was previously suggested that greenstones in the Pilbara Craton regionally underwent post-depositional, low-grade burial metamorphism and the metamorphic grade generally increases toward the south (Fig. 1) (Smith et al., 1982). However, geochronological studies (e.g., isochron age) indicate that the greenstones were isotopically reset after the eruption and register thermal events with various ages (Nelson et al., 1992; Alibert and McCulloch, 1993). In-situ dating of metamorphic phosphates in sedimentary rocks was conducted to constrain the precise timing of the thermal events, which revealed that Pilbara Craton was affected by ca. 2.4 Ga and ca. 2.2 Ga thermotectonic events (Rasmussen et al., 2001, 2005). Furthermore, the precise U–Pb ages of the phosphates show a younging trend from south to north, leading to the model of a southward increasing metamorphic grade of the Pilbara Craton associated with the migration of metamorphic front from the collisional zone in the southern margin of the Pilbara Craton (Fig. 1) (Rasmussen et al., 2005). However, Thorne and Tyler (1996) pointed out that the metamorphic mineral isograd pattern reported by Smith et al. (1982) appears to reflect the fold pattern, with lower grade rocks in the synclines and higher grade rocks in anticlines. Thus, more detailed studies on a local scale are essential to reveal the cause of metamorphic events. Here we present a stratigraphy-related investigation of the metamorphism in the Hardy syncline, Hamersley Province, Pilbara Craton, the results from which indicate an opposite direction of increasing grade as against the general southward increasing grade of the regional metamorphism.
1342-937X/$ – see front matter © 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2010.01.002
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2. Geological setting The Pilbara Craton is subdivided into two tectonic components: 3.5 to 2.8 Ga granite–greenstone terrains and <2.8 Ga unconformably overlying volcano-sedimentary successions (Fig. 1). The latter is made up of three stratigraphic components, the Fortescue Group, Hamersley Group, and Turee Creek Group, in ascending order (Thorne, 1990; Thorne and Seymour, 1991). The Fortescue Group consists mainly of mafic volcanics, volcaniclastic rocks, and sedimentary rocks. The Hamersley Group unconformably overlies the Fortescue Group, and consists mainly of banded iron formation (BIF), shale, carbonate, and tuff, interlayered with mafic and felsic igneous rocks. After the deposition, the southern Hamersley Province underwent folding and deformation (mainly WNW–ESE trend) during the Ophthalmian orogeny associated with collision along the southern margin of the Pilbara Craton (Fig. 1) (Blake and Barley, 1992). The metamorphic
grade of the Fortescue Group decreases northward on a regional scale (Smith et al., 1982). The Hardey Syncline, southern Hamersley Basin is one of the main folds related to the Ophthalmian orogeny (Fig. 1). The study area is located in the northern limb of the Hardey Syncline and exposes the Fortescue and Hamersley Groups (Figs. 1 and 2). The lower part of the area is composed mainly of pillow basalts, massive sheet flows, hyaloclastites, reworked hyaloclastites, basaltic komatiite, and minor dikes, often intercalated with thin-bedded cherts and felsic tuffs, which correspond to the Fortescue Group (Figs. 2 and 3; Boongal Formation, Pyradie Formation, Bunjinah Formation, and Jerinah Formation in ascending order). Sub-spherical to ellipsoidal pillows are closed-packed and range from 50 to 200 cm across. They have a dark-colored chilled margin without any trace of vesicles. The way-up structure of the pillows points to the south (Komiya, 2004). Although the greenstones of the Fortescue Groups are part of the continental
Fig. 1. Simplified geological map of the Pilbara Craton (modified after Trendall, 1990). Broken lines indicate metamorphic boundaries estimated mainly from secondary mineral assemblages of greenstones in the Fortescue Group; ZI, prehnite–pumpellyite zone; ZII, prehnite–pumpellyite–epidote zone; ZIII, prehnite–pumpellyite–epidote–actinolite zone; ZIV, (prehnite)–epidote–actinolite zone (Smith et al., 1982). In-situ U–Pb ages of metamorphic phosphate minerals in sedimentary rocks are also plotted (Rasmussen et al., 2005).
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Fig. 2. Geological map of the study area (modified after Komiya, 2004). Secondary mineral assemblages of basaltic greenstones are also plotted. Grid number corresponds to the geological map of Thorne et al. (1995).
flood basalt, the greenstones in the study area erupted under the sea (e.g., Thorne and Trendall, 2001). Although the contact between the Fortescue and Hamersley Groups is generally unconformable, the Hamersley Group in the study area conformably overlies the Fortescue Group (Thorne and Tyler, 1996). The Hamersley Group in the study area consists of mainly sedimentary rocks of BIF, chert, tuff, and felsic, basaltic, and komatiitic volcanic rocks (Figs. 2 and 3). The BIF (Marra Mamba Iron Formation, Mount McRae Shale, and Brockman Iron Formation) is intercalated with minor tuffaceous layers, shales, cherty conglomerate and a thin greenstone sheet (Komiya, 2004). Repeated basaltic and komatiitic flows (Weeli Wolli Formation) conformably overlie the BIF and are intercalated with thin BIF layers, indicating the frequent hiatus and submarine eruption of basalt/basaltic komatiite volcanism. The overlying rhyolite (Woongarra Rhyolite) shows cooling joints, amygdules, flow banding, and breccias. The Weeli Wolli Formation and the Woongarra Ryholite are considered to have resulted from a LIP activity (Barley et al., 1997). The uppermost sedimentary rocks of the study area comprise alternating BIF and tuffaceous layers with monolithic chert layers (Boolgeeda Iron Formation). In the study area, the strike of sedimentary bedding is NW–SE, while the dips are mostly 50–60°S (Fig. 2). The depositional age of the Woongarra Rhyolite is constrained by the U–Pb age of zircons, which is dated at 2449 ± 3 Ma (Barley et al., 1997). The upper part of the Marra Mamba Formation (bottom of the Hamersley Group) and the Jeerinah Formation (top of the Fortescue) are dated at 2597 ± 5 Ma (Trendall et al., 1998) and
2629 ± 5 Ma (Nelson et al., 1999; Trendall et al., 2004), respectively. The depositional environment is interpreted as a sedimentary basin in an open rift valley (Thorne and Seymour, 1991). 3. Secondary mineral parageneses The volcanic rocks in the study area were weakly metamorphosed to yield secondary minerals such as clay minerals, quartz (Qz), calcite (Cc), chlorite (Chl), K-feldspar, sericite, and biotite in felsic rock, and chlorite, epidote (Ep), actinolite (Act), prehnite (Prh), albite (Ab), oligoclase (Olg), biotite, quartz, sericite, and calcite in basaltic and komatiitic rocks (Fig. 4A, B, and C). Under such low-grade metamorphism, the original igneous textures are ubiquitously preserved. Rhyolite has quartz and feldspar phenocrysts, and aggregates of quartz and feldspar with felsitic texture in the groundmass (Fig. 4A). Quartz phenocrysts include melt inclusions that were replaced by chlorite/ smectite and sericite. K-feldspar phenocrysts are replaced by sericite and secondary K-feldspar, chlorite/smectite, and quartz (Fig. 4A). Plagioclase phenocrysts are replaced by albite, chlorite/smectite, and calcite. Minor biotite phenocrysts are replaced by chlorite/smectite and/or secondary biotite. Among basaltic rocks, the fine-grained samples have intergranular, or intersertal textures, whereas coarse-grained samples display ophitic textures. The original igneous minerals in a groundmass of fine-grained volcanic rocks are partly or completely replaced by secondary minerals such as chlorite, prehnite, epidote, actinolite, quartz, sericite, biotite, and
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Fig. 3. Simplified stratigraphic column section with metamorphic zones A to C of the study area. The lithologic patterns are same as in Fig. 2.
calcite (Fig. 4B and C). Many igneous clinopyroxenes are well-preserved in some coarse-grained samples. Olivine phenocrysts in basaltic rocks are completely replaced by aggregates of quartz, chlorite, and/or epidote. Igneous plagioclase is replaced by epidote, albite, prehnite, sericite, and calcite. Basaltic komatiite has a spinifex texture with clinopyroxene dendrites that tends to preserve igneous clinopyroxene in the core and is rimmed by chlorite and/or actinolite. Interstitial plagioclase is replaced by chlorite, albite, epidote, sericite, and minor K-feldspar and biotite. We subdivided the study area into three zones from A to C on the basis of the secondary mineral assemblages of basaltic greenstones (Figs. 2 and 4); note that the secondary minerals of rhyolite were excluded from the following description because of the difference in effective bulk composition for secondary minerals. Zone A is characterized by the assemblage Chl ± Prh + Ep ± Act + Qz + Ab ± Cc (Fig. 5). The mineral assemblage of Zone B is Chl + Ep + Act + Qz + Ab ± Cc and is characterized by the absence of prehnite compared to Zone A (Fig. 5). Zone C is characterized by the first emergence of calcic plagioclase with oligoclase composition; the mineral assemblage is Chl + Ep + Act + Ab + Olg + Qz ± Cc (Fig. 5). The systematic change in the secondary assemblage from Zones A to C indicates that the assemblage changes stratigraphically downwards.
Fig. 4. Microscopic images of volcanic rocks in the study area, showing secondary minerals. (A) Rhyolite in the upper part of the volcanic unit, (B) Basaltic rock in the Hamersley Group, (C) Basaltic rock in the uppermost part of the Fortescue Group.
4. Mineral chemistry The chemical compositions of secondary minerals in basaltic greenstones were analysed with an electron probe analyzer (JEOLJXA-8800 M) at Tokyo Institute of Technology. All analyses were performed with an accelerating voltage of 15 kV, a 12 nA beam current, and a counting time of 10–40 s.
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samples; subhedral to irregular-shape oligoclase grain (<5 μm) is surrounded by host albite. For example, plagioclase in 95BR107 has a bimodal composition of 2.2–8.3 and 15.2–23.8% of An content (Fig. 6). In summary, the plagioclases in Zones A and B are one phase with albite composition, but two phases of plagioclase with a peristerite gap appears in Zone C (Fig. 6). 4.2. Prehnite
Fig. 5. Secondary mineral parageneses of basaltic greenstones in the study area.
Prehnite occurs only in Zone A, replacing the interstitial glass and/or plagioclase. Assuming that the total iron is ferric, the variation of the XFe3+ (=Fe3+/(Fe3+ + Al)) is illustrated in Fig. 7. The XFe3+ of prehnite tends to be lower than that of coexisting epidote.
4.1. Plagioclase
4.3. Epidote
Secondary plagioclase is ubiquitous in all zones, whereas the igneous calcic plagioclases were completely replaced by secondary sodic plagioclase, sericite, calcite, and/or epidote. The anorthite (An) content of the plagioclases from Zones A to C is shown in Fig. 6. In Zones A and B, all of the plagioclases are albite, and the An content ranges from 0.4 to 6.5%; plagioclases in Zones A and B do not show a clear compositional difference in An content. In Zone C, the An content of plagioclase is clearly higher than those in Zones A and B, and albite and oligoclase coexist with a peristerite gap in some
Epidote is one of the most common Ca–Al silicates in the greenstones of the study area. It occurs as granular aggregates in the groundmass with igneous textures and replaces plagioclase phenocrysts accompanied with chlorite in a few places. The variation of XFe3+ of epidote, assuming no ferrous iron and a formula of 12.5 oxygens, is illustrated in Fig. 7. Epidote in the study area shows a wide compositional variation in terms of XFe3+ but has relatively homogeneous composition within a single zone (Fig. 7). In Zone A, XFe3+ of epidote ranges from 0.27 to 0.31, and epidote coexisting with prehnite shows higher XFe3+ than
Fig. 6. Frequency diagrams for anorthite (An) content of secondary plagioclase from Zones A to C. A peristeritic gap between albite and oligoclase is also shown by broken lines.
Fig. 7. Frequency diagrams for XFe3+ in the prehnite and epidote from Zones A to C.
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that of prehnite. All the epidotes in Zones B and C have lower XFe3+ (0.10–0.19) as compared to those in Zone A. The epidotes in the Fortescue Group greenstones from the whole Pilbara Craton have a similar XFe3+ variation (Smith et al., 1982). 4.4. Calcic amphibole Colorless to pale-green Ca-amphiboles occur in Zones A, B, and C. The amphibole replaces the rim of relict clinopyroxene. Chemical compositions of Ca-amphiboles in Zones A to C are plotted on the amphibole classification diagram (Fig. 8A and B). The Fe3+ content of amphibole was estimated from the stoichiometry based on a formula of 23 oxygens (Terabayashi, 1993). The Si content of amphibole in Zone A ranges from 7.71 to 7.95, and the XMg (=Mg/(Mg + Fe)) ranges from 0.54 to 0.62 (Fig. 8A). The amphibole in Zone B has a Si content from 7.59 to 7.97 and an XMg from 0.51 to 0.77, while the composition of amphibole in Zone C ranges from 7.66 to 7.92 in Si content and from 0.50 to 0.64 in XMg. Based on the Ca-amphibole discrimination diagram of Leake et al. (1997), all amphiboles are classified as actinolite. Ca-amphibole plotted on the Hallimond diagram (Hallimond, 1943) does not show systematic change in the edenite, pargasite, and tschermakite components from Zones A to C (Fig. 8B). 4.5. Chlorite Chlorite replaces igneous olivine, clinopyroxene, plagioclase, and interstitial glass. The chemical compositions of chlorites from Zones A to C were plotted on Hey's diagram (Hey, 1954), assuming no ferric iron, on the basis of 28 oxygens in the formula (Fig. 9).
Fig. 9. Chemical composition of chlorite in the study area along with compositional variations of chlorites from the 3.2–3.0 Ga Cleaverville area, Pilbara Craton, Western Australia (Shibuya et al., 2007a).
The chlorites have large compositional variations even within a single zone, but are relatively homogeneous in terms of the XFe (=Fe/ (Fe + Mg)) in each sample. Chlorite in Zone A has smaller compositional variation than that in Zones B and C. The XFe of chlorite in Zone A ranges from 0.51 to 0.56, and the Si content from 5.38 to 5.99. The XFe and Si content of chlorite in Zone B ranges from 0.39 to 0.62 and 5.32 to 6.00, respectively. Chlorite in Zone C has XFe ranging from 0.39 to 0.63 and Si content ranging from 5.47 to 5.92. Chlorite in the study area mostly overlaps the composition of chlorites from hydrothermally metamorphosed greenstones in the 3.2–3.0 Ga Cleaverville area (Shibuya et al., 2007a).
5. Discussion 5.1. Low-pressure metamorphism
Fig. 8. Composition of Ca-amphiboles from Zones A to C, plotted on (A) Leake's diagram (Leake et al., 1997) and (B) Hallimond's diagram (Hallimond, 1943).
5.1.1. Mineral assemblage and composition Here we discuss the compositional changes of secondary minerals related to the metamorphic grade in the study area. In general, the An content of plagioclase increases with metamorphic grade, and two plagioclases have a peristerite gap at around 12% of An content under low-pressure conditions (e.g., Maruyama et al., 1982). Similar compositional changes of plagioclase were reported from the Horokanai ophiolite in Japan (Ishizuka, 1985) and Karmutsen metabasites of the contact metamorphic aureole in Vancouver Island, Canada (Terabayashi, 1993), indicating that the increase in the An content of plagioclase in the metabasite represents an increase in metamorphic grade (Ishizuka, 1985; Terabayashi, 1993). In the study area, the plagioclases in Zones A and B are in one phase with albite composition, but two phases of plagioclase with a peristerite gap appear in Zone C (Fig. 6). Therefore, the increase in the An content of plagioclase in the study area indicates that the metamorphic grade increases from Zone B to C under low-pressure conditions. Furthermore, XFe3+ of epidote in Zone A is clearly higher than that in Zones B and C (Fig. 7). This indicates the lower metamorphic grade of Zone A compared to that of Zones B and C because of the general decrease in XFe3+ of epidote with increasing metamorphic grade (e.g., Ishizuka, 1985). In addition, the XFe of chlorite reflects not only the metamorphic grade but also the XFe of effective bulk composition for secondary minerals (e.g., Hayashi et al., 2000). Therefore, the variation of the XFe within the single zone was probably generated by a difference in the effective bulk composition for secondary minerals.
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The transition from greenschist to amphibolite facies is subdivided into a higher-pressure albite–epidote–amphibolite subfacies and a lower-pressure actinolite–calcic plagioclase subfacies (e.g., Miyashiro, 1973; Maruyama et al., 1983). Therefore, the appearance of calcic plagioclase precedes that of hornblende with increasing temperature at low-pressure (<2 kbar; e.g., Maruyama et al., 1983). A calcic plagioclase-bearing assemblage without hornblende is also reported from the prehnite–actinolite facies of low-pressure contact metamorphism (Terabayashi, 1993) and on-land ophiolites (Liou and Ernst, 1979; Ishizuka, 1985; Shibuya et al., 2007b), which also indicates lowpressure conditions. In the study area, amphibole is the only actinolite (Fig. 8a and b), but calcic plagioclase is present in Zone C (Fig. 6). Therefore, the assemblage of Olg + Act + Ep + Chl in Zone C indicates low-pressure conditions (<2 kbar). 5.1.2. Metamorphic reactions and conditions Two metamorphic reactions that define the boundaries between Zones A to C are investigated based on the calcite-free assemblages. The pressure–temperature (P–T) conditions of the following reactions were calculated with THERMOCALC 3.25 (Holland and Powell, 1998 and its update). A model CaO–MgO–Al2O3–SiO2–H2O system, in the presence of excess quartz and H2O, is assumed. The diagnostic mineral assemblage changes from Chl+Ep+Prh+ Act+Ab+Qz in Zone A to Chl+Ep+Act+Ab+Qz in Zone B. The metamorphic reaction between Zone A and B is defined by the disappearance of prehnite from Zone A to B. This can be represented through the following reaction: 5prehnite þ chlorite þ 2quartz ¼ 4clinozoisiteinepidote þ actinolite þ 6H2 O ð1Þ This reaction has a steep Clapeyron slope on a P–T diagram (Fig. 10), which means that the metamorphic temperature of Zone A is less than approximately 300 °C. Furthermore, the prehnite-bearing assemblage in Zone A indicates prehnite–actinolite facies metamorphism, which strongly suggests low-pressure conditions (<2–3 kbar; Liou et al., 1985; Beiersdorfer and Day, 1995).
Fig. 10. Schematic P–T diagram for metamorphism of basaltic rocks in the study area. Lines 1 and 2 indicate P–T conditions for the metamorphic reactions between Zones A and B, and Zones B and C, respectively. The dark gray band indicates estimated P–T conditions (A–C for Zones A to C). For comparison, a petrological grid is also shown (Liou et al., 1985; Beiersdorfer and Day, 1995). ZEO, zeolite facies; PA, pumpellyite–actinolite facies; PP, prehnite–pumpellyite facies; PRA, prehnite–actinolite facies; BS, blueschist facies; GS, greenschist facies; EA, epidote–amphibolite facies; AP, actinolite–calcic plagioclase facies; AM, amphibolite facies.
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The reaction between Zones B and C, defined by the appearance of oligoclase in the metabasite, can be attributed to the following reaction: chlorite þ 6clinozoisite þ 7quartz ¼ actinolite þ 10anorthite in plagioclase þ 6H2 O
ð2Þ This reaction represents decomposition of epidote and formation of anorthite in secondary plagioclase. It is therefore suggested that the assemblage Chl + Ep+ Act+ Olg in Zone C was formed through reaction (2) and at temperatures >330 °C (Fig. 10). However, this assemblage occurs widely in Zone C, which is not consistent with the discontinuous reaction (2). This should be derived from the assumption of a FeO-free model system. The addition of FeO into the system makes reaction (2) a continuous reaction, and results in a wide temperature range depending on the variation of the FeO/MgO ratio in the system. In addition, the assemblage of Chl + Ep + Act + Olg is generally formed under transition from greenschist facies to actinolite–calcic plagioclase facies, which indicates low-pressure conditions (<2 kbar; Maruyama et al., 1983; Liou et al., 1985). 5.2. Metamorphism of the study area It is well known that the Pilbara Craton underwent post-depositional metamorphic and/or hydrothermal events in whole or part (Smith et al., 1982; Macfarlane and Holland, 1991; Nelson et al., 1992; Alibert and McCulloch, 1993; Erel et al., 1997; Rasmussen et al., 2001, 2005; Rasmussen, 2005). Based on the secondary mineral assemblages of the Fortescue Group greenstones, Smith et al. (1982) suggested that metamorphic grade of the Pilbara Craton increases southward, which was considered to reflect an increasing depth of the burial metamorphism (Fig. 1). However, the precise in-situ dating of metamorphic phosphates in sedimentary rocks revealed that the Pilbara Craton recorded two major thermal events at ca. 2.4 Ga and 2.2 Ga in regional scale (Fig. 1) (Rasmussen et al., 2005). The cause of the former event has been still uncertain but Rasmussen et al. (2005) suggested that the latter metamorphic age corresponds to the Ophthalmian orogeny based on the alignment of the dated phosphate fabrics. This event was considered to be associated with the collision along the southern margin of the Pilbara Craton (Blake and Barley, 1992). Thus, the general southward increasing metamorphic grade of the Pilbara Craton reported by Smith et al. (1982) was reinterpreted to be related to migration of metamorphic front from the collisional zone, spanning a 70 millionyear period from ca. 2215 Ma nearest the collisional margin in the south, to ca. 2145 Ma toward the craton interior in the north (Fig. 1) (Rasmussen et al., 2005). However, this interpretation may be not necessarily be consistent with the metamorphic zones reported by the Smith et al. (1982) because the metamorphic grade increases generally increases southward (from ZI to ZIII) but lower grade zone (ZIII) appears again in the south of the highest zone (ZIV) (Fig. 1). In the study area, the increase of the metamorphic grade from Zone A to C indicate that the metamorphic grade increases northward, which is opposite direction to the general southward increasing grade of the regional metamorphism but probably corresponds to southward decreasing metamorphic grade in the southernmost part of the Pilbara Craton (Fig. 1) (Smith et al., 1982). If the metamorphism had occurred during the Ophthalmian orogeny due to the collision from the south, metamorphic grade would decrease toward the north. It is therefore suggested that the metamorphism of the Hamersley basin does not correspond to a simple southward increasing grade. More importantly, the increasing metamorphic grade from Zone A to C also indicates an increase in metamorphic temperature with stratigraphic depth. The stratigraphy-related thermal structure could not be recorded if the metamorphism had occurred after folding due to the Ophthalmian orogeny. Therefore, the preservation of stratigraphy-related metamorphism strongly suggests that the metamorphism was caused by a thermal
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event before the Ophthalmian folding. Previously, Thorne and Tyler (1996) presumed that the isograd pattern in the Hamersley Basin reported by Smith et al. (1982) appears to reflect the fold pattern, with lower grade rocks in the synclines and higher grade rocks in anticlines. This indicates that the stratigraphically lower units show higher metamorphic grades and that the metamorphic pattern was folded by the Ophthalmian orogeny. Hence, the stratigraphy-related metamorphism reported here supports the prediction by Thorne and Tyler (1996). Although geochronological study on the metamorphic minerals indicates that thermal event prior to the Ophthalmian orogeny was alternatively the 2.4 Ga unknown thermal event (Rasmussen et al., 2005), there has been no geological evidence for the 2.4 Ga thermal event in the Hardey Syncline. However, pre-Ophthalmian magmatic activity was recorded in the Hardey syncline as dolerite sills that intruded into the Turee Creek Group, which is an evidence for ca. 2.2 Ga continental rifting before the beginning of the Ophthalmian orogeny (Müller et al., 2005). The dolerite sill has a baddeleyite age of 2208 ± 10 Ma (Müller et al., 2005), which is identical to the age of 2216 ± 13 Ma for metamorphic phosphate in sedimentary rocks near the study area (Fig. 1) (Rasmussen et al., 2005) within the errors. Furthermore, the mafic sill was folded by the Ophthalmian folding, indicating that the rifting event preceded the Ophthalmian folding (Müller et al., 2005). Therefore, the stratigraphy-related, low-pressure metamorphism of the study area was likely caused by the 2.2 Ga continental rifting. As pointed by Müller et al. (2005), the age of 2.2 Ga is a global peak in LIP activity (French et al., 2004) as observed in ca. 2210 Ma (2217.2 ± 4 and 2209.6 ± 3.5 Ma) Nippissing sills that intruded the Huronian Supergroup of the Superior Province, Canada (Noble and Lightfoot, 1992) and in 2222 ± 13 Ma volcanic rocks of the Ongeluk and Hekpoort formations in the Transvaal Supergroup of the Kaapvaal Craton, South Africa (Cornell et al., 1996). Such a vigorous magmatic activity should have generated a high geothermal gradient in the continental crust, which could cause the stratigraphy-related, low-pressure metamorphism observed in the study area.
6. Conclusions (1) The metamorphic grades of the Hamersley and the Fortescue Groups in the Hardey Syncline are newly defined in this paper. The secondary mineral assemblages and compositions of basaltic rocks from the study area indicate that the metamorphic grade ranges from prehnite–actinolite facies (Hamersley Group), through greenschist facies (Fortescue Group), to a transition between greenschist facies and actinolite–calcic plagioclase facies (Fortescue Group). (2) The metamorphic grade in the study area increases northward, namely stratigraphically downward, which is also not consistent with the model that the regional metamorphism was caused by the Ophthalmian orogeny (Rasmussen et al., 2005). However, this finding is consistent with a southward decreasing metamorphic grade in the southernmost part of the Pilbara Craton (Smith et al., 1982). Furthermore, the low-pressure type metamorphic facies series and the correlation between the metamorphic grade and stratigraphy suggest that the metamorphism of the study area was caused by a thermal event before the Ophthalmian folding. (3) The mafic sills with an age of 2208 ± 10 Ma in the Hardey Syncline are considered to be associated with a continental rifting event immediately prior to the Ophthalmian orogeny (Müller et al., 2005), which is interpreted as the most probable cause of metamorphism in the study area because the metamorphic age reported near the study area (2216 ± 13 Ma; Rasmussen et al., 2005) is identical to the age of intrusion. The low-pressure metamorphic condition and the increasing temperature with stratigraphic depth also support this model.
Acknowledgements We thank S. Johnson and an anonymous reviewer for improving the paper. This research was partially supported by the 21st Century COE Program “How to build habitable planets,” Tokyo Institute of Technology, sponsored by the Ministry of Education, Culture, Sports, Technology and Science, Japan. T.S. is grateful for a Research Fellowship from the Japan Society for the Promotion of Science for Young Scientists.
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