Origin and global tectonic significance of Early Archean cherts from the Marble Bar greenstone belt, Pilbara Craton, Western Australia

Precambrian Research 125 (2003) 191–243

Origin and global tectonic significance of Early Archean cherts from the Marble Bar greenstone belt, Pilbara Craton, Western Australia Yasuhiro Kato∗ , Kentaro Nakamura Department of Geosystem Engineering, University of Tokyo, Tokyo 113-8656, Japan Received 28 February 2002; accepted 5 February 2003

Abstract Five sections of bedded chert in mafic-ultramafic rocks of the Archean Warrawoona Group in the Marble Bar greenstone belt, Pilbara Craton, were analyzed in order to understand their depositional environment and to provide some constraints on Early Archean tectonics. The sections are divisible into two types based on their field occurrence, mineralogy and geochemistry; thicker ones (A and B) that overlie Fe-rich, low-K tholeiites and thinner ones (C1, C2, and C3) overlying komatiitic basalts. The thickest, ferruginous section (A; 45 m thick) is the Marble Bar Chert of the Towers Formation, and is interpreted to have been an in situ precipitate derived from a high-T hydrothermal solution emanating from a mid-oceanic ridge (MOR). The following geochemical features are similar to those of modern hydrothermal iron-rich sediments at a MOR: (i) P, V, Zn, and Y are positively correlated with Fe, (ii) a positive Eu anomaly
(normalized to chondrite) decreases from 6.6 to 1.3 and the magnitude of a negative Ce anomaly decreases from 0.6 to 1.0 as REE and LREE/HREE increase. The 13 m thick B-section in the Apex Basalt, dominated by SiO2 and containing significant amounts of Ba (up to 4330 ppm), originated from a low-T MOR hydrothermal solution. This section is characterized by an association with massive black/gray silica veins that were hydrothermal feeders in normal fault zones in the spreading center. Geochemical evidence from the greenstones underlying the B-section indicates that they are of MORB origin. In contrast to A and B, C1-, C2-, and C3-sections are intercalated conformably with komatiitic basalts of the Apex Basalt and are 3–6 m thick. The geochemical signatures of these three sections suggest that they were most likely formed by low-T, weak hydrothermal activity that may have been associated with hot-spot volcanism. They show strong enrichment of Cr and Ni, reflecting a significant contribution of komatiitic basaltic detritus during sedimentation. The presence of volcaniclastic chert in their uppermost beds indicates a decrease of hydrothermal silica precipitation due to waning hydrothermal activity. Among the three sections, the uppermost (C3) exhibits greater enrichment in Zr, Nb, Hf, and Th, and higher Th/Sc and (La/Yb)N relative to the lower C1- and C2-sections. This could mean that the depositional environment of C3 was relatively closer to a continent or to island arcs composed of granitoids and/or their volcanic equivalents. Siliceous mudstones and sandstones of the uppermost clastic rocks (T-section; Panorama Formation) have
geochemical signatures analogous to those of felsic plutonic/volcanic rocks. High Th/Sc, negative Eu anomalies, and high REE in some siliceous mudstones from this unit imply that differentiated granitoids (or volcanic equivalents) were already exposed (erupted), at least locally, in the Pilbara at this time. This study shows these Warrawoona Group cherts were deposited in a variety of environments ranging from a mid-oceanic spreading center to

∗ Corresponding author. Tel.: +81-3-5841-7022; fax: +81-3-3818-7492. E-mail address: [email protected] (Y. Kato).

0301-9268/$ – see front matter © 2003 Elsevier Science B.V. All rights reserved. doi:10.1016/S0301-9268(03)00043-3

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a convergent plate boundary via a hot-spot. This variation was most likely due to horizontal plate motions which accordingly support the operation of plate tectonics in the Early Archean. © 2003 Elsevier Science B.V. All rights reserved. Keywords: Chert; Archean; Pilbara; Geochemistry; Hydrothermal; Plate tectonics

1. Introduction Chert is a common, albeit minor, sediment throughout the geological record from the Archean to the Cenozoic. Phanerozoic biogenic bedded cherts are mainly composed of radiolarian, diatoms, and/or sponge spicules, but minor hydrothermal cherts are locally deposited on mid-oceanic ridge (MOR) basalt (e.g. Adachi et al., 1986). The geochemical characteristics of these Phanerozoic cherts have been welldocumented (Shimizu and Masuda, 1977; Rangin et al., 1981; Sugisaki et al., 1982), because the cherts are markedly resistant to weathering and alteration, and are thus suitable for geochemical studies. Although cherts are generally monomineralic and thus their major element compositions are simple, trace and rare earth elements reflect minor constituents such as hydrothermal precipitates, volcanic detritus, and terrigenous clasts, and hence are a key indicator of depositional environment. From this viewpoint, Murray et al. (1991, 1992) reported appreciable variations of trace and rare earth elements in Phanerozoic cherts deposited in environments ranging from a MOR to a trench. Archean cherts are commonly shown to have been deposited from hydrothermal solutions (e.g. Paris et al., 1985; Sugitani, 1992), but their depositional settings and environments are not tightly constrained. By analogy with Phanerozoic cherts, increased understanding of Early Archean chert chemistry may provide some constraints on the depositional environment at the Earth’s surface at this time, and may help elucidate the debate on Archean plate tectonics. Whether or not a plate tectonic regime is capable of explaining the Early Archean rock record is controversial (Hamilton, 1998; de Wit, 1998), although Late Archean plate tectonics seems to have gained an unshakable consensus (e.g. Calvert et al., 1995; Skulski and Percival, 1996). Hamilton (1998) postulated that plate tectonics did not operate in the Archean because of a much higher heat loss from a hotter mantle, but de Wit (1998)

concluded that the onset of modern plate tectonics dates back to the Hadean–Archean transition between 4.0 and 4.2 Ga. Recently we provided geochemical evidence from 3.2 to 3.1 Ga banded iron formations (BIFs) indicating that plate tectonics was already operating in the Middle Archean (Kato et al., 1998). There are, however, as yet no data on the geochemistry of cherts that can be used to constrain Early Archean tectonics. The tectonic evolution of granite–greenstone terrains in the eastern Pilbara Craton has been also hotly debated. According to one model, the development of the dome-and-syncline map pattern in the eastern Pilbara is considered to have been caused by solid-state diapirism associated with convective overturn of the upper to middle crust, and thus is explained by non-plate tectonic regime (Collins et al., 1998; Van Kranendonk et al., 2001a). In contrast, van Haaften and White (1998) and Kloppenburg et al. (2001) have invoked a regional thrust-accretion tectonic process for the eastern Pilbara Craton. Our geochemical approach for Archean cherts and clastic sedimentary rocks in the Marble Bar greenstone belt may yield useful information about its tectonic evolution. In addition, increased knowledge of Archean chert formation may also help solve the origin of life, because Earth’s earliest putative microfossils have been discovered in these and stratigraphically equivalent cherts (Awramik et al., 1983; Schopf et al., 2002). A much better understanding of the source, depositional environment, and geological setting of Archean cherts is needed to resolve these primary problems of the early Earth. The Early Archean cherts in this study area show a great variety of thickness, color, association with greenstones, lithology, and mineralogy, and have a wide range of chemical compositions. In this contribution, we provide geochemical data on Early Archean cherts and clastic rocks from the Marble Bar greenstone belt in order to define their depositional environments and to infer Archean plate tectonics.

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2. Geological setting The studied chert sections are from the Towers Formation and Apex Basalt in the Warrawoona Group (Hickman, 1983; Van Kranendonk et al., 2001a). The Warrawoona Group consists of the Talga Talga Subgroup (including Duffer Formation) and Salgash Subgroup (Towers Formation, Apex Basalt, and Panorama Formation) in ascending stratigraphic order (Fig. 1; Van Kranendonk et al., 2001a). The felsic volcanic Duffer Formation has zircon U–Pb ages of 3471–3463 Ma (Thorpe et al., 1992; McNaughton et al., 1993). The felsic volcanic Panorama Formation has been dated at 3454 ± 1 Ma in the northern part of

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the Marble Bar greenstone belt (Thorpe et al., 1992). Therefore, the depositional age of the studied chert sections is regarded as between 3463 and 3454 Ma. The uppermost sandstone/mudstone of the present study is assigned to the Panorama Formation. Hickman (1983, 1990) considered that the Warrawoona Group was deposited as a tabular succession on sialic crust, but the recognition of subhorizontal shortening by thrust duplication and new geochronological data have led Krapez (1993) and Barley (1993) to propose a more detailed stratigraphic framework for the Pilbara Craton in the plate tectonic regime. More recently, however, a hot and great controversy on the tectonic evolution of the Warrawoona Group

Fig. 1. Location of study area in the geological map of the eastern Pilbara Craton (after Van Kranendonk et al., 2001a). Box represents study area shown in Fig. 2.

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has arisen again; passive, gravity-driven, solid-state diapirism (e.g. Van Kranendonk et al., 2001b) versus thrust-accretion tectonics (plate tectonics; e.g. van Haaften and White, 2001). The depositional settings and environments of the Warrawoona Group are also controversial. Barley (1992, 1993) postulated a volcanic-arc or near-arc setting from the composition of tholeiitic and calcalkaline rocks and from the style of hydrothermal alteration associated with barite-rich Zn–Pb–Cu massive sulfide mineralization and porphyry/epithermal Cu–Zn–Au deposits. Condie (1997) proposed an oceanic plateau setting for tholeiite/komatiite assemblages of the Warrawoona Group. In contrast, Maruyama et al. (1991) and Kitajima et al. (2001) suggested that the tholeiitic basalts were remnants of oceanic crust accreted to an island-arc and/or continental margin arc. Green et al. (2000), however, reported that the Coonterunah and Warrawoona Groups in the Pilgangoora greenstone belt were erupted onto continental basement, and that the basalts assimilated up to 25% crustal material, and hence eliminated volcanic arc, oceanic plateau, and ocean-floor settings. Figure 2 shows a geological map of the study area, based on 1:5000-scale geological mapping. The Duffer Formation, dominated by dacites, is exposed in the east. The Salgash Subgroup is mainly composed of pillowed and massive basalts with minor dolerite and many intercalated chert beds. Pillows consistently indicate a stratigraphic facing to the west. Chert beds that strike NNW-SSE and dip steeply to the E (overturned) have a great variety of field and petrographic features. The Mount Roe Basalt (ca. 2770 Ma) unconformably overlies the Salgash Subgroup with a basal conglomerate, and is separated from the Panorama Formation by a N-S trending high-angle normal fault (Fig. 2). The following is a description of the Salgash Subgroup in ascending order. There is a relatively large-scale shear zone (70–100 m thick) in the altered (silicified/carbonatized) greenstones immediately underlying the Marble Bar Chert Member (A-section in Fig. 2) in the Towers Formation. Although the shear zone is best observed at Marble Bar Pool, it can be traced northwest and southeast in the mapped area. The foliation strikes NNW-SSE and dips steeply to the E, parallel to bedding of cherts. The sheared greenstones are intensely weathered and eroded in some places. Within the ∼20 m thick zone from the

basal contact of the Marble Bar Chert Member, an anastomosing network of massive black silica veins occurs in the sheared altered greenstones. These silica veins have also experienced deformation, although non-deformed thin (<5 cm) veins of later stage cut deformed veins in some places. Damage due to the shearing becomes weak, away from the base of the Marble Bar Chert. The sense of shear of this zone, the amount of displacement, and the timing of formation have yet to be defined. The thickest, more than 50 m, section of multicolored chert (A-section) is the Marble Bar Chert Member in the Towers Formation (Hickman, 1983). Although rather small-scale (20 cm to several meters), a bedding-parallel high strain zone occurs between the A-section cherts and overlying greenstones. Another relatively thick chert section (B), ∼13 m thick, is exposed about 1 km west of the A-section. The B-section is characterized by an association with underlying massive black/gray silica veins. These silica veins diagonally intersect the base of the bedded chert (B-section), but never penetrate up through it. The veins can be traced downwards into the underlying greenstone unit for about 500 m in structural depth where they have an average width of 5–10 m and a maximum width of 30 m. The greenstones have experienced strong hydrothermal alteration around these silica veins. Field evidence strongly suggests that the massive silica veins were originally hydrothermal feeders through which bedded chert-forming components were supplied onto the seafloor. The ∼1 km thick volcanic rocks underlying the B-section is mainly composed of pillowed and massive basalts. The basalts are Fe-rich, low-K tholeiites and have MORB-type geochemical patterns (Nakamura and Kato, 2001). A small-scale, bedding-parallel shear zone (several meters thick) occurs at the top of the B-section. The volcanic unit above the B-section differs from the underlying units, consisting of pillowed and massive komatiitic basalts with local spinifex textures (pyroxene blades up to 10 mm long) and thin (<6 m) bedded cherts. Chert beds seem to have developed on silicified lava flow tops and upward silicification has occurred within each lava flow unit. Only relatively thick (>50 m) lava flows have chert caps, whereas thin lava flows lack overlying bedded cherts. Komatiitic basalt flows are intensely silicified where they are intruded by abundant <100 m long and several

Y. Kato, K. Nakamura / Precambrian Research 125 (2003) 191–243 Fig. 2. Geological map of the Marble Bar area. Sampling sites of six sections (A, B, C1, C2, C3, and T) and the least altered greenstones are shown in this map. ABTH: tholeiitic basalts underlying the B-section, ABKO: komatiitic basalts associated with thin chert sections.

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decimeters wide silica veins. In addition to the A- and B-sections, three chert sections overlying komatiitic basalts were selected for analysis; C1 (6 m thick), C2 (3 m thick) and C3 (4.4 m thick) in ascending stratigraphic order (Fig. 2). The topmost part of the Salgash Subgroup in the study area, stratigraphically correlated with the Panorama Formation (DiMarco and Lowe, 1989), is composed of alternating beds of sandstone and siliceous mudstone >180 m thick (T-section in Fig. 2). The sandstones are persistent in extent and thickness, and each sandstone bed is marked by a sharp base and a less sharp top. The sandstones are poorly sorted and mostly massive, but in places show coarse-tail grading. These field observations indicate that the topmost clastic rocks are of turbiditic origin.

3. Analytical techniques Constituent minerals were determined by microscopic observation, X-ray diffractometry using Cu-K␣ radiation, and electron microprobe analyzer (JEOL8900) operated at 15 kV and 12 nA. Details of the analytical techniques for major, trace, and rare earth elements are given in Kato et al. (1998, 2002). Major elements were analyzed by Rigaku 3270 X-ray fluorescence (XRF) at the Ocean Research Institute, the University of Tokyo. Trace elements and REEs were analyzed using a Yokogawa Analytical Systems PMS-200 inductively coupled plasma mass spectrometry (ICP-MS) at the Geological Survey of Japan (GSJ). Powdered splits (100.00 mg) were dissolved by HNO3 –HF–HClO4 digestion in tightly sealed 12 or 25 ml Teflon PFA screw-cap beakers, heated for 3–7 days on a hot plate under 160–180 ◦ C, then evaporated nearly to dryness. Although sample solutions obtained by dissolution of the residue with 5 ml (1 + 1) HNO3 were partly diluted to 1:1000 as usual, the chert solutions were mostly diluted only to 1:100 to raise the trace element and REE signal intensity of the ICP-MS, while considering the major element matrix. Very high concentrations of some trace elements (e.g. Cr, Zn, and Ba) in some parts of the samples were determined by additional analyses of properly diluted split solutions. Moreover, REE concentrations in some Ba-rich B-section samples were carefully determined by preparing solutions through

the ion exchange method in order to avoid the serious overlap of 135 Ba16 O with 151 Eu. The Eu anomaly is defined quantitatively as Eu/ Eu∗ = EuN /[(SmN )(GdN )]1/2 where Eu∗ is the hypothetical concentration that strictly trivalent Eu would have (Taylor and McLennan, 1985). The subscript “N” indicates chondrite-normalized values. When Eu is depleted or enriched, the Eu/Eu∗ ratio is less than one, or vice versa. The Ce anomaly is defined as Ce/Ce∗ = CeN /[(LaN )(PrN )]1/2 . The degree of light REE (LREE) enrichment relative to heavy REEs (HREEs) is presented as the ratio of chondrite-normalized La to Yb, (La/Yb)N .

4. Samples The A-section is very thick (>45 m) and has red, brown, light green, gray, black, and white layers (Fig. 3). Brown and black cherts are predominant in the lower horizon, light green cherts are present in the middle horizon, and red cherts become more dominant up-section. Each layer varies from 1 mm to 2 cm in thickness. Layering is generally distinct. Black, gray, and light green cherts are mainly composed of microcrystalline quartz, siderite, dolomite, and black aggregates (probably organic compounds). Brown cherts mostly consist of quartz and goethite. Red cherts contain quartz, hematite, and siderite, and white cherts are composed solely of quartz. The B-section is ∼13 m thick and dominated by translucent gray and black cherts. Each layer is mostly 1–2 cm thick. Layering is clear in the lower horizon, but unclear in the middle and upper horizons. Gray and black cherts consist dominantly of quartz and black aggregates or clots, and sometimes contain illite with a trace of barite. EPMA analysis shows that illite generally contains ∼0.2 wt.% Cr2 O3 , and sometimes Ba up to 5 wt.%. Mineralogically, mafic volcanic rocks of the Apex Basalt underlying the B-section (hereafter designated ABTH; Apex Basalt tholeiitic basalts) are dominated by secondary phases of chlorite, muscovite, quartz, carbonate minerals, epidote, actinolite, albite, and opaque minerals such as pyrite. Although relict pyroxene phenocrysts are locally preserved in the least altered massive basalts, actinolite and chlorite mostly replace pyroxene. Other primary phases are comp-

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Fig. 3. Stratigraphic columns of six sections.

letely replaced by secondary minerals. Olivine pseudomorphs after chlorite and carbonate minerals (siderite, ankerite, and calcite) have been frequently recognized in pillowed basalts. Plagioclase phenocrysts and laths in groundmass are entirely replaced by muscovite, epidote, quartz, albite, and fine opaque minerals. The C1-, C2-, and C3-sections are thin (3–6 m) and are mutually similar in stratigraphy and mineralogy. They are mainly composed of light green and black/

gray cherts with uppermost brown or green volcaniclastic cherts. Black/gray cherts are dominant in the lower horizon, and light red cherts are locally present in the middle horizon. These sections are characterized by a predominance of light green cherts and the presence of uppermost volcaniclastic cherts, unlike the A- and B-sections. Each layer varies from 1 mm to 3 cm in thickness, although layering is often unclear. Light green cherts are composed of quartz and illite

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with minor black aggregates or clots. Black/gray cherts consist of quartz and black clots with minor illite, and volcaniclastic cherts consist of quartz, illite, chlorite, rutile, dolomite, and goethite. The C2-section is characterized by the presence of chlorite as a major constituent. Rutile is more significant in the C3-section. Illite contains elevated amounts of Cr (usually 0.5– 1.5 wt.% Cr2 O3 and 6 wt.% at maximum). Chlorite also contains significant Cr (up to 2.6 wt.% Cr2 O3 ). The least altered komatiitic basalts associated with these thin chert sections (ABKO; Apex Basalt komatiitic basalts) were analyzed. Systematic samples from the entire unit of komatiitic basalts were not obtained (Fig. 2), because there is no relatively fresh basalt sample in the upper part of the unit due to intense weathering that is probably related to the basal unconformity of the Mount Roe Basalt. All samples for the present study are characterized by extensive development of secondary chlorite, quartz, carbonate minerals, and actinolite with minor epidote, albite, and opaque minerals. Primary pyroxene is partly overprinted by actinolite and chlorite. Olivine pseudomorphs after chlorite ± actinolite ± carbonate minerals can be recognized in coarse-grained samples. Plagioclase is completely replaced by epidote, albite, and fine opaque minerals. Opaque minerals in the matrix include pyrite and minor sphalerite. The T-section consists of alternating sandstone and siliceous mudstone. Siliceous mudstones are generally gray to light gray, although massive black mudstones rarely occur. Bedding is generally poorly developed, but locally very fine bedding is observed. The mudstones are composed of very fine-grained quartz, illite, and minor rutile. Sandstones are more dominant up-section (Fig. 3), and vary from fine to coarse, mostly 0.1–0.5 mm in diameter, and consist of quartz and illite with small amounts of rutile and zircon in mineralogy. Volcanic rock fragments, volcanic quartz showing a bipyramidal habit and corrosion embayments, and feldspar now replaced by illite are present in the coarse-grained sandstones (see also DiMarco and Lowe, 1989). Sedimentary rock fragments such as black chert are also present in these coarse-grained sandstones. Illite in this section contains approximately 3 wt.% Na2 O and negligible Cr2 O3 (<0.05 wt.%), and differs markedly from that in underlying chert sections: B-section illite contains ∼0.2 wt.% Cr2 O3 , and sometimes Ba up to 5 wt.%;

Cr contents of illite in the C1-, C2-, and C3-sections are much higher (up to 6 wt.% Cr2 O3 ).

5. Geochemistry Table 1 shows the average concentrations of each lithology in the six sections, along with standard deviations (±1σ) for the major elements. Average data for the least altered associated greenstones (ABTH and ABKO) are also included in this table. All data for individual samples (291 cherts, 19 volcaniclastic cherts, 14 sandstones, 12 mudstones, 6 tholeiitic basalts, and 6 komatiitic basalts) are given in the Appendix A–G. 5.1. Major and trace elements 5.1.1. A-section The Marble Bar Chert is mainly composed of SiO2 and Fe2 O3 ∗ , reflecting a ferruginous lithology (Figs. 4 and 5); average SiO2 and Fe2 O3 ∗ contents are 83.5 and 10.2 wt.%, respectively. The cherts in the lower part of the section are generally more ferruginous (Fig. 4). Al2 O3 contents are low (0.6 wt.% on average), suggesting that detrital materials were starved during deposition of the A-section. TiO2 and K2 O contents are also minor (Figs. 4 and 6). MnO contents are well correlated with Fe2 O3 ∗ and the maximum MnO content is 1.26 wt.%. The A-section contains appreciable MgO (up to 3 wt.%) and CaO (up to 1.9 wt.%), derived from carbonates. Manganese is also likely to be mostly contained in carbonates (mainly siderite and dolomite). The A-section cherts have higher P2 O5 (up to 0.17 wt.%), relative to other sections (Fig. 4), and there is a broad correlation between P2 O5 and Fe2 O3 ∗ . Figure 7 shows average trace element concentrations of six sections normalized to post-Archean average Australian shale (PAAS). These trace elements are generally concentrated in detrital components in sediments, although V, Ni, and Y are moderately enriched in hydrothermal iron-rich sediments. In the top figure, average patterns of ABTH and ABKO are shown. For the sake of comparison, the Panorama Formation dacites (PFDA; Barley et al., 1998) and the ca. 3450 Ma Shaw Batholith granitoids (SBG; Bickle et al., 1983, 1993) as well as Archean komatiite (KO), basalt (BA), TTG, and granite (GR) compiled by Condie (1993) are given in this figure. KO and BA are

Table 1 Average compositions of each lithology A-section

B-section

C1-section

Ferruginous chert (n = 60)

Chert (n = 83)

Chert (n = 51)

83.48 ± 14.2 0.04 ± 0.03 0.55 ± 0.50 10.24 ± 9.63 0.311 ± 0.307 0.86 ± 0.93 0.23 ± 0.33 0.46 ± 0.50 0.04 ± 0.02 0.021 ± 0.029 3.97 ± 3.76

Sc (ppm) V Cr Co Ni Cu Zn Rb Sr Y Zr Nb Mo Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Th U

1.52 24.9 18.7 8.61 38.5 10.5 29.2 1.60 10.2 4.47 3.00 0.20 n.d. 56.8 1.38 2.57 0.34 1.45 0.32 0.20 0.31 0.07 0.48 0.10 0.37 0.08 0.42 0.08 0.08 0.02 0.05 0.11

REE Ce/Ce∗ Eu/Eu∗ (La/Yb)N

8.12 0.86 2.61 2.13

n.d.: Not determined.

98.25 ± 1.81 0.05 ± 0.09 0.63 ± 1.00 0.54 ± 1.02 0.004 ± 0.001 0.04 ± 0.02 0.04 ± 0.01 0.08 ± 0.11 0.16 ± 0.26 0.008 ± 0.010 0.36 ± 0.29 2.17 16.2 58.0 1.81 12.4 6.38 11.3 3.34 17.8 2.25 13.3 0.67 0.71 629 1.62 3.38 0.39 1.72 0.36 0.12 0.34 0.06 0.35 0.07 0.22 0.04 0.23 0.03 0.26 0.05 0.22 0.18

± ± ± ±


8.39 0.09 1.22 1.38

8.95 1.00 1.10 6.72

96.52 ± 2.48 0.17 ± 0.13 2.07 ± 1.82 0.17 ± 0.41 0.005 ± 0.001 0.05 ± 0.02 0.04 ± 0.00 0.02 ± 0.02 0.49 ± 0.48 0.005 ± 0.003 0.51 ± 0.21 5.71 66.1 490 3.86 12.6 12.0 58.1 17.2 16.1 1.22 31.1 0.86 0.79 61.0 0.56 1.21 0.15 0.64 0.14 0.04 0.14 0.03 0.16 0.03 0.11 0.02 0.12 0.02 0.21 0.05 0.06 0.26

± ± ± ±

11.9 0.13 0.20 6.58

3.36 1.10 1.02 4.26

72.24 ± 8.97 0.49 ± 0.14 7.80 ± 2.92 8.11 ± 0.79 0.189 ± 0.126 2.22 ± 1.63 2.14 ± 2.82 0.09 ± 0.06 1.49 ± 0.59 0.032 ± 0.008 5.21 ± 4.19 38.2 185 4633 76.7 1508 209 10624 69.8 18.0 7.16 27.4 1.20 0.36 119 0.79 1.93 0.32 1.63 0.60 0.22 0.80 0.19 1.15 0.24 0.73 0.11 0.71 0.11 0.64 0.09 0.11 0.24

± ± ± ±

2.78 0.55 0.17 4.25

9.52 0.91 0.99 0.72

C3-section

Chert (n = 48)

Volcaniclastic chert (n = 5)

96.45 ± 2.51 0.08 ± 0.08 1.56 ± 1.08 0.77 ± 1.09 0.011 ± 0.011 0.19 ± 0.22 0.05 ± 0.011 0.01 ± 0.00 0.26 ± 0.22 0.004 ± 0.004 0.65 ± 0.31 2.80 50.3 379 11.6 51.2 12.8 89.1 9.70 1.77 2.99 7.66 0.54 0.45 41.0 0.55 1.08 0.16 0.82 0.25 0.10 0.34 0.07 0.43 0.09 0.29 0.05 0.27 0.04 0.17 0.03 0.08 0.41

± ± ± ±

2.41 0.07 0.11 0.16

4.55 0.86 1.08 1.86

75.26 ± 7.87 0.55 ± 0.26 7.76 ± 2.74 10.62 ± 3.02 0.109 ± 0.048 2.38 ± 0.94 0.11 ± 0.06 0.04 ± 0.01 0.31 ± 0.25 0.015 ± 0.004 2.88 ± 0.83 30.0 188 1589 62.3 321 323 442 8.57 11.6 58.5 39.0 2.64 0.51 45.3 5.38 10.6 1.73 10.0 3.87 1.56 6.91 1.48 8.60 1.84 5.58 0.82 4.59 0.73 1.19 0.17 0.37 0.27

± ± ± ±

7.91 0.10 0.21 1.61

Chert (n = 49) 95.79 ± 2.94 0.25 ± 0.37 2.30 ± 1.73 0.40 ± 0.94 0.012 ± 0.020 0.07 ± 0.02 0.04 ± 0.01 0.03 ± 0.01 0.61 ± 0.49 0.008 ± 0.006 0.53 ± 0.27 5.53 61.5 665 19.6 44.7 40.7 66.3 18.5 4.83 2.49 22.3 1.72 1.43 34.2 1.54 3.99 0.52 2.35 0.51 0.15 0.39 0.07 0.39 0.08 0.24 0.04 0.23 0.03 0.33 0.10 0.18 0.10

63.7 ± 34.1 0.76 ± 0.12 0.95 ± 0.07 0.98 ± 0.59

10.5 ± 17.9 1.09 ± 0.31 1.17 ± 0.22 3.66 ± 2.26

T-section Volcaniclastic chert (n = 10) 79.86 ± 4.44 1.43 ± 0.47 10.26 ± 2.53 2.85 ± 3.17 0.156 ± 0.192 0.27 ± 0.12 0.14 ± 0.11 0.05 ± 0.02 2.92 ± 0.74 0.042 ± 0.009 2.04 ± 0.61 26.9 236 951 35.0 79.5 170 13.1 104 19.3 10.6 91.2 10.8 1.29 124 7.35 19.5 2.58 11.3 2.64 0.83 2.22 0.44 2.36 0.48 1.35 0.21 1.26 0.18 2.64 1.15 1.46 0.51 52.7 ± 38.1 1.00 ± 0.13 1.07 ± 0.07 3.55 ± 1.59

Sandstone (n = 14) 85.29 ± 3.66 0.28 ± 0.08 10.14 ± 2.58 0.35 ± 0.40 0.005 ± 0.002 0.12 ± 0.11 0.09 ± 0.02 0.40 ± 0.19 1.66 ± 0.67 0.010 ± 0.014 1.68 ± 0.34 5.18 33.8 25.3 1.63 20.1 9.48 17.8 52.9 103 7.94 88.3 3.50 0.47 424 11.9 23.0 2.45 9.27 1.65 0.61 1.52 0.25 1.27 0.25 0.74 0.12 0.69 0.11 1.93 0.41 2.16 0.77 53.7 ± 9.69 0.99 ± 0.04 1.18 ± 0.16 12.6 ± 5.45

Greenstone Mudstone (n = 12) 86.70 ± 3.68 0.23 ± 0.12 9.25 ± 2.59 0.09 ± 0.064 0.004 ± 0.001 0.04 ± 0.03 0.08 ± 0.02 0.33 ± 0.22 1.82 ± 0.61 0.020 ± 0.03 1.44 ± 0.33 4.55 22.1 15.2 0.76 6.03 9.08 8.18 56.0 110 8.72 88.3 4.46 0.65 506 15.7 31.3 3.18 11.2 1.94 0.61 1.75 0.27 1.51 0.30 0.93 0.15 0.98 0.14 2.31 0.68 5.85 1.83 69.9 ± 38.1 1.02 ± 0.02 1.08 ± 0.20 10.5 ± 1.93

Tholeiitic basalt (ABTH) (n = 6) 50.05 ± 2.39 0.80 ± 0.23 15.73 ± 0.76 12.46 ± 1.30 0.172 ± 0.011 8.88 ± 2.45 9.40 ± 2.26 1.75 ± 0.33 0.18 ± 0.07 0.059 ± 0.018 5.50 ± 2.68 39.5 325 316 76.1 281 237 111 6.96 96.3 15.3 42.5 1.86 0.21 49.5 2.34 6.18 0.92 4.87 1.61 0.55 n.d. 0.40 2.55 0.55 1.70 0.27 1.67 0.25 1.11 0.15 0.22 0.05 26.0 ± 6.98 0.98 ± 0.03 0.92 ± 0.07 0.92 ± 0.04

Komatiitic basalt (ABKO) (n = 6) 51.38 ± 1.78 0.54 ± 0.03 10.89 ± 0.88 11.34 ± 1.06 0.178 ± 0.016 12.52 ± 2.92 10.67 ± 0.62 2.29 ± 1.04 0.06 ± 0.03 0.037 ± 0.005 3.26 ± 0.86 39.4 203 1196 63.5 326 81.8 69.8 0.59 144 12.8 29.0 0.90 0.52 12.9 1.64 4.46 0.68 3.62 1.18 0.43 n.d. 0.33 1.96 0.42 1.28 0.21 1.30 0.20 0.76 0.06 0.16 0.04

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SiO2 (wt%) TiO2 Al2 O3 Fe2 O3 ∗ MnO MgO CaO Na2 O K2 O P2 O5 LOI

C2-section Volcaniclastic chert (n = 4)

19.4 ± 2.29 0.98 ± 0.02 0.94 ± 0.12 0.83 ± 0.15

REE values are given as sum totals of fourteen REEs after unfilled REE data are supplemented by interpolation.

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Fig. 4. Stratigraphic variations of major elements of six sections.

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Fig. 5. SiO2 vs. Fe2 O3 ∗ (left) and SiO2 vs. Al2 O3 (right) for six sections. IT: illite-trend; CT: chlorite-trend.

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Fig. 6. TiO2 vs. Al2 O3 (left) and K2 O vs. Al2 O3 (right) for six sections. Symbols are the same as those in Fig. 5.

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Fig. 7. Average trace element patterns of six sections normalized to post-Archean average Australian shale (PAAS). Normalizing PAAS values are from Taylor and McLennan (1985). ABTH: tholeiitic basalts underlying the B-section, ABKO: komatiitic basalts associated with thin chert sections, PFDA: Panorama Formation dacites (Barley et al., 1998), SBG: Shaw Batholith granitoids (Bickle et al., 1983, 1993). Archean komatiite (KO), basalt (BA), TTG, and granite (GR) compiled by Condie (1993) are also shown in this figure.

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more enriched in Sc, V, Cr, Co, and Ni than GR and TTG. Moreover, KO is characterized by two distinct peaks of Cr and Ni. GR and TTG, on the other hand, exhibit enrichment of Zr, Nb, Hf, and Th relative to KO and BA. The patterns of ABTH and ABKO are generally similar to that of BA and KO, respectively, although enrichment of Cr and Ni in ABKO is less striking than KO. There is a general similarity among PFDA, SBG, GR, and TTG, although PFDA is slightly more enriched in Sc, V, Cr, and Ni than others. Although the normalized values are one order of magnitude lower due to dilution by chert, the normalized pattern of the A-section is similar to that of ABTH, suggesting a minor contribution of ABTH detritus in the A-section. There are positive correlations between Fe2 O3 ∗ and several trace elements in the A-section. Correlation coefficients of these trace elements with Fe2 O3 ∗ are 0.58 (P2 O5 ), 0.57 (V), 0.77 (Zn), and 0.55 (Y), which are statistically significant on a 99% confidence level. The positive correlation between Fe2 O3 ∗ and these trace elements is not recognized in other sections. 5.1.2. B-section Although the B-section consists exclusively of SiO2 (98.25 ± 1.81 wt.%), there are two distinct trends seen in Fig. 5. One trend is indicated by samples containing Fe2 O3 ∗ (up to 6.4 wt.%), identical to the A-section trend. The other trend comprises samples containing Al2 O3 up to 4.0 wt.% (illite-trend shown as IT in Fig. 5). Illite is responsible for the containment of Al2 O3 , although present locally in the lower and middle horizons (Fig. 4). Although K2 O is present only in trace amounts (Fig. 6), some samples exhibit an excellent correlation between Al2 O3 and K2 O caused by the presence of illite. In contrast to the A-section, the B-section contains only trace MnO, MgO, and CaO (Fig. 4 and Table 1). The B-section cherts are locally enriched in Ba (up to 4330 ppm), which is attributed to Ba-bearing illite and barite. The trace element pattern is generally similar to that of ABTH, like the A-section, although there is slight enrichment of Cr and Ni (Fig. 7). 5.1.3. C1-, C2-, and C3-sections The cherts of the C1-, C2-, and C3-sections are chiefly composed of SiO2 and Al2 O3 (Fig. 5). The average Al2 O3 contents of the cherts (except for the

volcaniclastic cherts) are 2.1 wt.% (C1), 1.6 wt.% (C2), and 2.3 wt.% (C3). The SiO2 contents gradually decrease and other oxides increase upward in these sections (Fig. 4). These compositional variations reflect a stratigraphic change showing that cherts containing volcaniclastic debris are abundant in the upper horizon of each section. The Fe2 O3 ∗ contents are mostly minor, although the volcaniclastic cherts contain significant amounts of Fe2 O3 ∗ (Figs. 4 and 5). The C2-section cherts often have moderate amounts of Fe2 O3 ∗ in the middle and upper horizons, and MgO contents are well correlated with Fe2 O3 ∗ (Fig. 4). The source of iron and magnesium for the C2-section is mainly chlorite. TiO2 enrichment is especially striking in the C3-section (Figs. 4 and 6), reflecting the greater amount of rutile. The K2 O contents of the C1- and C3-sections correlate well with Al2 O3 , indicating that both Al and K are derived from illite (Fig. 6). Some cherts of the C2-section show this K2 O–Al2 O3 correlation, but other cherts and volcaniclastic cherts are more depleted in K2 O due to the presence of chlorite as a major aluminous mineral (chlorite-trend shown as CT in Fig. 5). Significant amounts of MgO (up to 4.6 wt.%) and CaO (up to 6.0 wt.%) are contained in the volcaniclastic cherts of the C1-section (Table 1), caused by dolomite. The C1-, C2-, and C3-sections are commonly characterized by Cr and Ni enrichment (Fig. 7), reflecting a significant contribution of volcanic detritus derived from komatiitic basalts (ABKO). The Cr concentrations of the C1 cherts vary from 10 to 2500 ppm. The C1 volcaniclastic cherts exhibit significantly larger enrichment in Cr (4600 ppm on average) than komatiitic basalts (Table 1), probably due to concentration of Cr-rich illite (up to 6 wt.% Cr2 O3 ). Volcaniclastic cherts are more enriched in Cr and Ni than cherts in the C1- and C2-sections (Table 1). However, the C3 cherts locally have more Cr (up to 6120 ppm) than the overlying volcaniclastic cherts (Appendix E). Enrichment of Zr, Nb, Hf, and Th is recognized in the C3-section compared with the C1- and C2-sections, implying a relatively larger contribution of felsic rocks such as GR, TTG, PFDA, and SBG (Fig. 7). 5.1.4. T-section Mudstones and sandstones are mainly composed of SiO2 , Al2 O3 , and K2 O (Figs. 4–6). The Fe2 O3 ∗

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Fig. 8. Average major element patterns of siliceous mudstone and sandstone in the T-section, normalized to PFDA. Abbreviations and the data sources are the same as those in Fig. 7. PFVS: volcaniclastic sandstones and siltstones in the Panorama greenstone belt (Cullers et al., 1993).

contents are close to zero, while Al2 O3 contents vary from 6.3 to 15.8 wt.% (Fig. 5). TiO2 and K2 O contents generally correlate with Al2 O3 (Fig. 6). The average K2 O content of the T-section is 1.73 wt.%. MnO, MgO, and CaO contents are very minor (Table 1). Clastic rocks of the T-section show a broad correlation between Na2 O and Al2 O3 , with the average Na2 O content of 0.37 wt.%. Major oxide abundances normalized to the PFDA are given in Fig. 8. Remarkable depletion in Fe2 O3 ∗ , MnO, MgO, CaO, Na2 O, and P2 O5 suggests the removal of these labile elements relative to immobile constituents (Al and Ti) during chemical weathering and alteration. Potassium metasomatism (Fedo et al., 1997) may be responsible for relative enrichment of K2 O in the T-section. The patterns of siliceous mudstones and sandstones are generally consistent with that of volcaniclastic sandstones and siltstones in the Panorama greenstone belt (hereafter designated PFVS; Cullers et al., 1993). However in more detail, the clastic rocks of the present study are more enriched in CaO and Na2 O, and more depleted in Fe2 O3 ∗ and MnO than PFVS. The trace element pattern of the T-section is broadly similar to that of GR, TTG, PFDA, and SBG (Fig. 7).

and is highly variable (1.31–6.61,
average = 2.61). The total REE concentration ( REE) is generally low, and ranges widely from 0.2 to 46.2 ppm, reflecting two contrasting lithologies (pure chert and iron-rich chert). The Ce/Ce∗ ranges from 0.60 to ∗ 1.04, indicating negative Ce anomalies.
The Eu/Eu is well negatively correlated with REE, whereas
the Ce/Ce∗ is positively correlated with REE (Fig. 10). The positive Eu anomalies decrease with
increasing REE, and simultaneously the magnitude of negative Ce anomalies gradually decreases toward zero anomaly (Ce/Ce∗ = 1.0). The variation of (La/Yb)N is striking (0.32–6.3), and the REE patterns vary greatly from HREE- to LREE-enriched. Furthermore, the REE patterns of individual samples
become more LREE-enriched with increasing REE (Fig. 9). In other noticeable correla
sections these tions among REE, Eu/Eu∗ , Ce/Ce∗ , and (La/Yb)N are not seen,
although a weak correlation between Eu/Eu∗ and REE is recognized in the C2 and C3 cherts (Fig. 10). The Eu anomalies are not correlated with Al2 O3 (Fig. 11), suggesting that the positive Eu anomalies in the A-section are not attributable to an aluminum-bearing phase such as feldspar.

5.2. Rare earth elements

5.2.2. B-section The B-section cherts generally show moderate LREE-enrichment, and several samples with low
REE show flat REE patterns (Fig. 9). The range
of REE is about the same as that of the A-section (average = 8.95 ppm), but relative HREE-depletion is obvious in the B-section, resulting in a higher

5.2.1. A-section Chondrite-normalized REE patterns of each section are given in Fig. 9. The A-section is characterized by the presence of conspicuous positive Eu anomalies. The Eu/Eu∗ is much higher than that of other sections,

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Fig. 9. Chondrite-normalized REE patterns of six sections in stratigraphic order. Broken lines in the C1-, C2-, and C3-sections represent volcaniclastic cherts. The REE patterns of the T-section are plotted separately based on their lithology. Abbreviations and the data sources are the same as those in Fig. 7. Chondrite values are after Masuda et al. (1973) and Masuda (1975).

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Fig. 10. Eu/Eu∗ vs.




REE (left) and Ce/Ce∗ vs.




REE (right) for six sections. Symbols are the same as those in Fig. 5.

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(La/Yb)N than the A-section (the average (La/Yb)N is 6.72, greater than that of the A-section = 2.13). Eu anomalies are generally small (average Eu/Eu∗ = 1.1), but positive anomalies up to 1.84 are recognized locally in the lower and upper horizons. The Ce/Ce∗ is 1.00 ± 0.13 on average, but varies from 0.81 to 1.81, with more than two thirds of the samples having subtle negative Ce anomalies (Fig. 10).

Fig. 11. Eu/Eu∗ vs. Al2 O3 for six sections. Symbols are the same as those in Fig. 5.

5.2.3. C1-, C2-, and C3-sections The REE patterns of the
C1-section are slightly LREE-enriched (Fig. 9), and REE contents are very low (3.81 ppm on average, maximum of 14.3 ppm). Weak positive Eu anomalies are seen in a few chert samples in the lower part of the section (up to 1.68), although average Eu/Eu∗ is 1.02. The Ce/Ce∗ values are variable (0.45–3.32; Fig. 10). The uppermost volcaniclastic cherts, shown as broken lines in Fig. 9, have flat or slightly LREE-depleted patterns with chondrite-normalized values of 2–4. Nearly flat patterns are predominant in the C2-section (Fig. 9), but (La/Yb)
N varies from 0.34 to 8.19 (average = 1.78). REE is 10.1 ppm on average, which is significantly higher than
that of the C1-section and is mainly due to the high REE contents of the uppermost volcaniclastic cherts (average = 63.7 ppm). Weak positive Eu anomalies (up to 1.57) are clearly recognized in the HREE-enriched patterns from the lower to middle parts of the section. Ce/Ce∗ ranges from 0.56 to 1.08, and a negative Ce anomaly is predominant, as in the A-section (Fig. 10). The uppermost cherts and volcaniclastic cherts are marked by HREE-enriched patterns and negative Ce anomalies (Fig. 9). The C3-section is characterized by predominance of smooth and distinct LREE-enriched patterns with high

REE (Fig. 9). The average REE in the whole section is 17.7 ppm. The lower part of this section shows flat or subtly LREE-enriched patterns with positive Eu anomalies (1.92 at maximum). The average Eu/Eu∗ is 1.15, indicating a small positive Eu anomaly. Several samples have significant positive Ce anomalies (up to 2.8), but most samples show no anomaly (∼1). The volcaniclastic cherts in the uppermost horizon
show noticeable LREE-enriched patterns with high REE in spite of a significant contribution from komatiitic basalt (striking enrichment of Cr and Ni). The average (La/Yb)N of the volcaniclastic chert reaches 3.55.

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These LREE-enriched patterns clearly differ from the LREE-depleted patterns of the volcaniclastic cherts in the C1- and C2-sections. 5.2.4. T-section The REE patterns of the T-section are strikingly LREE-enriched and are similar to that of GR, TTG, PFDA, and SBG except for a positive Eu anomaly (Fig.
9). Average (La/Yb)N is 11.6 ± 4.26.
Average REE is 61.2 ± 27.5 ppm, but the REE contents reach a maximum of 147 ppm. Eu/Eu∗ shows a moderate variation from 0.75 to 1.54, averaging 1.13 ± 0.18 (Fig. 10). Although more than half the samples have positive Eu anomalies, relatively large negative Eu anomalies (0.75 and 0.81) exist in two mudstone
samples. These two mudstones have the highest REE contents (144 and 147 ppm). There is a positive correlation between Eu/Eu∗ and Al2 O3 in all samples (Fig. 11). This suggests that the positive Eu anomalies of the T-section are chiefly due to an aluminum-bearing phase, which may have been originally detrital feldspar. The Ce/Ce∗ values are near unity, indicating the absence of a Ce anomaly (Fig. 10).

6. Discussion 6.1. Origins and depositional environments 6.1.1. A- and B-sections: hydrothermal sediment at a MOR There are marked similarities in geochemical signatures between the A-section cherts and modern hydrothermal iron-rich particulates and sediments from MOR. It is well known that certain elements such as P, V, and As in modern hydrothermal plume particulates exhibit excellent linear correlations with Fe (e.g. Trocine and Trefry, 1988; German et al., 1991; Feely et al., 1994). Particulate Y and Zn concentrations also increase with an increase of particulate Fe, although Y and Zn show positive and negative departures from simple linear correlations due to subsequent uptake and release after particulate formation, respectively (German et al., 1991). Positive correlations of these elements with Fe in the particulate are essentially inherited in the iron-rich sediments (e.g. Marchig and Gundlach, 1982), although minor redistribution of these elements may occur during post-depositional

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processes (Schaller et al., 2000). Thus, the fact that P, V, Zn, and Y correlate with Fe in the A-section cherts, although rather scattered, may imply that the A-section ferruginous cherts are analogous to modern hydrothermal iron-rich sediments deposited at MOR. A similarity in correlation among REE indices is far more striking. Ruhlin and Owen (1986) and Olivarez and Owen (1989) demonstrated that the Eu anomalies of the East Pacific Rise hydrothermal iron-rich
sediments decrease with increasing REE, and simultaneously the magnitude of negative Ce anomalies decrease. An identical and more distinct correlation was reported for iron oxyhydroxide suspended particulates from the TAG hydrothermal vent field (German et al., 1990). Hydrothermal iron-rich particulates scavenge REEs partly from hydrothermal solutions at an early stage during particulate formation, and mostly from seawater at a later stage of formation. The continuous REE overprinting from seawater results in attenuation of the positive Eu anomaly which is a distinctive hydrothermal signature (Michard and Albarède, 1986;
Campbell et al., 1988), and also in enhancement of REE. As the magnitude of negative Ce anomalies in seawater near ridge crests is larger than that of open-ocean seawater (Klinkhammer et al., 1983), sediments and suspended particulates near the crest are characterized by larger negative Ce anomalies. Subsequent scavenging and overprinting of seawater-derived REEs bring about a reduction in the size of negative Ce anomalies closer to that
of open-ocean seawater, along with an increase of REE. Moreover, continuing preferential uptake of LREEs relative to HREEs results in increasing LREE/HREE with increasing
REE in the sediments and particulates (Piper and Graef, 1974; German et al., 1990). This type of correlation in REE indices is absent in modern iron-rich hydrothermal sediments from volcanic arcs (Savelli et al., 1999), back-arc seamounts (Binns et al., 1993), and hot-spots (Hekinian et al., 1993), although the cause of its absence in these settings seems uncertain. The above-mentioned correlation among REE indices of modern MOR hydrothermal sediments and particulates is identical to that of the A-section cherts, although absolute values differ considerably. Eu/Eu∗ in the A-section decreases from 6.6 to 1.3 and Ce/Ce∗ increases from 0.6 to 1.0 (i.e. the magnitude of the
negative Ce anomalies decreases) as REE and LREE/HREE increase (Fig. 10). These geochemical

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lines of evidence from the REEs and the trace elements mentioned above suggest that the A-section ferruginous cherts are the counterpart of modern MOR hydrothermal iron-rich sediments. Enrichment of silica in the A-section cherts relative to the modern iron-rich sediments is most likely attributed to the widespread inorganic precipitation of hydrothermal silica in the Archean ocean. Hydrothermal silica deposits are rarely formed in the modern ocean system, because amorphous silica is undersaturated due to the active formation of siliceous microorganisms. In contrast, hydrothermal silica is presumed to have been easy to precipitate in the Archean ocean owing to lack of silica-fixing microorganisms. It should be noted that the A-section cherts have negative Ce anomalies. Modern hydrothermal sediments exhibit huge negative Ce anomalies inherited from oxygenated deep-sea water by REE scavenging. The fact that indisputable negative Ce anomalies, although rather small, are present in the A-section could mean that some sort of Ce oxidation mechanism, even though on a very small scale, already existed in the Early Archean ocean, and hence that the ocean and atmosphere were not entirely anoxic at least partly. Therefore, we believe that the atmospheric oxygen level in the Early Archean was between the levels defined by two extremely contrasting models (anoxic model by Holland, 1999 versus fully oxic model by Ohmoto, 1997). Some parts of the B-section cherts are ferruginous and are identical to the A-section cherts (Fig. 5). The overall trace element pattern of the B-section is also similar to that of A-section (Fig. 7). However, there are several definite differences in geochemical features between the A- and B-sections. Among them are differences in the Eu anomaly, Ba concentration, and Fe2 O3 ∗ content which have important implications for their origin. Sverjensky (1984) demonstrated that the Eu2+ /Eu3+ redox equilibria in solutions strongly depend on temperature. Europium in high-T (>250 ◦ C) aqueous solutions is dominated by Eu2+ , and thus conspicuous positive Eu anomalies are present in hydrothermal solutions above that temperature. In contrast, at low-T (100–150 ◦ C), hydrothermal solutions with oxidation states near that of magnetite/hematite or sulfide/sulfate could contain significant amounts of both Eu2+ and Eu3+ , resulting in weak positive Eu anomalies. Therefore, it is most likely that the differ-

ence in the magnitude of the positive Eu anomalies between the A- and B-sections reflects differences in the temperatures of the hydrothermal solutions responsible for their formation. The A-section cherts with remarkable positive Eu anomalies were derived from high-T (greater than 250 ◦ C) hydrothermal solutions, whereas the B-section cherts with subtle positive Eu anomalies were derived from low-T (much less than 250 ◦ C) solutions. The presence of barite in the B-section may also be taken as evidence for origin from a low-T hydrothermal solution. In modern hydrothermal systems at MOR, opal-barite rocks at the East Pacific Rise, 21◦ N are regarded as products of low-T (∼100 ◦ C) white smokers (Haymon and Kastner, 1981). Moreover, the formation temperature of barite-silica chimneys from the Sumisu Rift was estimated by Urabe and Kusakabe (1990) to be less than 150 ◦ C. Experimental studies of basalt–seawater interaction at elevated temperatures have indicated that dissolved iron in hydrothermal solutions interacting with basalt increases exponentially with increasing temperature (Seyfried and Janecky, 1985). Thus, high-T hydrothermal solutions enriched in iron most likely produce iron-rich hydrothermal sediments such as the A-section cherts. In contrast, the B-section cherts were likely precipitated from a lower-T solution less enriched in iron. The detrital component is very minor in both the A- and B-sections. The trace element pattern of the B-section has weak positive Cr and Ni peaks and somewhat resembles those of the C1-, C2-, and C3-sections. Cr∗ is plotted against Sc∗ in order to constrain the source of detrital component more tightly (Fig. 12). Sc and Cr distributions in the sediments have been proposed as good indicators of the detrital source (e.g. Taylor and McLennan, 1985; Condie and Wrokiewicz, 1990). Clearly, the detrital components in the A- and B-sections are estimated to be mainly derived from the tholeiitic basalts underlying the B-section (ABTH). Greenstones underlying the B-section (ABTH) are Fe-rich, low-K tholeiites (SiO2 = 47–53 wt.%, Fe2 O∗3 = 11–14 wt.%, MgO = 7–12 wt.%, K2 O = 0.1–0.3 wt.%; Table 1). On a MORB-normalized trace element variation diagram (Fig. 13), the ABTH exhibits a relatively flat pattern with elemental abundances close to unity except for highly mobile elements such as K, Rb, and Ba. Chondrite-normalized

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Fig. 12. Cr∗ vs. Sc∗ . Cr∗ is the recalculated Cr concentration of each sample assuming that each has 15 wt.% Al2 O3 as a detrital component. If a given sample has 0.3 wt.% Al2 O3 and 8 ppm Cr, Cr∗ is recalculated as 400 ppm (8 × 15/0.3 = 400). The symbol and bar represent an average and range of ±1σ, respectively. Abbreviations and the data sources are the same as those in Figs. 7 and 8.

REE patterns of the ABTH are flat with slight depletion of LREEs. The ABTH plots in the modern N-MORB field on the Th–Hf–Ta and Zr–Nb–Y discrimination diagrams (Fig. 14). These geochemical constraints strongly suggest that the ABTH is MORB in origin. Although it is known that some greenstones in the Warrawoona and Coonterunah Groups are contaminated by crustal material (Condie, 1994; Green et al., 2000), geochemical features of the ABTH are quite different from those of the contaminated tholeiitic basalts (Figs. 13 and 15). Particularly, low Th/Ta, (La/Yb)N , and (La/Sm)N , and high Nb/U of the ABTH clearly indicate considerably low or no crustal contamination (Fig. 15). The absence of continental detritus and the small contribution of MORB-derived detritus indicate that the depositional site of the A- and B-sections was truly oceanic and remote from a continent or island arcs composed of felsic plutonic/volcanic rocks. The massive silica veins associated with the B-section were hydrothermal feeders producing the B-section cherts on the seafloor, emplaced up normal fault zones

in the MOR. In conclusion, these geochemical and geological lines of evidence consistently indicate that the A- and B-sections were in situ high-T and low-T hydrothermal sediments at a MOR, respectively. 6.1.2. C1-, C2-, and C3-sections: hydrothermal sediment at a hot-spot The field occurrence, mineral assemblages, and geochemical signatures of the C1-, C2-, and C3-sections are essentially identical to each other. Due to the presence of chlorite, the C2-section has a somewhat different major element composition. The existence of weak positive Eu anomalies in the lower horizons of the three chert sections suggests that these sections were essentially derived from a low-T hydrothermal solution precipitating silica. The remarkable enrichment of Cr and Ni indicates that these sections have a significant detrital component of komatiitic basalt (Figs. 7 and 12). The associated komatiitic basalts (ABKO) are characterized by relatively high MgO contents (10–16 wt.%) with elevated Cr (up to 1700 ppm) and Ni (up to 480 ppm) contents

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Fig. 13. MORB-normalized trace element patterns (left) and chondrite-normalized REE patterns (right) for tholeiitic (ABTH) and komatiitic basalts (ABKO). Shaded area represents the range of tholeiitic and komatiitic basalts from the Coonterunah and Warrawoona Groups in the Pilgangoora greenstone belt (Green et al., 2000). Normalizing MORB values are after Sun and McDonough (1989) and chondrite values are after Masuda et al. (1973) and Masuda (1975).

(Table 1). The SiO2 and Fe2 O3 ∗ contents are 49–53 and 10–13 wt.%, respectively. MORB-normalized trace element patterns of the ABKO are relatively flat, and chondrite-normalized REE patterns are generally flat or slightly LREE-depleted (Fig. 13). These geochemical features are different from those of the contaminated komatiitic basalts of Green et al. (2000). The ABKO has relatively low Nb/U (Fig. 15) due to slight depletion of Nb (Fig. 13). However, the (La/Sm)N does not increase with decreasing Nb/U, and hence data from the ABKO do not fall within the mixing trajectory between the least contaminated basalt (LB) and granodiorite (GD) defined by Green et al. (2000; Fig. 15a). This geochemical feature might imply that low Nb/U of the ABKO was not attributed to crustal contamination. In fact, significantly low Th/Ta and (La/Yb)N of the ABKO suggest that the maximum level of crustal contamination was less than 1% (Fig. 15b).

The striking differences between the cherts associated with tholeiitic basalts (A- and B-sections) and the cherts associated with komatiitic basalts (C1-, C2-, and C3-sections) suggest that the styles of hydrothermal activity responsible for each type of chert may have differed fundamentally. The interaction between the upper parts of the komatiitic basaltic lava flow and seawater at the time of eruption was probably responsible for the intense silicification of flow-tops (upward silicification). Large-scale eruptions introduced considerable volumes of silica into the hydrothermal solution, giving rise to formation of overlying chert on the seafloor. The hydrothermal activity waned gradually after volcanism, leading to an increased contribution of volcanic detritus and thus forming volcaniclastic chert in the uppermost horizon. The lower temperature of the hydrothermal solution, and the bulk-rock chemistry of komatiitic basalt that is more magnesian and less ferrous than Fe-rich, low-K tholeiites may have

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Fig. 14. (a) Th–Hf–Ta and (b) Zr–Nb–Y discrimination diagrams for tholeiitic basalts underlying the B-section (ABTH). Compositional fields are after (a) Wood (1980) and (b) Meshede (1986). N-MORB: normal mid-oceanic ridge basalt, E-MORB: enriched mid-oceanic ridge basalt, WPT: within-plate tholeiitic basalt, WPA: within-plate alkaline basalt, VAB: volcanic-arc basalt, CAB: calc-alkaline basalt, IAT: island-arc tholeiitic basalt.

contributed to the formation of Si-dominant chert on the komatiitic basalts. The thick cherts of the A- and B-sections were likely formed by strong hydrothermal activity at a divergent boundary where large-scale hydrothermal circulation was developed through normal fault zones. In contrast, the much thinner silica-dominant chert sections (C1, C2, and C3) were most likely formed by

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Fig. 15. (a) (La/Sm)N vs. Nb/U and (b) Th/Ta vs. (La/Yb)N . Data for tholeiitic and komatiitic basalts of the Coonterunah and Warrawoona Groups in the Pilgangoora greenstone belt are from Green et al. (2000). Mixing models of (a) least contaminated basalt and local granodiorite and (b) primitive mantle (solid line) or depleted mantle (broken line) and Archean upper continental crust show estimates of the percentage contamination. Mantle and crustal end-members: (a) GD: granodiorite, LB: least contaminated basalt (Green et al., 2000), (b) UC: Archean upper continental crust, DM: depleted mantle, PM: primitive mantle, EM1 and EM2: enriched mantle, HIMU: high-mu mantle (Zindler and Hart, 1986; Sun and McDonough, 1989; Weaver, 1991; Condie, 1993, 1997).

relatively weak and low-T hydrothermal activity in a different site where such normal fault zones were lacking. The plausible candidate for depositional settings for these thin chert sections overlying the komatiitic basalts may be a hot-spot. It is generally accepted that Archean komatiites erupted at a hot-spot (e.g. Jarvis

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and Campbell, 1983; Campbell et al., 1989; Storey et al., 1991). There are significant differences in lithological assemblage and geochemical features between the tholeiitic basalt (ABTH) and komatiitic basalt (ABKO) units, as mentioned above. Furthermore, the presence of shear zone between them (Fig. 2) suggests that depositional settings and environments of these different units might have been separated spatially. 6.1.3. T-section: clastic sediment at a convergent boundary All geochemical data presented for the T-section indicate that the bulk-rock geochemistry is quite similar to that of felsic plutonic/volcanic rocks. The trace element pattern of the T-section is characterized by relative enrichment of Zr, Nb, Hf, and Th, which is identical to that of felsic rocks such as GR, TTG, PFDA, and SBG (Fig. 7). The REE pattern is remarkably LREE-enriched and similar to that of these felsic rocks (Fig. 9). Significantly, Archean shale is enriched in Cr and Ni and has a significant contribution from komatiite and basalt (e.g. Taylor and McLennan, 1985; Wronkiewicz and Condie, 1987). However, the T-section mudstones are anomalously depleted in Cr and Ni (Fig. 12 and Table 1), in spite of the extensive development of komatiitic basalts underlying the T-section. This suggests that the komatiitic component did not contribute substantially to the T-section. Moreover, the T-section is much more depleted in Cr than PFVS and PFDA (Fig. 12). Because the similar removal of mobile elements due to weathering and alteration is recognized in the T-section and PFVS (Fig. 8), more striking depletion of Cr in the T-section is not attributed to the effects of weathering and alteration. This geochemical constraint indicates that a contribution from dacitic rocks was not significant in the T-section and more felsic rocks were source materials for the T-section. The positive Eu anomalies in more than half the samples were probably caused by local accumulation of detrital feldspar (now converted to illite). This type of positive Eu anomaly in clastic rocks has been reported for some Archean shales (e.g. Nance and Taylor, 1977; Kato et al., 1998). The significant negative Eu anomalies of two siliceous mudstones having
high Th/Sc and the highest REE suggest that differentiated and fractionated granitoids and/or their volcanic equivalents were exposed (erupted) at least

partly in the Early Archean Pilbara. The present result supports that recycling of differentiated continental crust had already occurred on the Earth’s surface in the Early Archean (Buick et al., 1995; Green et al., 2000). In summary, the siliceous mudstones and sandstones were probably deposited as turbidites near an island arc and/or continental margin arc where differentiated granitoids and/or felsic volcanic rocks supplied clastic materials. 6.2. Implications for Early Archean tectonics The sedimentary rocks in the present study exhibit a great variety of modes of occurrence and bulk-rock chemistry, reflecting their varied depositional environments. Both the A- and B-sections are most likely to have been deposited at MOR. The geochemical differences between these two sections are probably caused by the differences in the nature of the chert-forming hydrothermal solutions (e.g. temperatures). Intense hydrothermal activity at MOR brought about voluminous, high-T, iron-rich hydrothermal solutions giving rise to the formation of the very thick ferruginous chert section (A). Relatively weak and low-T hydrothermal activity was responsible for the thinner, Si-dominant, Ba-containing chert section (B). Geochemical evidence from the greenstones associated with these thick cherts consistently demonstrates that the greenstones are of MORB origin and entirely oceanic without contamination by continental crust (Fig. 15). This finding is not inconsistent with geologic evidence that the tholeiitic basalt unit has a tectonic contact with the underlying felsic succession (Duffer Formation; Fig. 2). These geochemical and geological constraints indicate that the Duffer Formation and Apex Basalt were not deposited continuously, and that the tholeiitic basalt unit in the lower Apex Basalt is a remnant of Early Archean oceanic crust. It is assured, however, that the Warrawoona basalts in the Pilgangoora greenstone belt were erupted onto pre-3.5 Ga continental basement (Green et al., 2000). Consequently, it is reasonably considered that tholeiitic basalts in the Warrawoona Group have varied sources and tectonic settings. The C1-, C2-, and C3-sections may have been deposited on an oceanic plateau or islands derived from hot-spot volcanism. Komatiitic basalts intercalated in these chert sections have bulk-rock chemistry much

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different from that of the Fe-rich, low-K tholeiites associated with the A- and B-sections. The distinctive chemistry of the komatiitic basalts is reflected in the bulk-rock chemical compositions of the intercalated

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cherts by inclusion of volcanic detritus (e.g. extreme enrichment of Cr and Ni). Komatiitic basalt volcanism and subsequent silica precipitation occurred repeatedly at a hot-spot volcano, and therefore the C1-, C2-,

Fig. 16. (La/Yb)N vs. Th/Sc. Felsic end-member (avg SBG∗ ) is the average of five differentiated felsic rocks (Sample No. D42-1, A195-1, C137-20, C137-37, A242-1) from Bickle et al. (1983). Numbers along line represent felsic rock-mixing ratios. Abbreviations and the data sources are the same as those in Figs. 7 and 8.

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Fig. 17. Schematic diagram showing changes of depositional environments in a reconstructed oceanic plate stratigraphy formed as a result of Early Archean plate tectonics.

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and C3-sections were deposited in this order during the interval of volcanism. The uppermost C3-section is more enriched in Zr, Nb, Hf, and Th than the lower C1- and C2-sections (Fig. 7), suggesting the increased contribution from felsic rocks in this section. In order to provide some quantitative constraints on the source of detrital material in the studied sections, (La/Yb)N is plotted against Th/Sc in Fig. 16. The ratio of Th/Sc is an excellent indicator of the detrital source in sediments (e.g. Taylor and McLennan, 1985; McLennan and Hemming, 1992). Because felsic plutonic/volcanic
rocks have much higher (La/Yb)N and REE than mafic rocks, an increased contribution from felsic rocks should bring about the rise of (La/Yb)N . The mixing line between felsic (average for differentiated granitoids in the Shaw Batholith; avg SBG∗ ) and mafic (arithmetic mean of ABTH and ABKO) end-members is given in this figure. The C1-section cherts show some departure from this mixing line probably because their REEs were derived from other sources (e.g. hydrothermal solutions) apart from the detrital source. The inferred detrital REE contents gradually increase from the C1- to the C3-sections via C2-section, and the C2- and C3-section samples plot along this mixing line. The contribution from the differentiated felsic rock (avg SBG∗ ) is estimated to be up to ∼40% in the C3-section. The same degree of Cr enrichment in the C1-, C2-, and C3-sections (Figs. 7 and 12) suggests that chemical compositions of komatiitic basalts did not change greatly during deposition of these chert sections. Therefore, it is assured that the possible secular change of volcanic activity producing the komatiitic basalts is not responsible for the increase of Th/Sc and (La/Yb)N . The depositional site of the C1-, C2-, and C3-sections is inferred to have changed and become closer to the source area where granitoids and/or their volcanic equivalents were exposed (erupted). Alternative explanations (e.g. sea-level change, hinterland uplift, onset or cessation of felsic volcanism) cannot be ruled out. However, similar geochemical changes corresponding with depositional positions have been reported for the Cenozoic cherts on the modern ocean floor, sampled by the Deep Sea Drilling Project and Ocean Drilling Program (Murray et al., 1992), and terrestrially exposed cherts in the Mesozoic and Cenozoic accretionary complex (Murray et al., 1991).

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By analogy, it seems valid to attribute the similar geochemical variations in these Archean cherts to a change of their depositional positions. The clastic rocks of the T-section generally plot on the mixing line, and the felsic rock-mixing ratio is estimated to vary from ∼40 to >90% in the T-section (Fig. 16). However, several mudstone samples show some departure from this mixing line and have lower (La/Yb)N , along with high Th/Sc. Moreover, these mudstones
have significant negative Eu anomalies and high REE contents. These geochemical constraints suggest that much more differentiated and evolved plutonic/volcanic rocks than dacite were a source for these mudstones. This is consistent with the very low Cr contents in the T-section (Fig. 12). On the other hand, the PFVS exhibit an increase in (La/Yb)N (Fig. 16). Therefore, sources for the clastic rocks in the Panorama Formation may have been fairly varied in the eastern Pilbara Craton. In conclusion, the uppermost mudstones and sandstones of the T-section were deposited as a turbidite flow near a continental margin, and the provenance of the T-section was not an undifferentiated island arc, but a differentiated and evolved continent. The great variation of the rock record as demonstrated by the present study cannot be explained by tabular deposition in a closed basin. These sedimentary rocks provide a record of varied depositional environments ranging from a mid-oceanic spreading center to a convergent plate boundary via hot-spot oceanic plateau/islands (Fig. 17). This continuous change of depositional environments was most likely caused by horizontal plate motions, supporting the operation of plate tectonics in the Early Archean.

7. Conclusions The sedimentary rocks in the study area came from a wide variety of depositional environments varying from a MOR to a convergent plate boundary via a hot-spot. The source for the thickest ferruginous chert section (A) was mostly a high-T hydrothermal solution enriched in iron and silica at a MOR. The ferruginous cherts were deposited as in situ hydrothermal precipitates on Fe-rich, low-K tholeiites. The geochemical signatures of the cherts are analogous to those of modern MOR hydrothermal sediments in

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that P, V, Zn, and Y are positively correlated with Fe, ∗ ∗ and that Eu/Eu
decreases and Ce/Ce increases with increasing REE and LREE/HREE. The B-section cherts contain predominant SiO2 and significant amounts of Ba, and thus likely originated from a low-T MOR hydrothermal solution. The B-section overlying Fe-rich, low-K tholeiites is associated with massive silica veins that were originally hydrothermal feeders in normal fault zones in the spreading center. Intense and widespread hydrothermal activity at a MOR probably produced well-developed thick chert sections such as the A- and B-sections on oceanic crust. Moreover, geochemical constraints show that the associated greenstones are MORB in origin. Geochemical and geological lines of evidence consistently demonstrate that the tholeiitic basalt unit having thick hydrothermal chert sections is an allochthonous block of obducted oceanic crust formed at a MOR. Chert sections intercalated conformably with komatiitic basalts (C1-, C2-, and C3-sections in ascending order) are much thinner. The cherts were likely derived from a low-T Si-precipitating hydrothermal solution that may have been derived from hot-spot volcanism. Due to lack of large-scale hydrothermal circulation through normal fault zones, the hydrothermal activity responsible for these thin chert sections was relatively weak. Enrichment of Cr and Ni indicates a significant contribution of komatiitic detrital component during precipitation. The uppermost chert section (C3) also exhibits enrichment of Zr, Nb, Hf, and Th, and high Th/Sc and (La/Yb)N , suggesting that its de-

positional site gradually neared a continental source area. The clastic rocks of the uppermost T-section were deposited as turbidites near a continental margin. Differentiated and evolved granitoids and/or their volcanic equivalents had already been exposed (erupted) at least locally, and recycling of them had occurred in the Early Archean. The sedimentary rock record of a variety of depositional environments can be best explained by horizontal plate motions, supporting the operation of plate tectonics in the Early Archean.

Acknowledgements We are very much indebted to B. Windley for his careful and critical review. Discussions with H.D. Holland, H. Ohmoto, A.H. Hickman, and M. Van Kranendonk were very helpful and fruitful. We thank Y. Watanabe, N. Imai, T. Ishii, and H. Yoshida for analytical facilities, and A. Ikezaki, T. Kawakami, K. Fujinaga, and E. Moriguchi for assistance with many of the analyses. Field collaboration with S. Maruyama, M. Terabayashi, T. Komiya, Y. Isozaki, and G. Kimura was of great assistance. We would also like to acknowledge M. Barley and C.M. Fedo for their constructive and helpful comments. This work was supported by the Ministry of Culture and Education of Japan through Grant-in Aid Nos. 07238105, 06041038, 08041102, 06740415, and 07740426, and by funds from co-operative programs (No. 40 in 1994, No. 28 in 1995, and No. 32 in 1997) provided by Ocean Research Institute, the University of Tokyo.

Appendix A. Major, trace, and rare earth element data for the A-section

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Appendix B. Major, trace, and rare earth element data for the B-section

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Appendix D. Major, trace, and rare earth element data for the C2-section

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