Early Archaean crustal evolution: evidence from ~3.5 billion year old greenstone successions in the Pilgangoora Belt, Pilbara Craton, Australia

MICHAEL GODFREY GREEN

Submitted in fulfilment of the requirements for the degree of Doctor of Philosophy

School of Geosciences Division of Geology & Geophysics The University of Sydney March 2001 Acknowledgements

This project was conceived by Roger Buick who recognised the significance of the Pilgangoora successions in addressing crustal growth processes. I hope that this thesis goes some way towards answering the questions that occupied so many of our discussions. Thanks for your encouragement, advice and belief. Support for this project was generally funded through ARC Grants, although Sipa Resources assisted with the 1997 field season and also provided aerial photographs, maps and other assistance; thanks Mike, Pete, Walter and John. Whole- rock geochemistry was undertaken in collaboration with Paul Sylvester and isotopic studies with Sam Bowring, Mark Schmitz and Jeff Vervoort. Support to attend conferences was provided by Edgeworth David Travelling Scholarships and the Geochemical Society. I have been greatly assisted during my studies by many of the technical staff at The University of Sydney, especially Tom, Rob, Phil, Dave, George, Nancy, Maria and Erica. Thanks also to the students and academic staff who enriched my life at Sydney. A special thanks to all who shared the wonderful days in the Pilbara, particularly Mike, Owen, Ian, Kevin and even Jochen. Thanks also to Dave, Richard and Peter from AGSO. To Natalie for putting up with all my absences, especially during the last 6 months when my mind was elsewhere, no amount of thanks can repay. Thanks to Mum and Dad and the rest of my family (including the Maloneys) for always believing and for always being there. CONTENTS

CHAPTER 1: INTRODUCTION 1.1 PREAMBLE 1 1.2 REGIONAL GEOLOGICAL SETTING 4 1.3 FIELD AREA 10 1.4 AIMS AND METHODS OF RESEARCH 10 1.5 TERMINOLOGY AND CONVENTIONS 12 1.6 THESIS ORGANISATION 12

CHAPTER 2: GEOLOGY 2.1 INTRODUCTION 13 2.2 PREVIOUS WORK 14 2.3 GEOLOGICAL INVESTIGATION 17 2.3.1 Geological overview 18 2.4 COONTERUNAH GROUP 21 2.4.1 Mafic rocks 21 2.4.2 Felsic rocks 23 2.4.3 Sedimentary units 24 2.4.4 Interpretation 25 2.5 CARLINDI GRANITOID COMPLEX 25 2.5.1 Granodiorite 26 2.5.2 Monzogranite 26 2.5.3 Inclusions 27 2.5.4 Dykes 27 2.5.5 Interpretation 28 2.6 COONTERUNAH-CARLINDI / WARRAWOONA UNCONFORM. 28 2.7 WARRAWOONA GROUP 29 2.7.1 Strelley Pool Chert (SPC) 29 2.7.1.1 Quartz-rich sandstone 29 2.7.1.2 Laminated carbonates 30 2.7.1.3 Black-white plane-laminated chert 31 2.7.1.4 Volcanic sediments 31 2.7.2 Basalts and gabbros 32 2.7.3 Mafic volcaniclastic units 33 2.7.4 Interpretation 33 2.8 POST-DEPOSITIONAL PROCESSES 36 2.8.1 Deformation 36 2.8.2 Metamorphism 37 2.9 SYNTHESIS 39 2.10 DISCUSSION 40 2.10.1 Pilgangoora Belt stratigraphy 40 2.10.2 Pilbara correlations 41 2.11 SUMMARY 42

CHAPTER 3: GEOCHRONOLOGY 3.1 INTRODUCTION 43 3.2 PREVIOUS WORK 44 3.3 ANALYTICAL METHODS 46 3.3.1 Sample preparation 46 3.3.2 Analytical conditions 47 3.3.3 Data processing 49 3.4 COONTERUNAH GROUP 54 3.4.1 Results 54 3.4.1.1 Central dacite, Sample 70649 54 3.4.1.2 Eastern rhyolite, Sample 70601 55 3.4.1.3 Lower rhyolite, Sample 70660 56 3.4.1.4 Granitic xenoliths, Sample 520798 57 3.4.2 Interpretation of Coonterunah geochronology 61 3.5 CARLINDI GRANITOIDS 62 3.5.1 Results 62 3.5.1.1 Unconformity microgranite, Sample 94058 63 3.5.1.2 Unconformity granodiorite, Sample 95037 64 3.5.1.3 Wilson Well granodiorite, Sample 153188 65 3.5.1.4 Shilliman Well granodiorite, Sample 153190 66 3.5.1.5 Xenolith host granite, Sample 100698 69 3.5.1.6 Gneissic tonalite xenolith, Sample 080698 70 3.5.1.7 Gneissic granite xenolith, Sample 090698 72 3.5.2 Interpretation of the Carlindi granitoids 73 3.6 WARRAWOONA GROUP 74 3.6.1 Results 74 3.6.1.1 Strelley Pool Chert Sandstone, Sample 98OB5002 74 3.6.1.2 North Pole Dome (#9 chert), Sample 94001 77 3.6.2 Interpretation 79 3.7 DISCUSSION 80 3.7.1 Geological history 81 3.7.2 Regional stratigraphy 85 3.7.3 Pre-Coonterunah zircons 86 3.8 SUMMARY 88

CHAPTER 4: ELEMENTAL GEOCHEMISTRY - MAFIC ROCKS 4.1 INTRODUCTION 89 4.2 PREVIOUS WORK 89 4.3 ANALYTICAL METHODS 90 4.3.1 ICP-MS versus pressed-powder XRF 91 4.3.2 Interpretation 95 4.4 MAJOR ELEMENTS 97 4.5 TRACE ELEMENTS 102 4.6 NORMALISED DIAGRAMS 108 4.7 INTERPRETATION 111 4.7.1 Relationships within and between basalt suites 112 4.7.2 Crustal component 113 4.7.3 REE models 115 4.7.4 Eccentric basalts 118 4.7.5 Contamination estimate 119 4.8 DISCUSSION 121 4.8.1 Previous Pilgangoora survey 122 4.8.2 Other Pilbara basalts 122 4.9 SUMMARY 125 CHAPTER 5: ELEMENTAL GEOCHEMISTRY - FELSIC ROCKS 5.1 INTRODUCTION 126 5.2 WHOLE-ROCK GEOCHEMISTRY 126 5.2.1 Major elements 127 5.2.2 Trace elements 130 5.2.3 Normalised diagrams 134 5.3 INTERPRETATION 138 5.3.1 Petrogenesis: Coonterunah felsic volcanics 138 5.3.2 Petrogenesis: Carlindi granitoids 143 5.3.2.1 Trondhjemite 143 5.3.2.2 Granites 152 5.3.2.3 Anomalous samples 154 5.4 DISCUSSION 155 5.4.1 Pilbara correlations 155 5.4.2 Basalt contaminant 159 5.5 SUMMARY 160

CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 6.1 INTRODUCTION 161 6.2 PREVIOUS WORK 162 6.3 ANALYTICAL METHODS 163 6.4 RESULTS 164 6.4.1 Sm-Nd system 164 6.4.2 Lu-Hf system 167 6.5 INTERPRETATION 168 6.5.1 Sm-Nd system 170 6.5.2 Lu-Hf system 172 6.5.3 Basalt components 172 6.5.4 Origins of the components 176 6.5.4.1 Mantle source 176 6.5.4.2 Crustal component 178 6.5.4.3 Combined Nd-Hf system 179 6.5.5 Origin of the granitoids 181 6.6 DISCUSSION 183 6.6.1 Pilbara evolution 183 6.6.2 Geological history 186 6.7 SUMMARY 188

CHAPTER 7: DISCUSSION - GEOLOGICAL SYNTHESIS 7.1 INTRODUCTION 189 7.2 CONTINENTAL BASEMENT 189 7.3 PRESENCE OF EARLY PILBARA CRUST 190 7.4 COMPOSITION OF EARLY PILBARA CRUST 192 7.5 FATE OF EARLY PILBARA CRUST 193 7.6 SUMMARY 194

CHAPTER 8: DISCUSSION - TECTONIC SETTING 8.1 INTRODUCTION 195 8.2 DIRECT OBSERVATIONS 196 8.3 DERIVED CHARACTERISTICS 200 8.4 POSSIBLE TECTONIC SETTINGS 206 8.4.1 Meteorite impact basins 207 8.4.2 Oceanic crust 208 8.4.3 Oceanic plateaux 208 8.4.4 Volcanic arcs 208 8.4.5 Back-arc basins 209 8.4.6 Rifts 210 8.4.7 Flood basalts 212 8.5 DISCUSSION 215 8.5.1 Mantle upwelling 215 8.6 SUMMARY 216

CHAPTER 9: DISCUSSION - MANTLE AND CRUSTAL EVOLUTION 9.1 INTRODUCTION 217 9.2 CONTINENTAL CRUST 217 9.2.1 Crustal genesis 218 9.2.2 Crustal growth 220 9.3 IMPORTANCE OF THE PILBARA 223 9.4 CRUSTAL GROWTH 224 9.4.1 Nd-Hf isotopic constraints 224 9.4.1.1 Mantle isotope evolution 225 9.4.2 Nb/U systematics 232 9.4.2.1 Mantle Nb/U evolution 234 9.4.3 Synthesis 235 9.5 CRUSTAL GENESIS 236 9.5.1 Nd-Hf isotopic correlation 236 9.5.1.1 Hf-paradox - complementary reservoir 237 9.5.1.2 Hf-paradox - reappraisal of BSE 239 9.5.1.3 Early Archaean Nd-Hf decoupling 241 9.6 SUMMARY 242

CHAPTER 10: CONCLUSIONS 243

REFERENCES 245

APPENDIX 1: SHRIMP U-Pb GEOCHRONOLOGY APPENDIX 2: WHOLE-ROCK ELEMENTAL ABUNDANCES APPENDIX 3: Nd-Hf ISOTOPES ABSTRACT

In the Pilgangoora Belt of the Pilbara Craton, Australia, the ~3517 Ma Coonterunah Group and ~3484-3468 Ma Carlindi granitoids underlie the £3458 Ma Warrawoona Group beneath an erosional unconformity, thus providing evidence for ancient emergent continental crust. The basalts either side of the unconformity are remarkably similar, with N-MORB-normalised enrichment factors for LILE, Th, U and LREE greater than those for Ta, Nb, P, Zr, Ti, Y and M-HREE, and initial e(Nd, Hf) compositions which systematically vary with Sm/Nd, Nb/U and Nb/La ratios. Geological and geochemical evidence shows that the Warrawoona Group was erupted onto continental basement, and that these basalts assimilated small amounts of Carlindi granitoid. As the Coonterunah basalts have similar compositions, they probably formed likewise, although they were deposited >60 myr before. Indeed, such a model may be applicable to most other early Pilbara greenstone successions, and so an older continental basement was probably critical for early Pilbara evolution. The geochemical, geological and geophysical characteristics of the Pilbara greenstone successions can be best explained as flood basalt successions deposited onto thin, submerged continental basement. This magmatism was induced by thermal upwelling in the mantle, although the basalts themselves do not have compositions which reflect derivation from an anomalously hot mantle. The Carlindi granitoids probably formed by fusion of young garnet-hornblende-rich sialic crust induced by basaltic volcanism. Early Archaean rocks have Nd-Hf isotope compositions which indicate that the young mantle had differentiated into distinct isotopic domains before 4.0 Ga. Such ancient depletion was associated with an increase of mantle Nb/U ratios to modern values, and hence this event probably reflects the extraction of an amount of continental crust equivalent to its modern mass from the primitive mantle before 3.5 Ga. Thus, a steady-state model of crustal growth is favoured whereby post ~4.0 Ga continental additions have been balanced by recycling back into the mantle, with no net global flux of continental crust at modern subduction zones. It is also proposed that the decoupling of initial e(Nd) and e(Hf) from its typical covariant behaviour was related to the formation of continental crust, perhaps by widespread formation of TTG magmas. CHAPTER 1: INTRODUCTION 1

Chapter 1: INTRODUCTION

1.1 PREAMBLE The solar nebula collapsed over 4.56 billi on years ago, and the planets, including the Earth, grew by accreting the residue (Chen & Wasserburg, 1981; Tera & Carlson, 1999). After 100 million years, the Earth had accumulated most of its present mass, had segregated its FeNi-rich core and the Moon had formed (Chen & Wasserburg, 1990; Righter & Drake, 1996; Lee & Halliday, 1995, 1996; Lee et al., 1997). By this time, a cool, solid crust would have also formed at the mantle’s interface with the hydrosphere and atmosphere. The modern Earth has two dist inct types of crust: oceanic and continental. Oceanic crust is composed of mafic and ultramafic volcanic and intrusive rocks derived directly from the mantle and is typically recycled back into the mantle within 200 million years of formation. Hence, oceanic crust is not readily preserved and so there is little direct evidence of ancient oceanic processes. In contrast, continental crust comprises a broad range of rocks, which on average are enriched in incompatible elements (those elements which preferentially enter magmas) relative to the mantle (Rudnick & Fountain, 1995; Rudnick, 1995; Hofmann, 1997). Continental crust has a low density, and so it is more buoyant and does not recycle into the mantle as easily. Therefore, continental crust has been preferentially preserved compared to oceanic crust, although the volume of preserved ancient continental crust is very small. For instance, the oldest- known terrestrial materials are microscopic <4.40 billion year old zircon crystals which are preserved in younger metamorphosed sandstones (Froude et al., 1983; Compston & Pidgeon, 1986; Wilde et al., 2001), and the oldest-known terrestrial rocks are the ~4.03 billion year old Acasta gneisses which are exposed over less than ~40 km 2 (Bowring & Williams, 1999). Moreover, although continental crust covers about 40 % of the terrestrial surface, rocks older than 2.5 Ga account for less than 15 % of the exposed continental area (Windley, 1995). The paucity of ancient continental rocks can be explained by two general models: 1) the total mass of ancient crust was small and has gradually increased to attain the modern mass, and 2) the total mass of ancient crust was large and the majority of it has been subsequently destroyed. In other words, preservation of ancient crust has been good or bad, respectively. The former are gradual growth models, whereby the CHAPTER 1: INTRODUCTION 2 preserved rock record more or less represents the formation of continental crust (Hurley & Rand, 1969; Reymer & Schubert, 1984; Taylor & McLennan, 1985, 1995; McCulloch & Bennett, 1993, 1994; Albarède, 1998; Collerson & Kamber, 1999), and the latter are steady-state models, whereby the preserved rock record is unrepresentative of actual crustal growth rates (Armstrong, 1968, 1981, 1991; Bowring & Housh, 1995; Sylvester et al., 1997). Moreover, gradual growth models imply that the continental mass is presently increasing, whereas steady-state models imply that the mass is staying constant. Thus, determining the modern flux of continental crust may discriminate between the two models. Modern continental crust is formed at subduction zones by volatile-induced partial melting of the mantle wedge above subducting oceanic lithosphere (Taylor, 1967, 1977; Ellam & Hawkesworth, 1988; Rudnick, 1995). Thus, the inevitable recycling of oceanic crust appears to create continental crust. Subducting slabs, however, are complex. For example, interaction with seawater causes chemical changes to the basaltic crust, and overlying clastic and organic sediments may also be subducted. Thus, some continental material is also recycled into the mantle during subduction. By calculating the difference between volcanic output and continental input, the net flux can be determined. This calculation, however, has proven difficult for individual subduction zones (McCulloch, 1989; Plank & Langmuir, 1993; Pearce et al., 1995), and with so many modern subduction zones, calculating the global flux is an almost insurmountable problem. Thus far, studies of the modern Earth have been unable to resolve whether the continental crustal mass is growing. Therefore, to discriminate between the crustal growth models may require studying the terrestrial record when the models predict the greatest dissimilarities in continental mass, that is, in the Archaean. In general, the composition of continental crust is complementary to depleted mantle, and so it is widely held that continental crust has been extracted from the mantle (Rudnick, 1995; Hofmann, 1997). As a result, crustal growth could be monitored by studying the preserved ancient crust or, somehow, the ancient mantle. Preserved Archaean crust is composed of gneissic and granite-greenstone terrains. Gneissic terrains have been extensively modified by high-grade metamorphism, melting and deformation, and so deciphering their original features and genesis is difficult. In contrast, many granite-greenstone belts have not experienced such intense modification and many original features have been preserved. Moreover, it is quite CHAPTER 1: INTRODUCTION 3 possible that gneissic terrains are the high-grade-metamorphic equivalents of granite- greenstone terrains. Therefore, an understanding of granite-greenstone genesis is critical for constraining Archaean processes. It should be noted, however, that the preserved Archaean terrains may not represent all modes of ancient tectonics. For example, irrefutable evidence of preserved Archaean oceanic crust has yet to be discovered, although it was probably extant throughout Earth history (Bickle et al., 1994). This point is important as the “absence of evidence is not evidence for absence” (Sagan, 1977). Granite-greenstone terrains are composed of granite ( sensu lato) plutons and greenstone belts. In general, granite plutons comprise over half of the exposed terrains and are internally complex as a result of multiple intrusive episodes. Greenstone belts are predominantly composed of metamorphosed mafic volcanic rocks, and so have abundant green minerals (amphibole, chlorite, epidote). They commonly contain other supracrustal rock-types such as komatiites, felsic volcanics and clastic and chemical sediments. In general, the greenstones form shallow arcuate synclines between the plutons. The relationship between the granites and greenstones is variable, as some greenstones have been faulted against granites, whereas elsewhere granites have intruded the greenstones. Although grouped together as a geological entity, there are marked differences between greenstone belts, and so many processes have been invoked to explain their formation. The current consensus recognises the need for multiple mechanisms (de Wit & Ashwal, 1997, and papers therein), although this does not resolve the broader issue of whether actualistic tectonics (essentially modern plate tectonics, but perhaps with different parameters related to higher Archaean heat flow) can adequately explain the evolution of all greenstone belts or whether the early Earth was influenced by non-actualistic tectonics. Greenstone belts, therefore, may provide critical information about a range of Archaean tectonic processes, and since they also contain abundant mantle-derived rocks, they can provide constraints on mantle evolution, and perhaps, crustal growth models. Although there is a wealth of geological, geochemical and geophysical information about Archaean greenstone belts, there is still much contention about the interpretation of many of these data. For instance, in the oldest-known greenstone belt (Isua, west Greenland, ~3.77 Ga, Nutman et al., 1996) it has been recognised that the basalts were sourced from previously depleted mantle (Hamilton et al., 1978, 1983; CHAPTER 1: INTRODUCTION 4

Blichert-Toft et al., 1999), but three widely different mechanisms have been proposed to explain this depletion: formation of continental crust (Bennett et al., 1993), extraction of a primordial mafic crust (Chase & Patchett, 1988), or crystallisation of a magma ocean (Albarède et al., 2000). Hence, resolving the mechanisms and rates of Archaean crustal growth are clearly not straightforward. The research reported here focuses on some extremely well-preserved ancient greenstone successions from the Pilbara Craton, Australia, which provide a critical reference point for evaluating older, more equivocal terrains. Geological and geochemical considerations suggest that these Pilbara successions were deposited on continental basement, probably as submerged flood basalt provinces. Moreover, their elemental and isotopic compositions provide clear evidence that these ancient greenstones were derived from a mantle as depleted in incompatible elements as the modern oceanic basalt source. It is concluded that the present mass of continental crust was extracted from the mantle during the Earth’s early history, thus favouring steady- state growth of continental crust.

1.2 REGIONAL GEOLOGICAL SETTING The Pilbara Craton, Western Australia, comprises a granite-greenstone terrain which is predominantly exposed in the northern third of the craton and an unconformably overlying volcanosedimentary succession (Hamersley Province) which is best preserved in the southern two thirds of the craton (Fig. 1.1). Granite-greenstone inliers within the Hamersley Province, particularly along the southernmost margin, provide evidence for a granite-greenstone basement to the Hamersley Province (Tyler, 1990). Seismic and gravity models are also consistent with such a basement (Fig. 1.2; Drummond, 1988). Outliers of Hamersley units in the north suggest that these units were once more extensive and probably covered the entire Pilbara Craton. The southern and southeastern margins of the Pilbara Craton are unconformably overlain by the Proterozoic Ashburton, Bresnahan and Bangemall Basins, the northeastern and northern margins by the Silurian-Devonian Canning Basin and the western margin by the Permian Carnarvon Basin (Fig. 1.1). The eastern margin is in contact with the Proterozoic Paterson Orogen. CHAPTER 1: INTRODUCTION 5

116° E 118° E 120° E 20° S

CANNING BASIN INDIAN OCEAN

GRANITE-GREENSTONE

CARNARVON BASIN

22° S HAMERSLEY BASIN PATERSON OROGEN Rat Hill Inlier Turkey, Rooney, Springo Inliers Rocklea Dome Cooninia Inlier Wyloo Dome Milli Milli Dome Billinnooka Inlier

ASHBURTON BASIN

BRESNAHAN Sylvania Inlier BASIN GASCOYNE COMPLEX 50 24° S BANGEMALL BASIN & YOUNGER ROCKS kilometers

Figure 1.1: Outcrop distribution of granite-greenstone basement and overlying Hamersley Basin in the Pilbara Craton. Adjacent terrains are also shown (adapted from Tyler, 1990).

The granite-greenstone terrain comprises large ovoid granitoid complexes which are surrounded by volcanosedimentary greenstone belts (Fig. 1.3). Seismic profiles show that the crust thickens to the south from 28 to 32 km (Fig. 1.2; Drummond, 1988). The general crustal structure and composition has been defined from seismic and gravity data which show that the greenstone belts are confined to the upper 5 km of the crust, the granitoid domes extend to nearly 15 km, and the bottom half of the crust has a density and seismic velocity similar to felsic granulite (Drummond, 1988). Other seismic structures within the crust have not been resolved. Unlike many other cratonic areas, such as southern Africa, no xenoliths from the lower crust or mantle have been found in the Pilbara, although alluvial diamonds have been discovered in the Nullagine area. Thus, lithological constraints on deep crustal structure and composition are not available in the Pilbara. The greenstone belts generally form tight, upright, shallowly plunging synclines such that the limbs of each belt are subvertical and parallel to the broad trend of the belt. The two main exceptions are the North Pole and McPhee Domes where dips are shallower at the domal centre and steepen outwards (Hickman, 1983). The greenstone successions have been metamorphosed from sub-greenschist to amphibolite facies, and in places, there is exceptional preservation of primary volcanic and sedimentary features. CHAPTER 1: INTRODUCTION 6

Elsewhere these primary features have been destroyed by localised deformation and accompanying alteration and metamorphism.

GRAVITY MODEL +400 s i m 5.86

p SEISMIC MODEL l e

B 6.0 - 6.2 o u g ) u

m 20 e 6.4 - 6.45 k r (

a

-400 h n calculated t o p m e

d 40 a l y 8.1 - 8.34

-800 ( m

measured m s - 1 100 ) 300 -1200 distance (km)

2.74 2.80 Ashburton 2.78 Bangemall Basin Basin Hamersley Province granite-greenstones + 20 + + + + + 2.86 + + + + + + + + + + 3.08 + + + + ++ + + + + ++ + + + + + ++ + + + + + + + + + + d + + + + + + + + + + + + + + + + + e + + + + + + + + + + + + + + + 40 p + + + + + + + + +

t + + + + + h

( k

m Felsic Granulite

60 ) 3.35 NOT TO SCALE

80 Garnet Granulite: Lower crustal root of dense metamorphic and/or volcanic rocks with the Capricorn Orogen

100 distance (km) 300

Figure 1.2: Interpreted density and seismic velocity models from the Pilbara Craton. Density in gcm -3, observed Bouguer (density = 2.67 gcm-3), Seismic velocity in kms -1. Cartoon shows interpreted regional geology based on the models. The vertical scale has been expanded in the cartoon; felsic granulite = 2.86 in gravity model and 6.4 - 6.45 in seismic model (from Drummond, 1988). All figures based on the same SSW-NNE cross-section.

The granitoids are extremely complex with a prolonged history of multiple intrusion and deformation. Generally, the earliest phases were trondhjemite and granodiorite, while later events were granite ( sensu stricto). Correlation between complexes is not straightforward, although widespread events between 3470 to 3400, 3325 to 3290, 3270 to 3235 and 2940 to 2920 Ma have been recorded (Nelson et al., 1999). The oldest-known phases are a ~3655 Ma gneissic xenolith in the Warrawagine (Nelson, 1999) and a ~3578 Ma gabbroic anorthosite ènclave in the Shaw Granitoid Complexes (McNaughton et al., 1988). These xenoliths suggest the importance of older crust in forming the younger granitoids. The relationship between the granitoids and greenstones is also not straightforward, although some felsic volcanism was broadly coincident with granitoid emplacement. Contacts are both intrusive and tectonic, and so the relationship between the granitoids and the greenstones has been interpreted to CHAPTER 1: INTRODUCTION 7 result from either horizontal tectonics (Bickle et al., 1980, 1985; Bettenay et al., 1981) or vertical diapiric emplacement (Collins, 1989; Collins & Van Kranendonk, 1999).

118° 120° + Indian Ocean + + Port Hedland + + + + + + + + + + + + + I + + + + + + + + + + + + + + + + + + + + + + + + + + + + + Karratha + + + + + + + + + + + + j + + + + + + + III + + + + + + + + + + + + + I + + + + + + + + + + + + + + + + + + + + + + + + IV + + + + + + + + + + + + + + + + + + + + + + + + + + + + a I + + + + + + + + i + + 21° + b + + + + + + + c Marble Bar + + + + + + + + + e + + + + + + + + + f +V + + + + + + + + + + + + + d + + + + + + + + + + + + + + + + + + + + + + II + g + + + + 0 50 100 + + h + + + + + + + + + + + + VI kilometers + + + m + + k + l + + + + + + + + + + + Fortescue Group + VII + n + + + + + + + + + + + + + + + + + + + Granitoid Complexes + + 22° + + + + I Carlindi + + II Yule Greenstone belts III Muccan a Mallina f North Pole k Coongan IV Warrawagine b Wodgina g North Shaw l Kelly V Mount Edgar c Pilgangoora h Soanesville m McPhee VI Corunna Downs d Pincunah i Marble Bar n Mosquito Creek VII Shaw e Lalla Rookh j Shay Gap Figure 1.3: Distribution of granites and greenstones in the northern Pilbara Craton. Main geological entities discussed in this thesis are labelled. Field area in diamond, see Figure 2.1.

North- to northeast-trending crustal-scale strike-slip faults dissect the craton (Fig. 1.3) and have been used to define tectonostratigraphic domains based on the sequence stratigraphy of <3325 Ma volcanosedimentary successions (Krapez & Barley, 1987; Krapez, 1993; Krapez & Eisenlohr, 1998). Recent geochronological studies, however, have shown that these strike-slip faults probably did not control the deposition of the older units (Buick et al., 1995; Barley et al., 1998). That aside, rocks older than 3300 Ma are restricted to the east Pilbara (Fig. 1.3). Supracrustal rocks in the greenstone belts have been assigned to the Pilbara Supergroup (Hickman, 1983). In the east Pilbara, this was initially subdivided into the volcanic-dominated Warrawoona Group and the overlying sediment-dominated Gorge Creek Group (Hickman, 1981, 1983). However, the recognition of more unconformities and the advent of precise geochronology have required many new stratigraphic divisions. For example, a new basal unit to the entire succession has been discovered in CHAPTER 1: INTRODUCTION 8 the Pilgangoora Belt (Buick et al., 1995). The general lithostratigraphy of the Pilbara Supergroup in the east Pilbara is shown in Figure 1.4 and summarised below.

DeGrey Includes Lalla Rookh Sandstone, Mosquito Creek Formation: Epiclastic rocks, carbonates, minor felsic volcanics. Local stratigraphic subdivisions are common, poor age constraints.

Gorge widespread dominantly fine-grained sedimentary succession, sandstone, siltstone, shale, conglomerate, BIF, Creek minor volcanics. Local stratigraphic subdivisions are common due to lack of age constraints. Kangaroo Caves Formation: dacite to rhyolite volcaniclastics and volcanics, minor chert ~3240 Ma Sulphur Kunagunarrina Formation: magnesian to komatiitic basalt, komatiite, chert, shale

Springs sandstone, shale, quartzite, some volcanics

P Leilira Formation:

U Six Mile Creek: tholeiitic to magnesian basalt, generally massive O

R Wyman Formation: porphyritic rhyolite, only known in Kelly Belt, ?correlative is Golden Cockatoo ~3325 Ma G

p magnesian to tholeiitic basalt & gabbro, pillows common, silicified mafic volcanics R Euro Basalt: h u E s o a

r silicified quartz-rich sandstone, carbonates & mafic volcanics

P Strelley Pool Chert: a g g l U n b felsic volcanics and volcaniclastics a Panorama Formation: S ~3458 Ma o u

S o S A Apex Basalt: magnesian to tholeiitic basalt & gabbro, pillows common, silicified mafic volcanics w R a

r dacite to rhyolite volcanic and volcaniclastics A Duffer Formation:

r ~3466 Ma B a a L p g

I magnesian basalt, pillows common, evaporites and clastic sediments

l Dresser Formation: W u P a o T r

g Mt Ada Basalt: tholeiitic to magnesian basalt & gabbro, typically massive, some pillows a b g l u a S McPhee Formation: tholeiitic to magnesian basalt & gabbro, some ultramafic units T h

a Double Bar Formation: tholeiitic basalt to gabbro, typically massive with rare pillows n u r Coucal Formation: massive tholeiitic basalt & gabbro, dacite to rhyolite volcanics, laminated chert ~3515 Ma e t n

o Table Top Formation: tholeiitic to magnesian basalt & gabbro, typically massive with rare pillows, o rare pyroxene-phyric basalt C

Figure 1.4: General lithostratigraphic column for Pilbara Supergroup in the eastern Pilbara. Ages are referred to in the text.

The ~3515 Ma Coonterunah Group forms the base of the Pilbara Supergroup and is composed of tholeiitic basalt and gabbro with minor felsic volcanic rocks, magnesian basalt, chert and clastic sedimentary units (Buick et al., 1995). The Coonterunah Group has only been found in the Pilgangoora Belt where it is at least 6 km thick, has a strike-length of ~50 km and is unconformably overlain by the Strelley Pool Chert and Euro Basalt of the Warrawoona Group. In most of the east Pilbara greenstone belts, the Warrawoona Group forms the base of the supracrustal succession. It is dominantly composed of tholeiitic and magnesian basalt and gabbro with significant felsic volcanic units and minor ultramafic and sedimentary units. In any single belt the exposed Warrawoona succession is less than 10 km thick, but the overall succession may be much thicker depending on the validity of correlations between belts. In the Marble Bar Belt, for example, the Duffer Formation is ~5 km thick, yet it is absent from the neighbouring North Pole Dome. CHAPTER 1: INTRODUCTION 9

Indeed, the absence of such important marker units in some greenstone belts has impeded lithostratigraphic correlations, and so the Warrawoona Group may be a mixture of broadly coeval, but distinct, greenstone successions. This is highlighted by the Duffer and Panorama Formations which have been dated at ~3465 and ~3457 Ma, respectively (Fig. 1.4; Thorpe et al., 1992a; McNaughton et al., 1993). Additional problems have been raised in the McPhee Dome and Kelly Belt where the Warrawoona Group has been dated at 3430 ± 3 and 3417 ± 9 Ma, respectively (Barley et al., 1998). Until detailed age constraints are available for all the greenstone belts, lithostratigraphic correlations should be cautious. The Sulphur Springs Group predominantly consists of magnesia n to tholeiitic basalt, dacite, andesite, rhyolitic volcanic and volcaniclastic units, epiclastic sediments and minor cherts. It is confined to the supracrustal belts immediately adjacent to the Strelley Granite, particularly in the Soanesville and Pincunah Belts. An eruption age of ~3235 Ma has been determined for the Kangaroo Caves Formation, which is coeval with the Strelley Granite (Buick et al., submitted). Two packages of rocks are known to have ages older than the Sulphur Springs Group and younger than the Warrawoona Group; the ~3325 Ma Wyman Formation in the southwestern Kelly Belt (Thorpe et al., 1992a; McNaughton et al., 1993) and the Golden Cockatoo Formation in the northeastern Yule Granitoid Complex, adjacent to the Pincunah Belt (Van Kranendonk, 2000). Presently, little is known about these rocks and their importance to the evolution of the Pilbara Craton. The Gorge Creek Group is predominantly composed of fine- to medium-grained sediments with some pillowed magnesian and tholeiitic basalt. Although the Gorge Creek Group is widespread throughout the Pilbara, its age, and hence regional correlations, are poorly constrained. For instance, the Cleaverville Formation in the west Pilbara may correlate with the Nimingarra Iron Formation in the east Pilbara, though this has not been validated by geochronology. Locally, the DeGrey Group unconformably overlies the Gorge Creek Group and is composed of fine- to coarse-grained epiclastic sediments with little volcanic input, except in the west Pilbara. The DeGrey Group includes the Lalla Rookh Sandstone, a thick package of fluvial and lacustrine sediments between the Pilgangoora Belt and North Pole Dome. The DeGrey Group also has a poorly constrained age in most greenstone belts. Given that the Gorge Creek and CHAPTER 1: INTRODUCTION 10

DeGrey Groups may have been deposited in small strike-slip basins, Pilbara-wide correlations may be unrealistic (Krapez & Barley, 1987; Krapez, 1993). The Mount Bruce Supergroup unconformably overlies the Pilbara Supergroup (granite-greenstones), and thus is somewhat younger. It has been divided into the Fortescue, Hamersley and Turee Creek Groups. Only the Fortescue Group is present in the northern part of the Pilbara Craton and is predominantly composed of subaerially erupted tholeiitic flood basalts and epiclastic sediments (Blake, 1984, 1993; Nelson et al., 1992). Deposition of the Mount Bruce Supergroup began ~2767 myr ago (Arndt et al., 1991).

1.3 FIELD AREA The study reported herein is confined to the Pilgangoora Belt, where the Coonterunah and Warrawoona Groups form a ~50 km long sinusoidal belt, and the adjacent Carlindi Granitoid Complex, where the granitoid was generally mapped within 10 km of the granite-greenstone contact (Fig. 1.4). Over 6 km of the Coonterunah and up to 4 km of the Warrawoona successions are exposed, and the unconformable contact between them is marked by a prominent chert ridge. The Coonterunah Group is exposed to the north of this ridge as low, undulating hills, whereas the Warrawoona Group forms steep ridges and moderately flat valleys. The Carlindi granitoid is very flat and exposures are generally restricted to within 2 km of the greenstone contact, near creeks or as isolated domes. In general, exposure of the greenstones is fairly good, whereas that of the granitoids is poor. Weathering has variably affected the rocks and typically the rocks have thin oxidised mantles. The most spectacular exposures are pavements within creek beds where oxidised mantles have not developed. In the absence of outcrop, particularly above the granitoid, float could not be confidently mapped due to the prevalence of sheet-wash colluvium.

1.4 AIMS AND METHODS OF RESEARCH The aims of the research reported herein are: 1) to determine the petrogenesis of the Coonterunah-Warrawoona rocks in the Pilgangoora Belt, 2) to constrain the tectonic setting in which they formed, 3) to determine the composition of the mantle from which they were derived, 4) to provide new evidence regarding the evolution of the Pilbara CHAPTER 1: INTRODUCTION 11

Craton, 5) to constrain the evolution of the early Archaean mantle and 6) to determine the mechanisms and rates of Archaean crustal formation. The project was primarily field-based, with geo chemical analyses following from rigorous geological control. Mapping was based on 1:25,000-scale colour aerial photographs of the Pilgangoora Belt flown for Sipa Resources in 1992 and 1:50,000- scale black and white aerial photographs of the entire region flown for the Western Australian Department of Land Administration in 1993. Geological maps were drafted using MapInfo at the The University of Sydney, with rock divisions and codes similar to those used by Sipa Resources. Drill-core of the mapped units does not exist, so surface samples were used for whole-rock elemental and isotope analyses. These were cleaned to remove oxidised mantles and vein material. Samples are labelled xxmmyy, where mm indicates the month and yy the year of sample collection. Crushing was performed at the Research School of Earth Sciences, Canberra in 1997 and at the Australian Geological Survey Organisation, Canberra in 1998. All equipment was precontaminated by sacrificing a portion of each sample, thus reducing the affects of cross-contamination between samples. Major and minor element oxides were analysed by XRF techniques on fusion disks at the Australian National University, Canberra, in 1997, and Memorial University, Newfoundland, Canada in 1998. Minor and trace elements were determined by solution nebulisation ICP-MS techniques (Eggins et al., 1997), and were undertaken at the Research School of Earth Sciences, Canberra in 1997 and Memorial University, Newfoundland, Canada in 1998. To check Zr concentrations, all samples were analysed using XRF pressed-powder techniques at Memorial University (see Chapter 4). Whole- rock Sm-Nd and Lu-Hf isotopic analyses were undertaken at the Department of Earth, Atmospheric and Planetary Science, Massachusetts Institute of Technology and Ecole Normale Supérieure de Lyon, France, respectively. Ages were obtained using zircon U- Pb methods on the SHRIMP II ion microprobe in Perth, Western Australia. Analytical results are presented in the appendices. Rock descriptions are based on field observations and thin-section petrography. Semi-qualitative mineral compositions were obtained using a Philips 505 scanning electron microscope at the Electron Microscope Unit, The University of Sydney. CHAPTER 1: INTRODUCTION 12

1.5 TERMINOLOGY AND CONVENTIONS Rock names are based on field criteria, and are so used throughout the thesis. Although in some circumstances purely chemical definitions may be more appropriate, field definitions are generally adequate and provide continuity through all phases of the study. Most rock units have been metamorphosed, but are generally referred to by their precursor name. For instance, metamorphosed basalts are referred to as basalts. Where a precursor cannot be conclusively identified, typical metamorphic terms are used. Standard chemical and isotopic conventions are used, and comparative analyses for normalisation are referred to in the text. Specific definitions of important terms are according to Bates and Jackson (1993). Structures are defined using the dip and dip azimuth convention.

1.6 THESIS ORGANISATION In Chapter 2, the geology of the northern Pilgangoora Belt and southern Carlindi Batholith is described and interpreted in order to determine the geological conditions during eruption, deposition, crystallisation and subsequent metamorphism and deformation. In Chapter 3, the eruption and crystallisation ages of various units are determined by using SHRIMP U-Pb analyses of zircons. The whole-rock elemental abundances of mafic units are presented in Chapter 4, whereas those from felsic volcanics and granitoids are presented in Chapter 5. Nd-Hf isotopic compositions are presented in Chapter 6, and a synthesis of these data incorporating previous studies is presented in Chapter 7. The tectonic setting for the Coonterunah and Warrawoona Groups is considered in Chapter 8. In Chapter 9, the formation and evolution of continental crust is discussed. The conclusions are presented in Chapter 10. Geochemical data are presented in the appendices. Maps showing sample localities can be found as MapInfo worksheets on the included CD. CHAPTER 2: GEOLOGY 13

Chapter 2: GEOLOGY

2.1 INTRODUCTION The field area investigated in this study forms an irregular ~30 x 40 km polygon in the central part of the Pilbara Craton. The northwest corner of this polygon is ~80 km SSE of Port Hedland whereas the southeast corner is ~45 km WNW of Marble Bar (Fig. 1.4). The field area is mainly on the Marble Bar 1:250,000-scale geological map sheet (Hickman & Lipple, 1978) with the northernmost portion on the adjoining Port Hedland sheet (Hickman & Gibson, 1981). The area is also covered by finer-scaled topographic map sheets (Fig. 2.1). The recently published North Shaw 1:100,000-scale geological map sheet covers a substantial part of the field area (Van Kranendonk, 2000). There are two main geological regions in the field area, supracrustal greenstones (~25 %) and granitoids (~75 %). The greenstones comprise the northern part of the irregularly-shaped Pilgangoora Syncline, which is bound by the Carlindi and Yule Granitoid Complexes and the Lalla Rookh-Western Shaw Fault (Hickman, 1983; combined Pilgangoora-Pincunah Belts in Fig. 1.3). The southern part of the Pilgangoora Syncline, however, is geologically distinct from the northern part and so has been divided along the Mount York Deformation Zone (Fig. 2.1) into the Pilgangoora and Strelley Belts, respectively (Krapez, 1993). These have been renamed Pincunah and East Strelley Belts, respectively (Van Kranendonk, 2000). Hence, the field area comprises most of the Strelley/East Strelley Belt. This subdivision is useful, but the term East Strelley Belt is confusing because the belt is west of the prominent Strelley Granite, and so the name may be mistaken for that immediately east of the Strelley Granite (Soanesville Belt, but also named Strelley Belt in some studies: Vearncombe et al., 1995; Vearncombe & Kerrich, 1999). It is proposed that the term Strelley should be confined to the Strelley Granite and the northern part of the former Pilgangoora Syncline be named the Pilgangoora Belt to be consistent with most published references (Barley, 1993; Buick et al., 1995; Neumayr et al., 1993a, b, 1998; Green et al., 2000) and the southern part be named Pincunah Belt (Fig. 2.1). Furthermore, it has been proposed that the name Pilgangoora Syncline be retained for the prominent fold closure in the northwestern part of the Pilgangoora Belt (Van Kranendonk, 2000), but this feature is more complicated than a syncline and so the name adds little to regional nomenclatural clarity. In the field area, the granitoids comprise the southeastern part of CHAPTER 2: GEOLOGY 14 the Carlindi Granitoid Complex, the largest of the Pilbara granitoid complexes (Fig. 1.3; Hickman, 1983).

119° 00' 119° 15'

Shaw River Coonterunah Pool

Wallaringa Carlindi 21° 00' 21° 00' Wodgina North Shaw Lalla Rookh Mine

Strelley Pool LWSZ Mount York Lalla Rookh Synclinorium HF

MYDZ Pilgangoora Belt North Pole Pincunah Belt Strelley Granite Dome 119° 00' 119° 15'

10 kilometers

Basalt Ultramafic Shear zone MYDZ Mount York Deformation Zone

Chert Granodiorite Strelley Pool Chert LWSZ Lalla Rookh-Western Shaw Shear Zone

Sediment Cover Perimeter of mapping area HF Hogback Fault

Figure 2.1: Published geology of mapping area (adapted from 1:250,000- scale Marble Bar sheet (Hickman & Lipple, 1978) and Port Hedland- Bedout Island sheet (Hickman & Gibson, 1981)). The Mount York Deformation Zone separates the Pincunah and Pilgangoora Belts. The Strelley Pool Chert marks the Coonterunah-Carlindi unconformity. Cross represents the corner of the 1:100,000-scale map sheets.

2.2 PREVIOUS WORK Geological and geochemical studies in the Pilgangoora Belt have generally been associated with craton-wide mapping or investigating local mineralisation. A number of studies on the regionally important Strelley Pool Chert have included the Pilgangoora exposures. The discovery of a major unconformity in the belt has important implications for Archaean tectonic processes and crustal growth and the research here stems from this discovery. CHAPTER 2: GEOLOGY 15

The general geology of the Pilgangoora Belt is shown on the 1:250,000-scale geological map sheets (Hickman & Lipple, 1978; Hickman & Gibson, 1981), along with the localities of known mineralisation (Lalla Rookh, Mount York district) and some structural measurements (Fig. 2.1). The greenstones form a U-shaped belt wrapping around the southern margin of the Carlindi granitoids and are truncated to the east by the Lalla Rookh-Western Shaw Shear Zone. The western part of the greenstones are folded and truncated by another shear zone. The Strelley Pool Chert (SPC) is a prominent feature which clearly cuts another chert (Fig. 2.1), although the significance of this contact was not realised until recently. North and stratigraphically below the SPC, the basalts were described as schistose and were considered to be part of the Talga Talga Subgroup (Fig. 1.4; Hickman, 1981, 1983; Glikson et al., 1986b). South and above the SPC, pillowed basalts younging southwards were considered to be part of the Salgash Subgroup (Hickman, 1983; Glikson et al., 1986b). The 1:250,000-scale maps do not identify the Coonterunah felsic volcanics and show the exposure of Carlindi granitoids to be more extensive than they actually are (cf. Fig. 2.1 & 2.3, also Van Kranendonk, 2000). Regional geochemical studies by the Geological Survey of Western Australia (GSWA) and the Bureau of Mineral Resources (now Australian Geological Survey Organisation - AGSO) examined major, minor and trace element abundances of 56 whole-rock samples from a stratigraphic traverse ~4 km west of Strelley Pool in the Pilgangoora Belt (Glikson et al., 1986b). Of these, 37 were from the Coonterunah Group (then Talga Talga Subgroup) and 19 from the Warrawoona Group (Euro Basalt). Major element abundances show that the successions have variably altered magmatic compositions, but are dominantly metamorphosed tholeiitic and magnesian basalts. No profound compositional changes were noted between the two successions. These results are discussed in Chapter 4. Galena 207Pb/206Pb ratios from the Lalla Rookh and Mount York mines indicated mineralising events at ~3.20 (3.19) and ~2.76 (2.74) Ga, respectively (Richards et al., 1981; bracketed ages are recalculations by Thorpe et al., 1992b). This work highlights the confusion regarding the stratigraphic position of the Lalla Rookh deposit (below the SPC), referring to a gold-bearing quartz vein cutting basaltic schist of the Salgash Subgroup (Richards et al., 1981). A Salgash Subgroup host was also recognised by other researchers (Thorpe et al., 1992b). Combined whole-rock and mineral separate Pb CHAPTER 2: GEOLOGY 16 isotope ratios and SHRIMP titanite U-Pb ages showed that some mineral deposits in the Mount York district were probably emplaced at ~2.89 Ga (Neumayr et al., 1998). This study also concluded that the Pb was probably not sourced from the host rocks but required extended ingrowth from a low m (238U/204Pb) source, perhaps old high-grade metamorphic basement (Neumayr et al., 1998). The Strelley Pool Chert (SPC) forms a prominent silicified sedimentary unit in the Pilgangoora Belt (Fig. 2.1), and although most original minerals have been destroyed, many primary textures and features are well preserved. The SPC was original composed of clastic, volcaniclastic and evaporitic facies that were deposited in a shallow subaqueous to subaerial setting, and include some of the oldest-known putative biogenic structures (Lowe, 1980, 1983; Buick & Barnes, 1984; DiMarco & Lowe, 1989a, b). A sandstone unit at the base of the SPC has been studied from sedimentological (O’Brien, 1999) and geochemical (Rasmussen et al., 1998) viewpoints. Importantly, the SPC, or an almost identical chert unit, has also been identified in the Coongan, North Shaw and Kelly Belts and the North Pole Dome (Lowe, 1980, 1983; Buick & Barnes, 1984; DiMarco & Lowe, 1989a, b; Van Kranendonk, 2000). In 1992, most of the Pilgangoora Belt was mapped by Sipa Resources at 1:25,000-scale, discovering minor Pb-Zn-Cu-Au occurrences but more importantly identifying extensive felsic volcanic units and a major unconformity. Zircon U-Pb dating of the felsic volcanics beneath the unconformity confirmed the existence of a new stratigraphic succession, the ~3515 Ma Coonterunah Group (Buick et al. 1995). This resolved the ongoing conflict regarding the stratigraphic host of the Lalla Rookh deposit. Moreover, the discovery of the unconformity highlighted the existence of emergent continental crust in the early Archaean. The North Pilbara National Geoscience Mapping Accord (NGMA), a collaboration by the GSWA and AGSO, began in 1995 with systematic mapping of selected 1:100,000-scale map sheets. As part of this work, airborne magnetic and gamma-ray spectrometry surveys were flown over the Pilgangoora-Carlindi region along east-west traverses at 400 m line-spacings and 80 m ground clearance (Mackey & Richardson, 1997a, b). These data showed that the southeastern part of the Carlindi Granitoid Complex has very low radiometric and magnetic responses that are not due to recent colluvial cover. It also showed a large number of mafic dykes. All the greenstone belts had a uniformly subdued response with a few continuous highly magnetic horizons. CHAPTER 2: GEOLOGY 17

The Mount York Deformation Zone, which separates the Pincunah and Pilgangoora Belts, was clearly traced. In the Pilgangoora Belt, the geological information used for the 1:100,000-scale North Shaw sheet relied heavily on the Sipa 1:25,000-scale mapping with some additional ground truthing (Van Kranendonk, 2000). The production of this map lead to substantial stratigraphic revision, including subdivision and formalisation of the Coonterunah Group (Van Kranendonk & Morant, 1998). In addition, a number of unpublished and ongoing economic geology projects have been undertaken in the Mount York district by University of Western Australia and University of Newcastle students.

2.3 GEOLOGICAL INVESTIGATION In this project, the Pilgangoora Belt and adjoining Carlindi Granitoid Complex were remapped during 1997-1999 using 1:25,000-scale colour and 1:50,000-scale black and white aerial photographs. The fine-scaled photographs only covered the greenstones and the margins of the granitoid, whereas the other photographs covered the entire field area. The main focus of the project was to determine the compositional differences between the greenstone successions either side of the recently discovered unconformity. This geochemical data was integrated with the mapping to constrain the compositional trends within each succession. It was realised after the 1997 field season, however, that the Carlindi granitoids were integral to greenstone evolution, and so the field area was extended to include them. Geographical co-ordinates were determined using a single 12-channel hand- held GPS (Garmin 12XL) and were collected at a time when the US Defense Forces were adding a random error spike to satellite signals (before May 2000). Therefore, GPS localities have an accuracy of 50 m, though localities were only recorded when the error spike had decayed and the signal was stable (up to 5 minutes). Aerial photographs were not orthographically converted and so parallax errors have also added to map errors. Map accuracy was improved by drafting onto the available topographic sheets, but the accumulated errors are probably still in the order of 100 m, which equates to 4 mm on the 1:25,000-scale map sheets. The maps were drafted in MapInfo® Professional 5.5 and can be located on the Appendix CD (see included readme.txt).

2.3.1 Geological overview CHAPTER 2: GEOLOGY 18

The Pilgangoora Belt contains the only known outcrops of the Coonterunah Group, where they are unconformably overlain by the upper units of the regionally extensive Warrawoona Group (Strelley Pool Chert, Euro Basalt, Fig. 1.4, 2.2). The Warrawoona basalts are in turn unconformably overlain by clastic sediments of the Gorge Creek Group (Fig. 2.2), although there are some thin, highly altered horizons immediately on top of the Warrawoona basalts which may be felsic volcaniclastic units of the Sulphur Springs Group (Fig. 1.4). In the Pilgangoora Belt, the Coonterunah and Warrawoona successions now uniformly dip subvertically, providing a cross-section of the stratigraphy through which the unconformity is exposed for over 50 km strike length (Fig. 2.2). Younging is consistently away from the Carlindi Batholith and there is no evidence of stratigraphic repetition. The maximum exposed thickness of the Coonterunah Group is ~6.5 km, but the complete succession is probably somewhat thicker as the top and base have been truncated. The upper Warrawoona Group is up to 3.5 km thick. The Carlindi granitoids extend northwards for over 30 km from where they are cut by the unconformity. The greenstones in the field area can be divided into three domains (Fig. 2.3) according to their preservation and relationship with the unconformity. The eastern domain contains well-preserved low-grade basalts and felsic volcanic and volcaniclastic units, and the unconformable contact between the Coonterunah and Warrawoona successions is angular. In the central domain, the rocks are similarly preserved but the unconformable contact is subparallel. Although continuous Coonterunah outcrop between the eastern and central domains is lacking, zircon U-Pb dating and trace element abundances of felsic volcanics in each domain has confirmed their stratigraphic equivalence (see 2.9.1). In the western domain, the successions are apparently also semi- conformable, but the rocks have higher metamorphic grades with well-developed tectonic fabrics and structures. The work reported here has concentrated on the eastern and central domains due to their exceptional preservation. CHAPTER 2: GEOLOGY 19

G o r g e C r e e k G r o u p : angular unconformity, bedded, medium-grained sandstone + S T R E L L E Y P O O L C H E R T

shale, age poorly constrained.

silicified mafic arenite (as below), abundant tabular

v

v intraclasts, minor mafic ash.

v v

v v v

P M a g n e s i a n b a s a l t : pillowed + massive v silicified mafic arenite (as below), common tabular v

v v

U flows, abundant large ocelli, vesicles + v v

v intraclasts, minor mafic ash. v v v v

amygdales, relict plagioclase + pyroxene. v v v v

O v

v

v

v

v silicified medium-grained mafic arenite, rounded

v v R v

v v v v v v mafic sand + angular glass, gypsum pseudomorphs. G black + white planar-laminated chert. A G a b b r o : massive, subparallel to basalts,

N coarse-grained, relict clinopyroxene.

v v v v

v

v

O v

v v

v v v

v

O v

v

v v

C h e r t : as below + intraclasts of mafic lutite. v v v v

v v v W v v silicified undulose- to planar-laminated carbonates, v v

A v pseudomorphs of large aragonite fans. v v v v v

v

v v v v v v R T h o le i i t ic b a s a lt : massive + pillowed flows, v

R large abundant vesicles + amygdales, sub-

A ophitic, relict pyroxene + plagioclase. v v v v v v v

quartz sandstone: medium- to coarse-grained, v v v v v v v v v v v v v v v v v v v v v v

W C h e r t : grey, fine-grained mafic ash, massive to mature, rounded, massive + cross-bedded, heavy laminated, siliceous gypsum pseudomorphs. mineral laminations, conglomerate clasts to 10 cm of chert, sandstone, basalt + felsic igneous. v v v + + +

v + + + Unconformity: non-planar, ~10 m eroded topography, marked P v v + + + alteration beneath unconformity (possible palaeosol), sand-filled U v + [fissures penetrate to 100 m of Coonterunah greenstones. + + O v + + + R v + v + G v + +

+

v v + + H F e l s i c v o lc a n ic ro c k s : massive, highly x v v + A vesicular, dacitic to rhyolitic breccias + v v x + hyaloclastites, 3515 ± 3 Ma (Buick et v N x x + al., 1995). v v v +

U v v v v x + R v v v v x+ M i c r o g r a n it e : massive, quartz phenocrysts in E v x + medium-grained quartz, microcline, plagioclase T v v v + groundmass, 3468 ± 4 Ma (Buick et al., 1995). v v v +

N x v v v + O T h o le ii t ic b a s a lt : massive + pillowed flows, v v v v x x + O scarce small vesicles + amygdales, sub- v v x + ophitic, clinopyroxene + plagioclase v v C v x x + relicts, isolated felsic xenoliths. v v v v v x + v x x v v x + v x x v v v x x x v x x x v v v x x x v v v x x x G a b b r o : massive, intrusive sub-parallel v v v x G r a n o d io ri t e : massive, coarse-grained plagioclase, to basalts, medium- to coarse-grained, x x microcline, quartz, 3479 ± 11 Ma (herein). v v v v x x x clinopyroxene + plagioclase relicts. v v v v x x x v v v x x x v v v x v v v x x v x x x v v v v x v v x x v v x x x v v v v x x x v v x x x v v v v v x x x v v v C h e r t: 10 cm to 5 m wide, black + white v v v x x x planar-laminated,magnetite-rich. v x x v v v v v v x x v v v v x C h e r t: black, white, brown planar- v v v x laminated, magnetite, quartz + v v v v v x cummingtonite, carbonate lamellae. v x v v v x v v v v x v v x v v v x v v v v x v v v x v v v x C h e r t: 10 cm to 5 m wide, black + white v v v x planar-laminated,magnetite-rich. v v x v v x v v v x v x v v x v v v v x v v v x x v v v x x v v v x v v v x v v x x

v v

x

v v vv v x

v v x x

v v v

v v x

v v x v v

v v v x

v v v x x

v

v v v

v v v x

v v x x 1 km

v

v v v x x x

v

v v v

P y r o x e n e - p h y r ic b a s a l t : elongate v v v

v x x x

pyroxene in massive fine-grained vv

v v v

matrix, some relict pyroxene. v x x x v

v v v x

v v x x

x M a g n e s i a n b a s a l t : massive + rare pillowed v v v x flows, abundant ocelli up to 5 cm across, x x x x v x rare small vesicles + amygdales. x x x x x x x x x x x x x

Figure 2.2: Schematic stratigraphic column showing the general relationships between units in the Pilgangoora Belt. CHAPTER 2: GEOLOGY 20

119° 00' 119° 15'

Coonterunah Pool

21° 00' Lalla 21° 00' Rookh Mine Strelley Pool

western eastern

central

119° 00' 119° 15'

10 kilometers

Coonterunah chert Warrawoona chert Ultramafic Cover Coonterunah felsic volcanics Warrawoona basalt Granodiorite Laterite gravels Coonterunah basalt

Figure 2.3: Summary of geological map of the field area at the same scale as Fig. 2.1. Strelley Pool Chert is at contact between the two successions. The three structural domains are marked. Complete map in Appendix CD.

The southern Carlindi Granitoid Complex compr ises granodiorite and monzogranite with some ultramafic, mafic and felsic dykes (Fig. 2.1-3). The granitoids can be reclassified using their normative feldspar compositions into trondhjemites and granites (Chapter 5), but these terms are not used in the chapter as they cannot be readily applied in the field. Geophysical images show no major compositional variations across the complex, in contrast to the other Pilbara granitoid complexes. Importantly, a microgranite body that intrudes a granodiorite is cut by the unconformity and so some of the granitoids must predate Warrawoona deposition (Fig. 2.2). This has been confirmed by zircon U-Pb dating (Buick et al., 1995; Nelson, 1999; Van Kranendonk, 2000; see Chapter 3). In the field area, the westernmost granitoids have developed tectonic fabrics that are subparallel to those in the greenstones and show similar folding. There are isolated patches of mafic and gneissic xenoliths at various localities in the Carlindi granitoid. CHAPTER 2: GEOLOGY 21

Geological descriptions of the Pilgangoora successions and Carlindi granitoids are provided below. The aim is to demonstrate the main features of and differences between the various units and to constrain their depositional settings.

2.4 COONTERUNAH GROUP The Coonterunah Group is predominantly composed of basalt and gabbro with significant felsic volcanic and volcaniclastic beds and minor sedimentary units. Well preserved primary features are present in the eastern and central domains where metamorphic grades are relatively low and deformation minimal. Rare pillows and flow layering in the basalts and normal grading in felsic volcaniclastic debris flows indicate a uniform younging direction away from the Carlindi granitoid. The entire Coonterunah succession has been intruded by Carlindi granitoids such that the base of the succession is missing. A hornblende-rich metamorphic aureole is common in the greenstones within <250 m of the granite-greenstone contact, but is not ubiquitous.

2.4.1 Mafic rocks Most mafic rocks are fine-grained (<0.5 mm), massive, predominantly composed of amphibole, chlorite and plagioclase, and have retained volcanic features such as pillows, amygdales, varioles and sub-ophitic textures (Plate 2.1). Thus, they are metamorphosed basalts. Coarser-grained equivalents (up to 2 cm) are identified as gabbros, and since the metamorphic grade of the gabbros is the same as the basalts and both basalts and gabbros are truncated by Carlindi granitoids (Fig. 2.2-3), then the gabbros must have crystallised before granodiorite intrusion and are thus deemed to be part of the Coonterunah Group. The gabbros are typically tabular and semi-conformable, and may be intrusive sills or the base of thick basaltic flows. A few gabbros have clear intrusive relationships with the volcanic units. Most of the basalts and gabbros have appreciable amounts of leucoxene and thus are TiO2-rich and classified as tholeiites. Thin units of variolitic (magnesian) and pyroxene-phyric (?komatiitic) basalt are exposed near the base of the Coonterunah succession. The classification of basalts according to these criteria is also reflected in their elemental abundances (Chapter 4). Tholeiitic basalts and gabbros are typically massive and contain actinolite or low- Al hornblende (30-50 %), albite (30-50 %), chlorite (10-20 %), quartz (10-20 %), CHAPTER 2: GEOLOGY 22 epidote (<10 %) and margarite (<10 %) with minor leucoxene, magnetite, ilmenite and carbonate. Relics of igneous clinopyroxene, orthopyroxene, plagioclase, hornblende and Cr-bearing magnetite are preserved in places (Plate 2.1F-G), but such preservation is rare, especially in the basalts. Although metamorphic recrystallisation has been relatively complete, igneous textures have been preserved by pseudomorphic replacement. For instance, sub-ophitic felted textures of intergrown plagioclase and pyroxene have been preserved in many tholeiitic samples, where pale laths of albite-altered plagioclase have intergrown with fine-grained actinolite, chlorite and albite, the latter minerals probably replacing pyroxene and glass (Plate 2.1C). Magnesian basalts are restricted to a ~250 m thick package near the base of the Coonterunah succession and contain abundant varioles up to 5 cm in diameter of very fine-grained (<0.1 mm) albite and actinolite (~50 % each) in a fine-grained (<0.5 mm) actinolite-albite (~70:30 %) matrix (Plate 2.1D-E). Accessory minerals include chlorite, clinozoisite, carbonate and FeTi-oxide minerals. Relict igneous minerals have not been found. A single <20 m thick pyroxene-phyric basalt unit immediately overlies the uppermost magnesian basalt and consists of very fine-grained (<0.1 mm) Fe-Ti-oxides and low-Al hornblende after elongate pyroxene needles (up to 10 cm long) in a very fine-grained (<0.1 mm) actinolite-hornblende-albite groundmass (Plate 2.2A-B). Rare pyroxene relics are preserved. Such units are commonly referred to as komatiitic basalts, but in this particular case, this name may not be appropriate (see Chapter 4.4). It is difficult to identify individual basalt flows because the mafic units are generally massive and have relatively uniform metamorphic compositions. However, there is limited evidence that some of the flows were quite thin. For example, <20 cm thick layers defined by the upward increase of vesicle abundances have been recognised (Plate 2.1B). Preserved pillows are rare and are generally less than two metres in diameter, with well-defined chilled rinds, but few vesicles or amygdales (Plate 2.1A). Indeed, vesicles and amygdales are relatively uncommon in the Coonterunah basalts, especially compared to the overlying Warrawoona basalts. A single small outcrop of massive tholeiitic basalt in the middle of the stratigraphic succession contains highly altered granitic xenoliths which are described in Chapter 3.4.1.4 (Plate 3.1). 2.4.2 Felsic rocks There are three main felsic packages in the Pilgangoora Belt. In the eastern domain, a single felsic unit is up to 1100 m thick and in the central domain, there is a CHAPTER 2: GEOLOGY 23

~200 m thick package immediately beneath and cut by the Coonterunah-Carlindi unconformity and a <50 m thick package ~2.0 km lower in the stratigraphy (Fig. 2.2-3). In the western domain, some thin fine-grained quartzofeldspathic schist horizons are probably deformed felsic volcanics. The three main felsic packages comprise volcanic and volcaniclastic rocks and are interleaved with basalts and cherts and intruded by gabbros. Volcaniclastic rocks are characterised by the presence of clasts, most of which have similar compositions to the host groundmass (Plate 2.2D-E), though some are clearly exotic (Plate 2.2C). Massive felsic units with no obvious clastic material have been classified as felsic lavas, although distinguishing massive felsic lavas from highly altered basalts can be difficult. In many such cases, a gradation between relatively unaltered and highly altered basalts is evident, but where this is not visible, primary features such as remnant ophitic textures in the basalts and abundant amygdales and perlitic fractures in the felsic lavas have been used for discrimination. All of the recognised felsic packages have been confirmed by their trace element compositions (Chapter 5). Felsic volcanics are massive, highly amygdaloidal and gener ally hypohyaline. They are up to ~100 m thick (immediately west of Strelley Pool), although it is unclear whether this thickness represents a single volcanic unit or a pile of thin units. Elsewhere the felsic volcanic units are less than 30 m thick and lie between basalt, chert or felsic volcaniclastic units. The felsic volcanics are predominantly composed of very fine- grained (<0.1 mm) quartz, albite, sericite and chlorite and minor carbonate, epidote and FeTi-oxides. This assemblage is partly the result of intense silica and chlorite alteration (Plate 3.1A). Despite such alteration, primary features are well preserved. Remnant microphenocrysts (up to 0.5 cm, <20 %) of quartz and altered feldspar, conchoidal and perlitic fractures in original glass and spherical amygdales up to 5 mm across are common (Plate 2.2F-G, 2.3A). The amygdales are composed of very fine-grained chlorite-rich rims mantled by axial growth of thin FeTi-oxide-rich quartz fibres with the original vesicle filled by chlorite and quartz, the common alteration assemblage (Plate 2.2G). Eutaxitic textures associated with welding of glass are also common (Plate 2.3A), and may imply that these units were emplaced hot, and so were probably lavas. The felsic volcaniclastic units are quite diverse, but are predominantly composed of highly angular felsic volcanic clasts in a felsic volcanic groundmass. Both the clasts and the groundmass are composed of quartz (10-30 %), albite (10-20 %), sericite (10- CHAPTER 2: GEOLOGY 24

40 %), chlorite (0-20 %), carbonate (<5 %), epidote and FeTi-oxides. Perlitic fractures, amygdales, eutaxitic textures and relict feldspar and quartz crystals are common, and so the rocks are compositionally similar to the adjacent volcanics. In places, however, there are clasts of basalt and black-white chert, indicating that fragmentation, at least for some units, preceded deposition. Individual beds are between ~1 and 100 m thick, and <5 mm thick plane-laminations are rare. In the eastern domain, the topmost volcaniclastic unit is <250 m across and ~25 m thick, forming a large lenticular body. It overlies a thin basalt layer (<1 m), which in turn overlies a ~5 m thick black-white plane-laminated chert. The unit is flanked by basalts, but the contacts are poorly exposed. This volcaniclastic unit comprises abundant poorly rounded black-white chert, basalt and felsic volcanic clasts up to 10 cm long supported in a fine-grained felsic volcanic matrix (Plate 2.2C, E). Clasts with a high aspect ratio are preferentially aligned subparallel to bedding and were probably orientated parallel to the flow direction. Clast size and abundance decreases towards the top of the unit and a fine-grained plane-laminated quartz-feldspar arenite caps the unit. These features are consistent with normal grading in a debris flow and provide the same younging direction as nearby pillowed basalts. This unit probably represents a cross- section through the channel of a debris flow.

2.4.3 Sedimentary Units Black-white plane-laminated ferruginous cherts are the most common sedimentary units in the Coonterunah succession. They are <10 m wide and predominantly composed of very thin (<1 cm) black bands of silicified magnetite and kerogen (±pyrite) with irregularly spaced, 1 to 40 mm thick bands of white chert of varying thickness (Plate 2.3B). The white chert bands are commonly transgressive and are probably related to post-depositional alteration. Some of the cherts are continuous for many kilometres, but others are <100 m long. One chert horizon contains a relatively unsilicified domain of grey to buff brown plane-laminated carbonate with fine-grained fibrous cummingtonite-grunerite (Plate 2.3C), suggesting that some cherts were originally carbonate sediments. The more magnetite-rich cherts may have precipitated from Fe-rich bottom waters. In the eastern domain, chert layers are folded at the metre- scale (Plate 2.3B), but the chert units themselves are not folded at the gross scale. CHAPTER 2: GEOLOGY 25

There is one small outcrop (<3 m wide, <50 m long) of finely laminated fine- grained sediments in the central domain (Plate 2.3D) which comprises chloritoid porphyroblasts within a fine-grained quartz, chloritoid, sericite and carbonate matrix. Its relationship with the adjacent units is unknown and no other exposures have been found, but it may have originally been a fine-grained clastic unit.

2.4.4 Interpretation The Coonterunah Group represents a bimodal volcanic succession dominated by tholeiitic basalts and gabbros. Deposition was on a relatively flat basement as demonstrated by the tabular continuity of many units such as the pyroxene-phyric basalt and cherts. Deposition was clearly subaqueous as shown by the rare but widespread presence of pillowed basalts. The scarcity of pillow forms suggests that deposition may have been dominated by low viscosity sheet flows (Cas & Wright, 1987). Determining whether felsic volcanic and volcaniclastic units were deposited subaqueously can be quite difficult (Cas & Wright, 1987), but as pillowed basalts are present stratigraphically above and below the felsic packages, then they were probably also deposited subaqueously. The presence of sparsely amygdaloidal basalts and scarcity of brecciation in highly amygdaloidal felsic volcanic units suggests that deposition was at high confining pressures and thus in deep water (Buick et al., 1995). The general lack of bedforms in the felsic volcaniclastic units and scarcity of clastic sediments supports a deep marine environment. The chert horizons probably represent chemical precipitates that were deposited during volcanic hiatuses.

2.5 CARLINDI GRANITOID COMPLEX The Carlindi granitoids intrude the entire Coonterunah Group such that the base of the succession is absent. Small granitic bodies penetrate several kilometres into the greenstones, but do not intrude the overlying Warrawoona succession. The granite- greenstone contact is irregular and sharp, with abundant greenstone rafts in the granitoid near the contact. Xenoliths are uncommon elsewhere in the field area, except at Coonterunah Pool. The granitoid is well exposed adjacent to the greenstones, particularly in the western part of the field area, but outcrop becomes more discontinuous away from the greenstone contact. Many ultramafic, mafic and felsic dykes intrude the Carlindi granitoid. The granitoids are predominantly composed of CHAPTER 2: GEOLOGY 26 quartz, plagioclase and microcline with quartz-feldspar ratios defining them as granodiorite or monzogranite. Normative classification shows them to be trondhjemites or granites (Chapter 5), but these terms are not used here as they are difficult to apply based on field criteria alone.

2.5.1 Granodiorite Carlindi granodiorites are light-coloured, massive, coarse-grained (<5 cm), holocrystalline and predominantly composed of plagioclase (40-55 %), quartz (25- 40 %), microcline (10-25 %), mafic minerals (5-10 %) and <1 % muscovite. The mafic assemblage includes chlorite, epidote, hornblende, FeTi-oxides, biotite and clinozoisite. Carbonates and rare sulphides are present in places. Very fine-grained crystalline inclusions (zircon, apatite) are widely disseminated. Plagioclase (An < 20) is commonly euhedral to subhedral with turbid interiors due to very fine-grained sericite (?paragonite) alteration. Concentrically zoned plagioclase is common with the zones typically defined by turbidity contrasts, with rims generally less turbid. Microcline is generally anhedral and contains inclusions of apatite and zircon. Quartz typically forms clear, ragged, serrated, anhedral grains with undulose extinction and minor subgrain development (Plate 3.1E). Chlorite is typically fine- to medium-grained, ragged and commonly associated with fine-grained epidote and clinozoisite. Prismatic green hornblende is abundant in some areas, but is rare overall. Hornblende is generally pleochroic, subhedral and poikilitic with fine-grained quartz and feldspar inclusions. FeTi-oxides are euhedral and randomly dispersed. Medium-grained, subhedral, interstitial muscovite is rare.

2.5.2 Monzogranite Carlindi monzogranites are light-coloured, massive, coarse-grained (<5 cm), holocrystalline and composed of plagioclase (30-40 %), quartz (30-40 %), microcline (25-30 %), mafic minerals (<5 %) and muscovite (<5 %). The mafic assemblage includes chlorite, biotite, epidote, FeTi-oxides and garnet. Fine-grained zircon and apatite inclusions are common. Minor carbonate and sericite alteration is widespread. In general, the textures are remarkably similar to the granodiorites, with zoned subhedral plagioclase and anhedral microcline and quartz. However, microcline is commonly poikilitic with abundant fine-grained plagioclase and quartz inclusions. CHAPTER 2: GEOLOGY 27

Notably, monzogranites contain medium-grained subhedral muscovite and one sample has trace amounts of embayed, corroded fine-grained garnet (Plate 2.3E). Near the Coonterunah-Carlindi unconformity, a quartz-phyric microgranite intrudes coarse-grained granodiorite, indicating some temporal variations among the granitoids. The microgranite contains medium-grained (1-2 mm across), well-rounded, subspherical quartz phenocrysts (~15 %) and subhedral prismatic orthoclase (~25 %). The quartz phenocrysts are rimmed by very fine-grained quartz-sericite (Plate 3.2A), perhaps a reaction texture.

2.5.3 Inclusions Apart from greenstone rafts near contacts with the Coonterunah Group, there are also small greenstone fragments (<20 cm) throughout the Carlindi granitoids, though they are not common. Small wisps (<30 cm) of biotite and chlorite are relatively common and may be relict inclusions. However, the most striking xenolith outcrop is at Coonterunah Pool (Fig. 2.1) where there are abundant fragments of granitic and mafic gneiss hosted by massive monzogranite. These are up to 10 m long, blocky with high aspect ratios and are generally orientated at ~330 °, parallel to nearby biotite wisps. Well-developed gneissic fabrics in the xenoliths are commonly sheared and folded, with these fabrics generally subparallel to the gross orientation of the xenoliths . This outcrop and the suite of xenoliths is described in detail in Chapter 3.

2.5.4 Dykes There are many ultramafic, mafic and felsic dykes in the field area. Most of the mafic dykes are partly altered dolerites, less than 5 m wide and trending NNE-SSW. A thick ~20 to 50 m wide ultramafic dyke intrudes the Coonterunah Group in the central domain, trends north into the granitoid and then continues west to where it is cut by a prominent NNE-trending fault (Fig. 2.1-3). This dyke may be exposed further west where an ultramafic unit intrudes the western greenstones along the Coonterunah- Warrawoona contact. However, smaller ultramafic intrusives are relatively common in the western domain. There are many <5 m wide quartz-albite-microcline-muscovite pegmatitic dykes in the western part of the granitoid. Thick (up to 10 m wide) quartz veins also intrude the granitoid and commonly form prominent ridges. CHAPTER 2: GEOLOGY 28

2.5.5 Interpretation The Carlindi granitoids are predominantly granodiorite and monzogranite, but distinguishing between these rock-types in the field has proven difficult. The presence of hornblende in the granodiorites and medium-grained muscovite and poikilitic microcline in the monzogranites may be useful discriminating criteria. Most of the granitoids were apparently emplaced after deposition of the Coonterunah Group, representing a discrete and significant magmatic event. Intrusive relationships suggest that the monzogranites may by younger then the granodiorites. The presence of granitic gneiss xenoliths, poikilitic microcline and xenocrystic garnet in the granites suggest that they were derived, at least in part, from a crustal source.

2.6 COONTERUNAH-CARLINDI / WARRAWOONA UNCONFORMITY In the eastern Pilgangoora domain, the SW-trending Coonterunah Group is truncated by the E-trending units of the Warrawoona Group (Fig. 2.1-3). The contact is sharp and highly irregular, with Coonterunah basalts and cherts in direct contact with Warrawoona sandstones (Plate 2.3F). No tectonic fabric has been developed parallel to the contact and the basalts either side of the contact have experienced different metamorphic grades. Coonterunah cherts project up to 10 m into the overlying Warrawoona succession and the Coonterunah basalts immediately adjacent to the contact are more altered and finer-grained than those a few metres further away. Sand- filled cracks emanating from the contact penetrate up to 100 m into the Coonterunah basalts. All of these features provide compelling evidence that the contact between the Coonterunah and Warrawoona Groups is an angular unconformity. In this instance, the Coonterunah succession was tilted and eroded such that the chert units formed small resistant ridges on the ancient land surface and the basalts developed a palaeosol with deep weathering fissures. Warrawoona sand filled the fissures and was draped over the palaeosol and the chert ridges. In the central and western domains, the Coonterunah and Warrawoona successions are subparallel, but the presence of the other features indicates that the unconformity is continuous for over 50 km strike length. A Carlindi microgranite is also cut by the unconformity, providing some relative timing constraints to the formation of the unconformity. Indeed, this microgranite intruded granodiorite which in turn intruded CHAPTER 2: GEOLOGY 29 the Coonterunah Group (Fig. 2.2), perhaps indicating that most of the Carlindi magmatism was post-Coonterunah and pre-Warrawoona.

2.7 WARRAWOONA GROUP Immediately overlying the unconformity is a ~20 m thick package of silicified clastic, chemical and volcaniclastic sediments known as the Strelley Pool Chert (SPC). These sediments are overlain by up to 3.5 km of basalts and gabbros which are interbedded with thin horizons of mafic volcaniclastic chert. These are the only exposed units of the Warrawoona Group in the Pilgangoora Belt and extend from the Lalla Rookh-Western Shaw Shear Zone in the eastern domain to the northernmost part of the western domain (Fig. 2.3). Primary sedimentary and volcanic features are very well preserved, except in the northern half of the western domain where deformation, alteration and metamorphism have been quite intense. A <50 m thick ultramafic unit intrudes along the Coonterunah-Warrawoona contact in the western domain and is not considered part of these successions (Fig. 2.1-3).

2.7.1 Strelley Pool Chert (SPC) The SPC is generally 15 to 30 m thick and comprises four main units: quartz-rich sandstone, silicified laminated carbonate, black-white plane-laminated chert and silicified volcanic sediments (Fig. 2.2). Apart from the black-white laminated cherts which are probably post-depositional features, these units form a consistent gross stratigraphy along the entire strike length of the SPC.

2.7.1.1 Quartz-rich sandstone The basal part of the SPC is a fine- to coarse-grained (0.1 to 1.0 mm) quartz-rich sandstone. It is typically 2-6 m thick, but is absent in places where silicified laminated carbonates immediately overlie the unconformity. The contact between the sandstone and Coonterunah volcanics and Carlindi granitoids is generally sharp and highly irregular (Plate 2.3F). Fissures filled with identical sand penetrate up to 100 m into the underlying units. The contact between the sandstone and overlying laminated carbonates is also typically sharp, although thin (<10 cm) quartz-rich sandstone beds are present in the bottom two metres of the laminated carbonates. CHAPTER 2: GEOLOGY 30

The sandstone is predominantly composed of medium- to coarse-grained, well- sorted, well-rounded, spherical quartz (60-80 %) in a siliceous cement. The quartz displays undulose extinction and contains abundant mineral and fluid inclusions (Plate 3.2F). The sandstone also contains clasts of black-white chert, tholeiitic basalt, felsic volcanics, vein quartz and reworked quartz sandstone, all of which are most evident in conglomerate lenses at the base. The siliceous cement typically contains very fine- grained sericite (Plate 3.2F), suggesting that it may have replaced a fine-grained matrix rather than filled voids between grains. The sandstone is grossly bedded (Plate 2.3G), although individual beds can be difficult to distinguish due to intense silicification. However, lateral and vertical grainsize variations, particularly in coarser sediments, indicate that individual beds are probably quite thin (<50 cm) and lenticular over tens of metres. Beds are typically massive, although in places heavy mineral laminations, normal graded beds and planar and trough cross-beds are developed. Lenticular conglomerate beds are quite common near the base of the sandstone and may show coarse-tail grading. Heavy minerals such as zircon, chromite-gahnite, leucoxene, ilmenite, rutile, pyrite, magnetite, monazite, sphene, chlorite and cassiterite are disseminated throughout the sandstone (Plate 3.2F) and in places are concentrated in laminations. The ages of some detrital zircons have been determined by SHRIMP U-Pb geochronology (Chapter 3.6.1.1). Casts of straight symmetrical ripples with peaked to rounded crests and broad troughs are preserved in some places (Plate 2.3H). Wavelengths and heights of the ripples are 6-8 cm and 0.5- 1.5 cm, respectively.

2.7.1.2 Laminated carbonates The thickest unit in the SPC is a plane- to wavy-laminated chert. Laminations are <1 cm thick and cannot be traced for more than a few metres. Intense silicification has generally replaced the original mineral assemblage, although there are domains where coarse-grained dolomite has survived (Plate 2.4A). These domains show progressively increasing degrees of marginal silica alteration, and so dolomite was probably an early component. The dolomite typically forms coarse-grained rhombs that are randomly arranged and probably represent an even earlier alteration episode. Because of its coarse fabric, textures and features at the sub-millimeter-scale have been destroyed. Towards CHAPTER 2: GEOLOGY 31 the base of the unit, thin (<10 cm) quartz-rich sandstone beds are intercalated, whereas towards the top thicker (<30 cm) mafic volcaniclastic beds interfinger. Within the laminated chert lie lenses with subvertical clusters of relict very coarse-grained, prismatic hexagonal crystals with pyramidal terminations (Plate 2.4B). The original mineral has been replaced by coarse-grained dolomite rhombs and silica, but the initial gross crystal morphology resembles that of aragonite. The wavy laminae arch over these lenses and the crystals do not penetrate the laminae. Thus, the clusters appear to be integral parts of the original unit and show that dolomitisation was post- depositional. The laminated carbonates also contain rare conical structures that are up to 15 cm in diameter and vary markedly in size over short distan ces (Plate 2.4C). The stacked lamina flexures consistently dome upwards and the internal laminations are similar to those in the plane- to wavy-laminated carbonates. Indeed, laminae from conical structures can be traced into the plane-laminated carbonates, suggesting a close relationship.

2.7.1.3 Black-white plane-laminated chert Thin black-white plane-laminated cherts are composed of silicified kerogenous laminae and are generally subparallel to the adjacent sedimentary layering. In places, however, these cherts are clearly transgressive and are probably later intrusive features. They are more common near the top of the SPC.

2.7.1.4 Volcanic sediments The topmost part of the SPC comprises grey-green silicified very fine- to coarse- grained sediments and megabreccias. The original mineral assemblage has been almost entirely replaced by silica, though primary textures and features are well preserved. Most of the sediment grains are highly angular and arcuate shards are common (Plate 2.4D), indicating that the sediments were derived from juvenile volcanic glass. Trace amounts of chlorite and leucoxene are widespread and suggest that the original sediments were probably mafic. Tracing individual beds is difficult due to the intense silicification, but where possible, they seem to form lenticular beds <1 m thick. Lutites are massive to plane-laminated with scarce coarse-grained fragments. The laminae are <1 mm thick and are defined by subtle colour contrasts, perhaps CHAPTER 2: GEOLOGY 32 indicating compositional variations. The lutites commonly contain randomly arranged, six-sided crystallites <1 cm long (Plate 2.4E) containing cores of the surrounding sediment. The gross morphology of the crystallites resembles gypsum and the sediment inclusions show that they formed by incorporative growth during diagenesis. Arenites are generally composed of fine- to medium-grained silicified fragments of basaltic glass and commonly display planar laminations and cross-beds. They are typically grain supported and may contain thin lutite beds with silica-altered gypsum crystallites. Megabreccias contain very large (up to 5 m) tabular fragments of lutite and arenite which are arranged subparallel to bedding.

2.7.2 Basalts and gabbros Above the SPC lies a ~3.5 km thick pile of magnesian and tholeii tic basalts. Major compositional changes define a gross stratigraphy and are coincident with thin horizons of mafic volcaniclastic chert (Fig. 2.2). The general stratigraphy is consistent for the entire extent of the Pilgangoora Belt (Fig. 2.3). There are a few semi- conformable gabbros within the succession that may represent the bases of basaltic flows, but other gabbros are clearly intrusive. Metamorphic recrystallisation has been relatively complete, but primary textures and features are extremely well preserved. Pillowed outcrops are common in both tholeiitic and magnesian basalts with exceptional preservation of pipe vesicles, carbonate- and silica-filled amygdales, magma evacuation holes, altered and chilled margins, interpillow hyaloclastite and spalled pillow-margin fragments (Plate 2.4F). Generally the pillows are exposed in cross- section, but rare three-dimensional exposures show that the pillows were subhorizontal tubes extending for over 5 m (Plate 2.4G). Ropey structures resembling those on surfaces of pahoehoe flows have also been noted, in the form of <30 cm thick chaotically folded cords of fragmented basalt in a basalt breccia (Plate 2.4H). Younging is consistently to the south, away from the Carlindi granitoid. Magnesian basalts contain abu ndant coarse-grained varioles (up to 5 cm) which are composed of fine-grained (<0.1 mm) actinolite-albite-clinozoisite (~45:45:10 %) and are set in a fine-grained (<0.1 mm) pale actinolite-chlorite-albite (~33:33:33 %) matrix. Tholeiitic basalts are fine-grained (<0.1 mm), massive and composed of chlorite (20- 40 %), albite (20-40 %), actinolite (10-30 %), prehnite (<20 %), clinozoisite (<10 %), quartz and carbonate and with widespread relict clinopyroxene, plagioclase, hornblende CHAPTER 2: GEOLOGY 33 and FeTi-oxides. Gabbros of both tholeiitic and magnesian composition have larger grain sizes (up to 2 cm) and better preservation of igneous plagioclase, clinopyroxene and Fe-Ti-oxides.

2.7.3 Mafic Volcaniclastic Units Interbedded with the basalts are thin (<5 m) tabular chert u nits composed of silicified mafic volcanic sediments. The units extend for the entire length of the Pilgangoora Belt and provide useful marker horizons in the deformed western domain. The cherts are composed of mafic lutite and arenite, and are compositionally identical to those at the top of the SPC. They also contain many ripples and ripple casts that are generally symmetrical and sinusoidal with peaked crests and have wavelengths and heights of 10-14 cm and 1.4-2.0 cm, respectively.

2.7.4 Interpretation In the Pilgangoora Belt, the Warrawoona Group was deposited on an angular unconformity underlain by the Coonterunah-Carlindi terrain. The basal Warrawoona unit is a quartz-rich sandstone containing clasts of black-white chert, tholeiitic basalt, felsic volcanics and vein quartz. These clasts have similar compositions to units in the underlying Coonterunah succession, their likely source. A significant fraction of the detrital quartz is coarse-grained, which is not a feature of the Coonterunah felsic volcanics. Moreover, detrital quartz is typically well rounded and so was probably even coarser before deposition. The most obvious quartz source is the Carlindi granitoids which contains 25-40 % coarse-grained quartz. The abundance and type of inclusions in the quartz supports such a granitoid source. If so, then substantial amounts of granitoid must have been eroded to account for the volume of coarse-grained quartz in the SPC sandstone. However, the present exposure of the unconformable contact with the Carlindi granitoid is quite limited (~750 m long), and so this probably does not represent the true three-dimensional extent of the contact. Fine-grained detrital quartz may have been derived from the Carlindi granitoid or Coonterunah felsic volcanics. Though most of the sediments in the SPC-sandstone were locally derived, a significant fraction of the eroded basement was not deposited locally. Most of the feldspar from the Coonterunah mafic and felsic volcanics and the Carlindi granitoid has been lost from the sedimentary system, as well as most of the mafic components. CHAPTER 2: GEOLOGY 34

However, relics of these sources remain in the heavy mineral fraction: zircon, monazite, sphene and cassiterite from felsic volcanics and granitoids, and rutile, leucoxene, ilmenite, magnetite, pyrite and chlorite from the Coonterunah mafic volcanics. Pyrite and magnetite may have also been derived from the black-white Coonterunah cherts. However, there are abundant Cr-Zn spinels (chromite-gahnite) in the SPC- sandstone with no obvious local source. Reconnaissance SEM studies show that the compositional range of the chromite-gahnite detritus is fairly continuous and so it was most likely derived from the same source, probably ultramafic to mafic rocks (Deer et al., 1975; Press, 1986; Morton, 1991). However, there are no ultramafic units in the Coonterunah succession and microsampling studies of the Coonterunah basalts with the highest Cr contents (<0.11 wt %) did not uncover any chromite, as the Cr was concentrated in magnetite. Moreover, whole-rock Zn contents are very low in the Coonterunah basalts (<100 ppm; Chapter 4). Thus, it is unlikely that chromite and gahnite were derived from the Coonterunah succession. Possible chromite and gahnite sources occur in the Marble Bar Belt (Gruau et al., 1987; Barley, 1993) and North Pole Dome (Green et al., in prep.) where komatiite units lie stratigraphically below the SPC, but the presence of chromite and gahnite has not been verified in these units. A number of features indicate that the SPC-sandstone was deposited in a shallow subaqueous high-energy environment, including the abundance of well-rounded, coarse- grained, spherical sediment and the presence of lenticular beds with planar and cross- bedding. The symmetrical morphology of the ripples indicates deposition in a wave- dominated environment, with the presence of coarse-tailed grading in conglomerate beds and the well-sorted nature of the sandstone suggesting a near-shore setting where sediment was reworked and fine-grained material removed by persistent wave action (Tucker, 1991). The onset of carbonate deposition marked a change of depositional dynamics to low energy with restricted clastic input. The presence of thin sandstone beds in the bottom ~2 m of the carbonates indicates periods when the earlier conditions briefly prevailed, though these clastic incursions became progressively less frequent and eventually ceased. The conical structures are quite similar to Ephyaltes stromatolites, although the critical microscopic features needed to conclusively prove their biogenic origin (Buick et al., 1981) would have been destroyed during dolomitisation. Nevertheless, their marked diversity in small areas (Plate 2.4C), uniform flexural CHAPTER 2: GEOLOGY 35 direction (pointing up) and the continuity of laminae into plane-laminated domains indicates that they formed early and were related to the rest of the carbonates. If biogenic, then the plane-laminated carbonates may have been algal mats. The presence of radiating crystal clusters resembling aragonite suggests that the carbonate unit formed in an evaporite setting, perhaps in a restricted basin behind a barrier that was occasionally breached allowing deposition of clastic sediments. Mafic volcaniclastic beds appear near the top of the carbonate unit and become progressively more common until the carbonates disappear. These beds mark a change to a volcanic-dominated system. As volcanic input increased, stromatolites and algal mats would have had more difficulty recovering from burial, eventually perishing. The presence of lenticular beds with abundant cross-bedding indicates that the depositional environment was still shallow subaqueous, which is further supported by the abundance of gypsum crystallites in lutites. The mafic volcaniclastics may represent reworked airfall sediments, and thus their presence does not require the removal of the basin barrier. The SPC is overlain by pillowed basalts, indicating continued subaqueous deposition. The abundance of vesicles and amygdales is consistent with a shallow water setting, but not diagnostic. However, interbedded mafic volcaniclastic horizons contain abundant cross-beds, gypsum crystallites and ripples, all of which indicate a shallow water setting. Moreover, the ripples suggest a swell-dominated environment (Tucker, 1991), and thus were probably formed in slightly deeper water than those in the SPC- sandstone.

2.8 POST-DEPOSITIONAL PROCESSES 2.8.1 Deformation There are no bedding-parallel deformational structures or stratigraphic repetitions evident in the Coonterunah or Warrawoona successions and so the present exposures are considered to represent true stratigraphic thicknesses. However, this may not be the case in the western domain where relationships are more equivocal. Before deposition of the Warrawoona Group, the Coonterunah succession was intruded by Carlindi granitoids and this composite terrain was uplifted and eroded prior to development of an unconformity. These events tilted the Coonterunah Group in the CHAPTER 2: GEOLOGY 36 eastern domain, as shown by its angular relationship with the Warrawoona Group. There are small folds in black-white plane-laminated Coonterunah cherts in the eastern domain that are not found in the adjacent basalts or the overlying SPC, and hence are local features which formed prior to Warrawoona deposition. The folds are generally tight, upright and moderately NE-plunging with consistent Z-shapes, though some plunge to the SW and others are moderately open (Plate 2.3B). The folds are consistent with development in a simple shear-dominated regime where strain is preferentially accommodated by weak plane-laminated units (Hobbs et al., 1976). Back rotation of the unconformity to palaeohorizontal shows that the folds originally verged steeply to the east (Van Kranendonk, 2000), away from the ancient topographic high of the Carlindi granitoid where it was cut by the unconformity. This vergence is consistent with folding by granitoid doming, and are similar to folds adjacent to other Pilbara granitoids (Collins. 1989; Van Kranendonk, 2000). However, as an antithetical fold set has not been found on the other side of the granitoid high, in the western domain of the Pilgangoora Belt, such an origin cannot be confirmed. The entire Pilgangoora Belt was also tilted such that bedding is now consistently subvertical. A large component of this tilting probably happened after deposition of the Gorge Creek Group (<3235 Ma, Buick et al., submitted) as these sediments are also steeply tilted. However, the Warrawoona-Gorge Creek contact is slightly angular, indicating some earlier deformation. There is no evidence that the Carlindi granitoids were decoupled from the greenstones during tilting, and so the pre-Warrawoona parts of the Carlindi granitoid must have been also tilted. Nevertheless, tilting was probably related to granitoid doming, perhaps associated with young granitoid intrusions or solid- state diapirism. Regional-scale N- to NE-trending faults dissect the supracrustal successions and penetrate the granitoid. The SPC sediments do not change thickness across these faults and so they were probably not syn-depositional. These faults may be related to post- Warrawoona tilting. In the Lalla Rookh shear corridor, numerous shear-related events of poorly constrained age formed a series of overprinted subparallel structures. Deposition and subsequent deformation of the adjacent Lalla Rookh sediments were probably associated with movement along this shear zone (Krapez, 1993). CHAPTER 2: GEOLOGY 37

Deformation was very complicated in the western domain where at least four discrete post-Warrawoona events are required to explain the structural relationships. A bedding-parallel schistosity was firstly developed in the greenstones and adjacent granitoids, and this was then folded into a regional-scale upright, steeply NE-plunging syncline. An axial planar cleavage was developed during this folding event and is associated with ~2.89 Ga Au mineralisation (Neumayr et al., 1998). On the eastern limb of the fold, there are a series of younger steeply SE-plunging folds, whereas a NE- trending fault truncates the western limb of the large fold.

2.8.2 Metamorphism Metamorphic conditions in low-grade mafic rocks are extremely difficult to constrain due to the presence of relict minerals and the difficulty in analysing fine- grained rocks and establishing chemical equilibrium (Beiersdorfer & Day, 1995). In addition, basaltic rocks require many components to describe their chemical systems and are composed of minerals with diverse solid solution relationships (Spear, 1995). Hence, only broad constraints have been derived here, with semi-quantitative SEM analyses used to confirm the general composition of minerals. The change from greenschist- to amphibolite-facies is transitional and marked by i) the transition from albite to oligoclase, ii) the change from actinolite to hornbl ende, iii) increasing MgO content of chlorite and its eventual disappearance, and iv) the gradual disappearance of epidote (Spear, 1995). The Coonterunah basalts contain actinolite, albite, chlorite and epidote, and so were largely recrystallised at greenschist-facies. However, the widespread presence of low-Al hornblende suggests that they experienced uppermost greenschist conditions, perhaps even transitional amphibolite-facies. Along the contact with the Carlindi granitoid, there are <250 m wide domains where hornblende is more prevalent at the expense of chlorite and epidote, and plagioclase is more calcic, thus indicating slightly higher-grade metamorphic conditions. These domains represent contact aureoles with the granitoid, but they are not ubiquitous along the contact. The change from prehnite-pumpellyite- to greenschist-facies is marked by the disappearance of prehnite and pumpellyite, and the increased presence of zoisite, epidote and actinolite (Spear, 1995). The Warrawoona basalts usually contain actinolite, chlorite, clinozoisite and albite, and so have experienced greenschist conditions. CHAPTER 2: GEOLOGY 38

However, the occasional persistence of prehnite indicates that they have experienced only uppermost prehnite-pumpellyite to lowermost greenschist-facies metamorphism. In the western domain, the presence of abundant hornblende in both the Coonterunah and Warrawoona basalts indicates that they experienced amphibolite-facies conditions. The Pilgangoora successions were probably metamorphosed by hydrothermal action during subaqueous eruption and burial in a hot volcanic pile. Sea-floor metamorphism is typically very heterogeneous, with greater fluid-rock interaction in the most porous units (Mevel, 1981; Evarts & Schiffman, 1983), whereas metamorphism within a volcanic pile is marked by progressively higher metamorphic grades towards the base of the pile (Stern & Elthon, 1979; Levi et al., 1982). However, the Pilgangoora successions show widespread, nearly complete metamorphic recrystallisation with no notable stratigraphic trends of metamorphic grade. Thus, any of these early localised features were probably overprinted during later regional metamorphic events. The Coonterunah succession was clearly metamorphosed prior to formation of the unconformity. The intrusion of the Carlindi granitoids can account for the metamorphic aureole at the Coonterunah-Carlindi contact, and may have provided the thermal input to metamorphose the entire Coonterunah succession. Metamorphism of the Warrawoona succession was related to regional thermal increases, perhaps during burial of the entire belt by later successions. However, thermal changes did not correspond with tectonism as metamorphic minerals do not define deformational fabrics, except in the western domain.

2.9 SYNTHESIS CHAPTER 2: GEOLOGY 39

The Coonterunah basalts were deposited in a deep subaqueous setting coevally with thick felsic volcanic and volcaniclastic units (Fig. 2.4A). Eruption was punctuated by hiatuses, allowing the formation of Fe-rich cherts and plane-laminated carbonates. The Coonterunah succession was intruded by Carlindi trondhjemites and then granites, perhaps inducing widespread upper greenschist-facies metamorphism (Fig. 2.4B). Intrusion may have also coincided with initial uplift and tilting of the Coonterunah succession. The combined Coonterunah-Carlindi terrain was later uplifted again such that the greenstones in the eastern domain were further tilted, and eroded such that the Carlindi granite was exposed (Fig. 2.4C). The now subaerial terrain developed a palaeosol, deep weathering cracks and some topography.

A. Deposition of the Coonterunah B. Intrusion of Carlindi granitoids Group in deep water causes doming and metamorphism

water of Coonterunah Group basalt VV V V V VV V V V V V V V V V chert V V V VV V V V V V V V V felsic volcanic V + V V basalt V x V V VV V V VV x x + + V V V V VV V V V V V V x x x + + + V x ? ? ? x x x x + + x

water V V V V V V V basaV lt V + V V V SPC V x x + x + V + V V V V V + x V x x x x + + x + + V x x V V x x x x x + + x x + + + + V x x x x x x x x x x x x x + + C. Uplift and erosion of D. Deposition in shallow water of Coonterunah- Carlindi terrain Warrawoona SPC and basalts forms an erosional unconformity onto the unconformity

Figure 4.4: Schematic diagram showing the early evolution of the Pilgangoora Belt. The entire terrain was then tilted and faulted such that the present map view largely reflects the ancient cross-section.

The Warrawoona succession was deposited onto this weathering surface (Fig. 2.4D), starting with a quartz-rich sandstone in a high-energy, near-shore, wave- dominated environment. Most of the sediment was probably derived locally, with limited exotic material. Conditions changed to low energy in a restricted basin with the development of possibly biogenic carbonate sediments and evaporites. Barriers protecting this basin were periodically breached allowing deposition of coarse-grained sand onto the carbonates. Mafic volcanic sediments entered the basin eventually CHAPTER 2: GEOLOGY 40 overwhelming carbonate precipitation. This was followed by the deposition of a thick succession of pillowed basalts, with minor hiatuses during which mafic sediments were deposited. Deposition of the Warrawoona succession was maintained in a shallow subaqueous setting. The Warrawoona succession was subsequently tilted and eroded to form a slight angular unconformity with the overlying Gorge Creek sediments. Burial beneath these younger successions probably coincided with regional prehnite-pumpellyite- to lowermost greenschist-facies metamorphism. The entire terrain was tilted again, perhaps due to doming in the Carlindi granitoid, such that bedding in the Warrawoona succession became subvertical. The whole terrain was cut by strike-slip faults and the margins of the belt were modified during various younger deformation events.

2.10 DISCUSSION 2.10.1 Pilgangoora Belt stratigraphy The Coonterunah succession was identified by Buick et al. (1995) who recognised the unconformable contact at the base of the SPC. The succession has since been upgraded to group status with the recognition of three formations (Van Kranendonk & Morant, 1998; Van Kranendonk, 2000). The formational subdivision relies on a correlation of the felsic unit in the eastern domain with that ~2.5 km stratigraphically below the unconformity in the central domain (Coucal Formation). However, the felsic unit immediately beneath the unconformity west of Strelley Pool was not recognised in the GSWA lithostratigraphic compilation (Van Kranendonk & Morant, 1998) and appears on the resulting 1:100,000-scale geological map as silicified and sericitised basalts (Van Kranendonk, 2000). It is here proposed that the felsic unit immediately beneath the unconformity correlates with that in the eastern domain, a relationship amply supported by geochronological and trace element data (Chapters 3 & 5). To accommodate this, the lithostratigraphy should be revised by adding two extra formations. A further consequence of this correlation is that the bulk of the felsic and overlying mafic units in the eastern domain (~2.5 km thick) must have been removed from the central domain prior to Warrawoona deposition. However, the SPC-sandstone (<6 m thick) contains the only clastic sediments that could have been derived from these units, and so the bulk of the eroded terrain must have been completely removed. CHAPTER 2: GEOLOGY 41

Moreover, the exposed part of the Coonterunah Group is up to 6.5 km thick in the central domain, and so adding the eroded units implies that the original stratigraphy may have been >9 km thick.

2.10.2 Pilbara correlations The Coonterunah Group is the oldest Pilbara supracrustal succession. It has only been identified in the Pilgangoora Belt and so there are no correlative sections. In contrast, the SPC and overlying basalts (Euro Basalt) have been identified elsewhere in the Pilbara Craton and represent the topmost units of the regional Warrawoona lithostratigraphy (Fig. 1.4). Thus, the SPC may represent a marker horizon that can be used to correlate between the Pilbara belts. However, recent geochronological studies (Barley et al., 1998; Nelson, 2000) suggest significant temporal diversity of units beneath the SPC, and to accommodate these new data it is possible that the SPC overlies an unconformity in other Pilbara belts. This issue is explored further in Chapter 3. Nevertheless, the SPC is exposed in the North Pole Dome and Coongan, Kelly and North Shaw Belts, where it is remarkably similar to and contains the same three lithostratigraphic units as the Pilgangoora exposures (Fig. 1.3; Lowe, 1980, 1983; Buick & Barnes, 1984; Van Kranendonk, 2000). This suggests that it does indeed represent an easily recognised regional stratigraphic marker horizon. Recently, Lowe (1994) recanted on earlier interpretations that the carbonates were biological (Lowe, 1980, 1983), but new discoveries of more complex conical morphologies in the North Pole Dome have added further support for biogenic origins (Hofmann et al., 1999). Regardless, shallow subaqueous deposition was clearly a common feature for these Warrawoona successions. The Pilbara granitoid complexes are predominantly composed of multiple, discrete, intrusive bodies that record prolonged magmatic histories (Bettenay et al., 1981; Hickman, 1983; Davy & Lewis, 1986; Collins, 1989; Bickle et al., 1993). Airborne radiometric images (Mackey & Richardson, 1997b) show that the southeastern part of the Carlindi complex has markedly lower K-Th-U counts than the Yule, Mount Edgar, Shaw, Corunna Downs and Strelley granitoids (Fig. 1.3). The subdued signature is present where the Carlindi is exposed and so is not due to colluvial cover. The magnetic image (Mackey & Richardson, 1997a) shows that the southeastern Carlindi, Mount Edgar, Corunna Downs and northern Shaw granitoids have low magnetic CHAPTER 2: GEOLOGY 42 responses (<-500 nT), whereas the Yule, Strelley and southern Shaw granitoids have high responses (>-100 nT). Magnetic response is not a function of age because the ~2851 Ma Cooglegong Granite (Nelson, 1998) has a low response, similar to >3400 Ma granitoids. Thus, the southeastern part of the Carlindi Granitoid Complex appears unique in having subdued radiometric and magnetic responses. This suggests that the Carlindi granitoids may represent discrete magmatic events not recorded in the other Pilbara granitoid complexes. Furthermore, those granitoids elsewhere that are broadly coeval with the Carlindi intrusions may have had different petrogenetic histories or be derived from different sources.

2.11 SUMMARY The Pilgangoora Belt contains the only known exposure of the Coonterunah Group, a basalt-dominated, bimodal volcanic succession deposited in a deep water setting. The Coonterunah succession was intruded by Carlindi granodiorites and monzogranites and metamorphosed to uppermost greenschist-facies. The combined terrain was tilted, uplifted and eroded, forming an erosional unconformity onto which the shallow-water sediments of the Strelley Pool Chert were deposited. The SPC sediments were covered by a thick pile of pillowed basalts during continued shallow- water deposition, and then tilted and metamorphosed to prehnite-pumpellyite- to greenschist-facies. The units above the unconformity are the uppermost units of the regional Warrawoona Group. Although the Pilgangoora successions were largely recrystallised during regional metamorphism, primary features and textures are very well preserved, thus making them important rocks for interpreting ancient processes and environments. CHAPTER 3: GEOCHRONOLOGY 43

Chapter 3: GEOCHRONOLOGY

3.1 INTRODUCTION The temporal distribution of rock units can be used to correlate geological events between terranes and constrain the rates of geological processes. Biostratigraphy provides a well-established, straightforward chronologic database for Phanerozoic rocks, but a paucity of indicator fossils makes Precambrian biostratigraphy difficult or impossible. Defining Precambrian ages has therefore relied heavily on radiometric dating using long-lived radioactive isotopes such as K, Rb, U, Th and Sm. Each isotopic system has its merits and defects, so it is critical to correctly define the aims of each geochronologic study. Moreover, preferential concentration of particular radioactive isotopes into specific minerals makes microsampling of rocks more likely to produce interpretable ages. For example, garnet Sm-Nd ages from a gneiss may be metamorphic, whereas zircon U-Pb ages from the same gneiss may be magmatic. Single-grain geochronology is commonly complicated by crystal zoning, though such problems can be resolved, at least in part, by using lasers or microprobes to examine discrete domains within individual grains. In this study, depositional ages of the Pilgangoora greenstones and magmatic ages of the Carlindi granitoid phases were required. Field relationships demonstrate relative age differences between both greenstone packages and the Carlindi granitoids. For instance, the Coonterunah Group was intruded by Carlindi granitoids and then both of these were unconformably overlain by the Warrawoona Group. Precisely dating these various events enables the tectonic evolution to be quantitatively constrained. For instance, how soon after granitoid emplacement was the Coonterunah-Carlindi terrain eroded and did granitoid emplacement continue during or after Warrawoona volcanism? Well-defined ages can also be used to correlate units within the Pilbara and thus determine the timing, extent and nature of craton-wide magmatism. This should provide a better understanding of the tectonic processes involved in the formation and preservation of the Pilbara Craton. Whole-rock isotopic systems commonly re-equilibrate and fractionate during metamorphism and so microsampling of preserved magmatic minerals has been chosen as the most reliable technique to determine magmatic ages in the Pilbara. The most accessible magmatic mineral is zircon because it commonly crystallises from CHAPTER 3: GEOCHRONOLOGY 44 intermediate to felsic melts, favourably fractionates the radioactive parents U and Th from the daughter Pb and is widely considered resistant to later thermal events. Although zircon does not commonly crystallise in mafic rocks, the dominant components of most Archaean greenstones, many Pilbara belts contain conformable felsic volcanic rocks that are amenable to zircon dating. Unfortunately, this is not the case for the Warrawoona Group in the Pilgangoora Belt, an issue which is addressed in this chapter. Since pre-magmatic zircon had been discovered in the Warrawoona Group in the North Pole Dome (Thorpe et al., 1992), zoned and heterogeneous zircons were expected in this study. Therefore, a microprobe technique (Sensitive High-Resolution Ion MicroProbe (SHRIMP)) was used to discriminate between possible heterogeneous zones in zircon grains.

3.2 PREVIOUS WORK Early workers deduced the immense age of the Pilbara greens tones after discovering a number of unconformities near Nullagine (e.g., Maitland, 1905), although this was not universally accepted for many decades (see Hickman, 1983, for a review). The first isotopic studies defined Rb-Sr ages of gneissic granites at 2986 ± 176 and 3059 ± 358 Ma (Compston & Arriens, 1968; de Laeter & Blockley, 1972) and post-tectonic granites at 2614 ± 93 and 2820 ± 64 Ma (de Laeter & Trendall, 1970; de Laeter & Blockley, 1972). In general, these and other early Rb-Sr ages (de Laeter et al., 1975; 1977; Oversby, 1976; Cooper et al., 1980) provided little resolution of granite evolution because their errors were large and the older Rb-Sr compositions were probably re- equilibrated and fractionated during metamorphism. Whole-rock Pb-Pb studies by Oversby (1976) provided smaller errors, but likewise, most ages were probably influenced by metamorphism. These whole-rock ages concur with hornblende and muscovite 40Ar-39Ar ages which record prolonged or punctuated high-temperature metamorphism between 3300 and 2840 Ma (Wijbrans & McDougall, 1987; Zegers, 1996). Model Pb-Pb ages of mineralisation, including galena-bearing stratiform deposits, were thought to date greenstone volcanism (<3506 Ma), but these are imprecise and model-dependent (Richards, 1977; Sangster & Brook, 1977; Richards et al., 1981; Thorpe et al., 1992b). As a result, many of the early studies were not capable of adequately dating the older magmatic events in the Pilbara due to large analytical errors and metamorphic modification of whole-rock isotopic systems. CHAPTER 3: GEOCHRONOLOGY 45

The first reliable greenstone and gneissic granitoid ages were obtained by conventional zircon U-Pb techniques (Pidgeon, 1978a, b; 1984). Zircon was separated from ~100 kg samples and analysed as batches such that each measurement pooled the age of many zircons. Thus, if the zircons were heterogeneous (old relict cores, young hydrothermal rims), then the batch method could produce meaningless mean ages. These first zircon U-Pb ages (3452 ± 16 Ma; Pidgeon, 1978a), however, are still within error of recent, more precise microprobe ages (3474 ± 7 Ma; Nelson, 2000). Importantly, whole-rock Rb-Sr analyses of the same samples produced significantly younger ages and confirmed that the Rb-Sr isotopes were modified during metamorphism (Pidgeon, 1978a). Meanwhile, excellent agreement had been established between Sm-Nd whole- rock isochrons and zircon U-Pb ages in the early Archaean Isua Greenstone Belt, Greenland (Michard-Vitrac et al., 1977; Hamilton et al., 1978). This led to Sm-Nd studies in the Pilbara which defined greenstone ages in the Marble Bar Belt at 3560 ± 32 Ma (Hamilton et al., 1981), 3556 ± 542 Ma (Jahn et al., 1981) and 3712 ± 98 Ma (Gruau et al., 1987). These are older than zircon U-Pb ages from the same rocks (Thorpe et al., 1992; McNaughton et al., 1993; Nelson, 1999, 2000) and suggest that assumptions of comagmatism or a single simple source are incorrect (discussed in Chapter 6). Therefore, Sm-Nd isotopes also fail to constrain the age of the early Pilbara magmatism. Improvements to conventional zircon U-Pb methods (Krogh, 1982a, b) and the coupling of an ion microprobe to a mass spectrometer (Compston et al., 1984) eventually led to widespread success in dating Archaean rocks. Ion microprobe studies by Williams et al. (1983) confirmed the antiquity of the Pilbara granitoids and McNaughton et al. (1988) showed that voluminous granitoid magmatism was coincident with Warrawoona felsic volcanism and that older phases had been preserved within the granitoids. Further ion microprobe (Williams & Collins, 1990; McNaughton, 1993) and conventional (Thorpe et al., 1992) studies started to demonstrate that the craton-wide stratigraphic framework was not as straightforward as proposed (Hickman, 1983) and that the magmatic evolution of the Pilbara was protracted, punctuated and complicated. Current geochronology studies of Pilbara magmatism rely on ion microprobe zircon U-Pb techniques, though this may be supplemented by the use of other isotopic systems. However, the presence of felsic volcanic units in most of the Pilbara greenstone CHAPTER 3: GEOCHRONOLOGY 46 belts and the ubiquitous presence of zircon in the granitoids means that the ion microprobe is usually more than adequate. In addition, zircon is relatively resistant to erosion and so is well preserved in sedimentary environments, thus providing a provenance indicator and maximum depositional ages. Moreover, given precise zircon U-Pb ages, other isotopic systems, such as Rb-Sr, Sm-Nd and Lu-Hf, can provide powerful constraints upon the composition of magmatic sources. For example, Bickle et al. (1993) used Rb-Sr data from the young phases in the Shaw Granitoid Complex to show that they were probably derived from partial melting of a heterogeneous source which separated from the mantle at ~3.5 Ga, possibly the gneissic granitoids. Likewise, to constrain the composition of the mantle source from which the Pilbara basalts were extracted, Sm -Nd and Lu-Hf can be used (see Chapter 6). In the Pilgangoora-Carlindi area, four ion microprobe zircon U-Pb magmatic ages have been previously reported: a Coonterunah felsic volcanic and a Carlindi quartz- phyric microgranite (Buick et al., 1995) and two other Carlindi granitoids (Nelson, 1999). These are discussed below, along with nine new analyses. One important sample from the North Pole Dome which was reported by Buick et al. (1995) is also considered here.

3.3 ANALYTICAL METHODS Ages were obtained using zircon U-Pb methods on the SHRIMP II ion microprobe at Curtin University, Western Australia. The methods used here follow Smith et al. (1998) and vary slightly from those used by Nelson (1999). These differences are minor and do not preclude direct age comparisons. The localities of dated samples are shown in Figure 3.1 and in the geological maps (Appendix - CD).

3.3.1 Sample Preparation Samples (~20 kg) were crushed and milled to pass through a disposable 60 mesh nylon sieve. This powder was washed, the fines decanted and the remaining sample dried. Heavy minerals were separated using tetra-bromoethane, and the mafic and magnetic minerals were removed with a Franz Isodynamic separator. The non-magnetic fraction was processed with di-iodomethane to obtain a final heavy mineral concentrate. Up to 100 representative zircon grains were handpicked and mounted in an epoxy disc with chips of the CZ3 standard zircon, a ~564 Ma gem-quality zircon from Sri Lanka CHAPTER 3: GEOCHRONOLOGY 47 which is free of observable zoning and relatively undamaged by radiation (Pidgeon et al., 1994). The mount was polished to allow analytical access to grain interiors, then cleaned, dried and coated with >99.999 % Au to yield an across-mount resistance of 5- 20 W.

Figure 3.1: Locality of geochronology samples.

3.3.2 Analytical conditions - An elliptical 25 x 20 µm O2 primary ion beam was used to sputter secondary ions from the surface of the zircon. The current at the sample surface was typically between 1.6-3.0 nA. The secondary ions were accelerated into the mass spectrometer and measured. Prior to data collection the beam was rastered across the sample surface for 2-5 minutes to remove surface Pb contamination from an area greater than the analysis area. A typical data acquisition cycle is presented in Table 3.1 and each CHAPTER 3: GEOCHRONOLOGY 48 measurement consisted of seven such cycles. Conditions were not varied between standards and samples.

Table 3.1: Mass analysis cycle for typical zircon data. Species Notional mass Count time (s) Typical count rate (c/s)

90 16 + Zr2 O 196 2 10600 204Pb+ 204 10 0.2 Background 204.04 10 0.1 206Pb+ 206 10 3700 207Pb+ 207 40 1000 208Pb+ 208 10 340 238U+ 238 5 2600 232Th16O+ 248 5 4800 238U16O+ 254 2 17000

Mass resolution (defined as M/DM at 1 % peak height) was typically 4500-4600. A stable hysteresis loop was attained at high resolution prior to data acquisition by sequentially precycling the masses for >5 minutes. Centering of the 196, 206, 238, 248 and 254 peaks was adjusted prior to beam measurement, whereas the other peaks were measured at fixed mass dispersions using empirically derived mass offsets from the 196 (for 204, 204.04) and 206 (for 207, 208) peaks. The mass offset between the 196 and 204 peak (8.170-8.165 amu) was determined by mass spectrum scans of a Pb-rich zircon and remained constant throughout an analytical session (up to 22 hours). Background ion counts were typically 0.040-0.045 amu above 204Pb. Delays of 1-7 seconds were required between each mass analysis with the wait-time dependent on the size of the mass step. Secondary ion beams were measured with an electron multiplier by ion counting. Ion counts were corrected for the intrinsic deadtime of the counting system (measured on galena, range of 32-36 ns). Hydride (H +) isobaric interference on Pb isotopes was insignificant due to sample evacuation for >12 hours prior to analysis. The 176 16 + 208 + high mass tail of Hf O2 was resolved at <1 % of the Pb peak height for the CZ3 standard ( 208Pb = 2.6 ppm). Total Pb sensitivity on the CZ3 standard was typically 10-17 CHAPTER 3: GEOCHRONOLOGY 49

- counts/sec/ppm Pb/nA O2 . Analysis notation is grain.x, where x is the xth analysis of the grain.

3.3.3 Data processing Ion counts were processed using algorithms associated with SHRIMP to obtain uncorrected isotopic ratios and absolute concentrations of U, Th and Pb. Uncertainties in the isotopic ratios were determined by counting statistics and augmented where there was excess scatter about linear trends (each cycle represents a point). The relationship between the measured Pb/U and UO/O ratios as determined on zircon by SHRIMP should follow a quadratic law (Compston et al., 1984), although more recent work suggests that this relationship may be simplified to a power law (Claou é-Long et al., 1995). Here, the U-Pb abundances were based on concurrent data for the CZ3 standard (206Pb/238U = 0.0914) and the empirical observation of a correlation between ln(206Pb*/238U) and ln(UO/U) with a slope of 2.0 (* denotes radiogenic). Long-term studies on the CZ3 zircon indicate that analytical reproducibility of U-Pb was unlikely to be <1 %, and so this is used as a minimum error in data reduction calculations. A fundamental step in U-Th-Pb age determination and a potential source of errors is the correction for non-radiogenic (common-) Pb. All recorded mass-204 counts (background corrected) were taken to be 204Pb+ and used for the common-Pb correction. Most data were corrected by the procedure designed by Compston et al. (1984), where the non-radiogenic component is assumed to be a surface-related laboratory contaminant with the composition of Broken Hill Pb-ore with 204Pb/206Pb = 0.0625, 207Pb/206Pb = 0.9619, 208Pb/206Pb = 2.2285 (Cooper et al., 1969). In general, the 204Pb+ counts were very small (ƒ206% < 0.7, where ƒ206% = (common 206Pb/ total 206Pb) x 100), except for 11 analyses, and so corrections were relatively insensitive to the common-Pb composition. However, if the measured 204Pb counts (background corrected) were greater than six times the average measured on the CZ3 standard during an analytical session (likewise corrected), then the common-Pb composition was assumed to equal that defined by the Cumming & Richards (1975) model 3 average at the apparent age of crystallisation. Uncertainties in the measurement of 204Pb+ were propagated to the corrected ratios. Common-Pb corrected data are presented in Appendix 1. All errors are quoted at ±1s and those in the text are at >95 % confidence limits. 3.3.4 Interpretation procedure CHAPTER 3: GEOCHRONOLOGY 50

The common-Pb corrected data were processed using Isoplot/Ex 2 (written by Ken Ludwig, Berkeley Geochronology Center) to calculate 206Pb/238U, 207Pb/235U and 207Pb/206Pb ages and plot concordia diagrams. Ages were calculated using the decay constants of Jaffrey et al. (1971): 238U ® 206Pb = 1.55125 x 10-10 year-1 235U ® 207Pb = 9.8485 x 10-10 year-1 In general, the calculated ages for each zircon were not equal ( 206Pb/238U < 207Pb/235U < 207Pb/206Pb), although some of the oldest ages were recorded by 206Pb/238U ratios. Preferred ages for interpretation are based on the 207Pb/206Pb ratio because this ratio is the least corrupted by the 204Pb correction and age errors are correspondingly small. Although 232Th/208Pb ratios were also measured, they are not reported here. On a conventional 207Pb/235U versus 206Pb/238U concordia plot (Wetherill, 1958) the majority of the analyses are either concordant or define trajectories indicating recent radiogenic-Pb loss. This is illustrated in Figure 3.2 where all the data are plotted as points (not error ellipses) to define a broad trend between 3520-3460 Ma (magmatic ages) and 0 Ma (Pb loss), although there are some younger ages. Concordia ages can thus be determined by forcing the lower intercept of an isochron through 0 Ma to reflect this recent Pb loss. For samples which may have experienced other thermal re- equilibration events, such as the Carlindi gneissic xenoliths, this procedure may be inadequate and is discussed where appropriate. The calculated errors for 207Pb/235U and 206Pb/238U are highly correlated (generally r > 0.9) which is reflected by the error ellipses for each analysis being elongate with a slope equal to the r-value. Populations have been regressed to a straight line (also forced through the origin to account for recent Pb loss) using a modification of the least squares quadratic method (York, 1969; Ludwig, 1998). The 207Pb/206Pb age is determined at the intercept of the regression line with the concordia curve. Analytical errors are propagated using a maximum-likelihood estimation algorithm (Titterington & Halliday, 1979), and no decay-constant errors are used for calculating the concordia curve intercept. CHAPTER 3: GEOCHRONOLOGY 51

0.8

3400 3600 3200

0.6 3000 U 8 3 2 /

b 0.4

P 2000 6 0 2

0.2

0 0 10.0 20.0 30.0 207 235 Pb/ U .

Figure 3.2: Discordance is dominantly a function of recent Pb-loss from zircons with an age between ~2900 and 3500 Ma. All data points plotted without error ellipses.

To define a meaningful isochron from a population of single analyses, it is crucial to show that the population errors can be explained within the limits of the analytical uncertainty. In other words, the scatter of data around the isochron should be solely related to analytical errors. If not, the added scatter must be due to inclusion of analyses which are not part of the true isochron population and so represent geological outliers. Examples of such outliers include inherited grains in a mainly magmatic population or grains with some unaccounted Pb loss. An index is therefore required to measure whether a population does indeed define a meaningful isochron. The mean square of weighted deviates (MSWD) is employed here (McIntyre et al., 1966; Wendt & Carl, 1991), although other statistical approaches have also been defined (Wendt, 1969; York, 1969; Brooks et al., 1972). Essentially, the MSWD measures whether the variance of the scatter around the isochron is equal to the variance of the analytical error. This is achieved by calculating the ratio between these variances (actual errors / expected errors), and hence, the MSWD has an F-distribution and produces a workable index. With large populations of both isochron data and replicate analyses of the standard (analytical error) there are three possible scenarios; i) MSWD = 1, the analytical error can account entirely for the isochron scatter, ii) MSWD < 1, the analytical errors have been overestimated, and iii) MSWD > 1, there are geological outliers. Therefore, with large sample populations the aim is to have MSWD £ 1. Zircon analyses, however, do not form large populations and so the critical MSWD value must be altered to allow for CHAPTER 3: GEOCHRONOLOGY 52 smaller sample sets. These can be derived using the F-distribution for the appropriate population sizes (degrees of freedom), and will change between samples depending on the number of data points defining the isochron and the number of replicate analyses of the CZ3 standard zircon (analytical error). For example, an isochron defined by 10 analyses and 10 standard replicates would have a critical MSWD of 3.07 at a 95 % confidence level (Brooks et al., 1972). A population value above this would imply geological errors. It has been suggested that an MSWD = 2.5 is an acceptable cut-off for isochron definition (Brooks et al., 1972), and a similar boundary is employed here. The MSWD index can thus be used to identify and remove statistical outliers. In some cases, however, a significantly low MSWD may be determined for a population which includes possible suspect points. An excellent example comes from felsic volcanic sample 70649, which has twenty overlapping analyses and two obvious and two probable visual outliers (Fig. 3.3). Removal of the 3 farthest outliers defines a statistically significant population with 21 points (MSWD < critical; Table 3.3). The age of the other two outliers are ~3.8s and ~2.7s from the population mean, respectively (based on sample and not population errors), so should they be included in the population? Probably not, since they are quite removed from the population mean and their statistical variability has been swamped by the other data. Removal has two consequences; it increases the population age and lowers the MSWD (improves the population statistics even further). In other examples the age error also decreases. Of course, the iterative removal of outliers could continue ad infinitum, and so an arbitrary cut off at 2 s is invoked. Interestingly, the fifth outlier was not obvious on the concordia diagram (Fig. 3.3), and so this numerical procedure cannot be abbreviated. This example also portrays the problem of skewness; nine of the ten farthest outliers were less than the population mean and so progressive removal kept increasing the age. When this skewness is less marked or absent the age stabilises because successive points average out, although the MSWD can keep decreasing. In summary, to supplement the MSWD index, the largest outliers within a significant population were assessed. The following procedure was consequently used to define a significant population: 1) all data were plotted on a conventional concordia diagram ( 207Pb/235U versus 206Pb/238U), 2) visual outliers from the dominant population were removed, CHAPTER 3: GEOCHRONOLOGY 53

3) the remaining data were replotted and a 207Pb/206Pb age, error and MSWD were computed, 4) individual analyses more than 2 s from the isochron age (using the 207Pb/206Pb error of the sample, not the population) were treated as outliers and removed from the population, 5) steps 3 and 4 were repeated until i) there were no remaining outliers, ii) removal of the outliers had no influence on the age or error of the population, or iii) removal did not substantially change the MSWD. Preference was given to large zircon populations rather than better statistical significance. This was deemed important to avoid the creation of many small populations which may not have any geological meaning. In general, the aim was to find dominant populations. Small secondary populations have been defined for a few samples but their pooled results are rarely statistically significant and may have little geological value.

Table 3.2: Calculations for sample 70649. The critical values are defined at the 95 % confidence level by the F-distribution of the number of points and 15 standard replicates (Marascuilo & Serlin, 1988). Ages are 207Pb/206Pb. No. of Age of analysis Isochron age MSWD critical points removed (Ma) (Ma) 24 3497 ± 35 73 2.31 23 3226 ± 5 3504 ± 21 26 2.32 22 3322 ± 5 3512.1 ± 7.1 3.0 2.33 21 3470 ± 8 3517.1 ± 4.1 0.54 2.35 20 3484 ± 9 3517.7 ± 4.1 0.32 2.36 19 3499 ± 7 3518.2 ± 4.2 0.21 2.37 18 3506 ± 8 3518.5 ± 4.2 0.17 2.38

Aside from their isotopic ages, another useful comparison within and between zircon populations is their U-Th elemental contents. For example, Coonterunah zircons have low U and Th concentrations relative to the Carlindi granitoids and so can be used as an added criteria for discriminating detrital zircons from a mixed source. Backscattered electron (BSE) and cathodoluminescence (CL) images can also be used to define zircon populations by comparing the gross grain morphologies (crystal habit, CHAPTER 3: GEOCHRONOLOGY 54 size, shape) or more specific features such as zoning or fracturing. In many cases, the images show that analyses were collected from metamict zones, and thus represent younger recrystallised portions of the zircon. Images from samples 520698, 080698, 090698 and 100698 were collected on a Phillips JEOL-6400 SEM at the University of Western Australia and are presented in the Appendix CD. The CL images are washed out because the Pilbara zircons are very metamict and have inherent low luminescence.

3.4 COONTERUNAH GROUP 3.4.1 Results Three samples were collected from felsic volcanic units to determine the age and age variation of Coonterunah deposition. One of these samples yielded only three zircons and provides no reliable age. The other two volcanic samples have large, well- defined zircon populations within error of each other and with few statistical outliers. A fourth sample consists of small granitic ènclaves within a Coonterunah basalt. The ènclaves yielded zircons with U-Pb ages younger than the interpreted age of the basalt. If the ènclaves are xenoliths, then these ages may reflect zircon recrystallisation during protracted greenstone metamorphism, and thus, original old ages have not been preserved.

3.4.1.1 Central dacite, Sample 70649 This sample was collected from a massive, amygdaloidal volcanic dacite in the central Pilgangoora domain, ~ 3 km southwest of Strelley Pool and immediately beneath the unconformity (Fig. 3.1). The rock was pale brown and silicified, but otherwise unaltered. Twenty-four analyses were obtained from 22 zircons of which 16 were concordant to slightly discordant (Fig. 3.3). The discordance pattern is consistent with several episodes of radiogenic Pb-loss, although most of this loss was probably recent. Nineteen analyses define a pooled 207Pb/206Pb age of 3518 ± 4 Ma (MSWD = 0.21) and this is interpreted to be the eruption age. These volcanic zircons have low U (58-173 ppm) and Th (27-104 ppm) concentrations. Four of the five unpooled analyses define a chord between this interpreted magmatic age and 1699 ± 4 Ma. This may signify a Proterozoic Pb loss event, but these analyses are highly discordant and so did not equilibrate at ~1699 Ma. Three of these CHAPTER 3: GEOCHRONOLOGY 55 analyses were obtained from grain #7, two of which have extremely high U and Th concentrations. No equivalent age has been measured in other Pilbara samples and these analyses probably reflect local disequilibrium zircon growth of no geological importance.

0.8 Intercept at 3518 ± 4 Ma MSWD = 0.21 3550 n = 19 3450

24.1 0.7 3350

3250 7.1

3150 U 8

3 3050 23.1 2 / 0.6 b 7.3 P 6 0 2 7.2

0.5

to origin 0.4 18 22 26 30 34 207Pb/235U

Figure 3.3: Concordia plot of sample 70649, central dacite, Coonterunah Group.

3.4.1.2 Eastern rhyolite, Sample 70601 A brecciated hyaloclastic rhyolite was collected from the eastern part of the Pilgangoora Belt (Fig. 3.1). The rock is mid-green and has undergone siliceous, chloritic and carbonate alteration. Angular blocky rhyolite fragments up to 10 cm across form a framework, separated by smaller fragments of spalled glass in a siliceous cement (Plate 3.1A). It yielded clear to pale-brown, euhedral to subhedral zircon with faint subhedral zoning and no distinct core. Twenty-seven grains were analysed and there are no highly discordant results (Fig 3.4). All 27 zircons define a chord signifying recent Pb-loss and have a pooled 207Pb/206Pb age of 3515 ± 3 Ma (MSWD = 0.25; Buick et al., 1995). This is interpreted to be the age of eruption. All grains are within 2 s of the population mean. These volcanic zircons also have very low U (38-101 ppm) and Th (19-78 ppm) concentrations, which are marginally less than for sample 70649. CHAPTER 3: GEOCHRONOLOGY 56

0.85 Intercept at 3515 ± 3 Ma MSWD = 0.25 n = 27 0.75 3600 3500

3400

U 3300 8 3 2 / 0.65 3200 b

P 3100 6 0 2

0.55

0.45 to origin 18 22 26 30 34 207Pb/235U

Figure 3.4: Concordia plot of sample 70601, eastern rhyolite, Coonterunah Group.

3.4.1.3 Lower rhyolite, Sample 70660 This sample was collected from the stratigraphically lowest felsic volcanic unit in the central Pilgangoora domain (Fig. 3.1). The sample was from a green, spherulitic rhyolite and yielded only three zircons; two are highly discordant and the other plots above concordia (Fig. 3.5). No reliable age could be established for this unit. CHAPTER 3: GEOCHRONOLOGY 57

0.85 Intercept at 3515 ± 3 Ma MSWD = 0.25 n = 27 0.75 3600 3500

3400

U 3300 8 3 2

/ 3200

b 0.65

P 3100 6 0 2

0.55

0.45 to origin 18 22 26 30 34 207Pb/235U

Figure 3.5: Concordia plot of sample 70660, lower rhyolite, Coonterunah Group.

3.4.1.4 Granitic xenoliths, Sample 520798 This sample consists of isolated granitic ènclaves within tholeiitic basalt and was collected in the central Pilgangoora domain, ~1 km stratigraphically beneath sample 70660 (Fig. 3.1). The exposure is ~3 x 4 m and represents the only known exposure of small granitic ènclaves within the Coonterunah basalts. The ènclaves comprise ~40 % of the outcrop and are generally less than 20 cm across. They are irregularly distributed and vary from diffuse irregular, amoeboid bodies with poorly defined, embayed, resorbed margins to more tabular, well-defined bodies (Plate 3.1B). There are neither chilled margins nor contact metamorphic features in the ènclaves or the host basalt. The ènclaves contain clear medium-grained, anhedral quartz and intensely sericite-carbonate altered, medium-grained subhedral feldspar (Plate 3.1C-D). Ghost plagioclase twinning and more altered cores are discernible, but rare. In general, the texture of the è nclaves is similar to the Carlindi granitoids, with subhedral feldspar enclosed by anhedral quartz, though the alteration and general metamorphism have obscured the finer features (Plate 3.1E). The host basalt is massive and contains very fine-grained hornblende, epidote, chlorite and leucoxene, not dissimilar to nearby basalt outcrops. The è nclave-rich exposure has no features to indicate that the host rock was definitely volcanic, but adjacent basalts have small spheroidal amygdales and an outcrop ~50 m up-section CHAPTER 3: GEOCHRONOLOGY 58 contains well-preserved amygdaloidal pillows. There are no intrusive features or metamorphic variations between these outcrops, and so the ènclave host is interpreted to be part of the Coonterunah succession. The ènclaves are fragmented, but since there is no evidence of post-magmatic deformation this is probably an early feature. The coarseness of the quartz and feldspar precludes the ènclaves forming as quenched felsic melts, although the subhedral feldspar crystals indicate that the ènclaves were once magmatic. Indeed, the texture indicates that feldspar crystallised first and that quartz was interstitial, consistent with typical granitoids. Hence, the ènclaves are probably granitic xenoliths which were fragmented while being incorporated into the basaltic melt. The sample inevitably included some of the enclosing basalt, although this was kept to a minimum. Whole-rock elemental analyses of the Coonterunah basalts (Chapter 4) show that Zr was highly incompatible, and thus zircon did not readily crystallise from these melts. Therefore, the zircons separated are probably representative of the è nclaves and not the basalt. Moreover, in these basaltic melts, individual xenocrystic zircons would have been out of equilibrium and, given enough time, they would have dissolved. The ènclaves, however, may have been xenolith arks within which older zircons survived. Analysed zircons are equant to elongate and rounded to subhedral. Their largest dimension is typically <140 µm long. Many of the equant grains have smooth irregular, cuspate margins, possibly due to resorption. Zoning is apparent in both BSE and CL images, and is not necessarily concentric. These zones are truncated in many grains (Plate 3.1F-G). Twenty-five analyses were obtained from 25 zircons of which 16 are concordant to near-concordant (Fig. 3.6). The discordance pattern is consistent with multiple Pb- loss events, including one recent event. The 16 concordant to near-concordant analyses span ~40 myr, and the oldest single zircon has a 207Pb/206Pb age of 3504 ± 4 Ma, significantly younger than the interpreted eruption age of the Coonterunah host. A population of 13 zircons, including two discordant analyses, has a pooled 207Pb/206Pb age of 3499 ± 4 Ma (MSWD = 0.19). Four other concordant analyses are ~2s younger than this age and their inclusion in the above population produces a pooled age of 3496

± 3 Ma (MSWD = 1.01). Older grains are typically equant and rounded (Plate 3.1F-G), whereas younger grains have high aspect ratios. There are no obvious contrasts in the CL or BSE images or the Th-U-Pb systematics of the different aged grains. CHAPTER 3: GEOCHRONOLOGY 59

0.8 Intercept at 15.1 3499 ± 4 Ma 30.1 MSWD = 0.19 34.1 3550 n = 13 3450

4.1 0.7 3350 3.1

3250 22.1 10.1 3150 21.1 U

8 3050

3 13.1 2

/ 0.6 24.1 11.1 17.1 b P 6 0 2

0.5

to origin

0.4 18 22 26 30 34 207Pb/235U

Figure 3.6: Concordia plot of sample 520798, granitic ènclaves in tholeiitic basalt, Coonterunah Group.

At face value, the ènclaves can be interpreted to have a crystallisation age of ~3499 Ma, and thus, are significantly younger than the host Coonterunah basalt (>3518 Ma). If so, the ènclaves may be intrusive and the textures may have been incorrectly identified. This age does not correlate with any known Carlindi intrusive, although many detrital zircons of equivalent age have been obtained from the Warrawoona sandstone immediately above the Coonterunah-Carlindi unconformity (sample 98OB5002) and single analyses from other samples are within error of this age. The ènclaves may thus reflect an important intrusive event at ~3499 Ma. However, many ènclaves have resorbed, nebulous margins which are consistent with ènclave inclusion while the basalt was molten. If so, they should be at least as old as the host basalt, maybe older. The host basalt is clearly part of the ~3517 Ma Coonterunah succession and so, in this case, the zircons would have deceptively young U-Pb ages. The data would then most easily be explained by isotopic re-equilibration at ~3499 Ma, perhaps during regional metamorphism. Concordant to near-concordant analyses, however, span ~40 myr, which suggests that such resetting would not have been associated with a single thermal pulse, but that there were protracted or punctuated thermal events between ~3504 and ~3465 Ma. In the central and eastern Pilgangoora CHAPTER 3: GEOCHRONOLOGY 60

Belt, the Coonterunah succession was metamorphosed prior to Warrawoona deposition, and so there must be an older regional metamorphic event. Zircon U-Pb ages from the Carlindi trondhjemite and granite indicate thermal pulses at ~3484 and ~3468 Ma, respectively, which may have been associated with greenstone metamorphism. Contact aureoles adjacent to large granitoid bodies demonstrate at least local affects from these intrusions. Thus, the probable timing of metamorphism is coincident with the range of ènclave ages. If the ènclaves are xenoliths and their zircons record younger metamorphic ages, then the zircons must be either newly crystallised metamorphic grains or older recrystallised xenocrysts. The oldest analysed grains are generally equant with irregularly cuspate to rounded boundaries, and, in places, grain boundaries truncate internal zoning (Plate 3.1E-F). These features are consistent with dissolution of larger grains. For these grains to be solely metamorphic would require growth and dissolution from fluids with fluctuating Zr saturation. Coonterunah basalt compositions do not show widespread, outcrop-scale Zr mobility (Chapter 4), and it is unlikely that local conditions would fluctuate enough to alternatively crystallise and dissolve zircon without metasomatic input. It has also been reported that many metamorphic zircons have elevated U concentrations with correspondingly small Th/U ratios (<0.1; Schaltegger et al., 1999; Vavra et al., 1999; Hartmann et al., 2000). Although the ènclave zircons have higher U contents (77-324 ppm) than the Coonterunah felsic volcanic zircons (38-173 ppm), their Th/U ratios (0.32-0.71) are not in the metamorphic range. In summary, there is little evidence that the ènclave zircons first crystallised during metamorphism. The most tenable interpretation is that the zircons are recrystallised xenocrysts and the ènclaves were derived from ~3517 Ma Coonterunah subvolcanic intrusions or >3517 Ma granitic basement. Since these xenocrysts have yielded only younger ages they must have been severely affected by metamorphism and their origins cannot be deduced from their ages. In addition, grain morphology and internal features cannot be used as discriminants because zircons derived from either source could have undergone resorption by basaltic magma (producing rounded and cuspate grains with transgressive boundaries) and later metamorphic recrystallisation. Compositional comparisons with local (eg. Carlindi) or global (eg. Heaman et al., 1990) zircon populations are probably also uninformative due to element mobility during recrystallisation. CHAPTER 3: GEOCHRONOLOGY 61

The validity of this interpretation depends on metamorphic recrystallisation erasing the original igneous U-Pb ages of the zircons. This has indeed been recorded in more evolved rocks, such as granites and granodiorites, but typically some of the original domains are preserved and there are recognisable metamorphic overgrowths (Pidgeon & Wilde, 1998; Hartmann et al., 2000). A similar outcome would be expected for zircons within granitic xenoliths because the xenoliths should protect their contents. Since the xenoliths are quite small, however, there may have been considerable disequilibrium with the basalt host. For example, uranium concentrations are ~4 orders of magnitude lower in the Coonterunah basalts than in the zircons, and so during recrystallisation the zircons may have lost both U and Pb. Indeed, the super-concordant composition of many of the zircons is consistent with U loss. Coincident zircon recrystallisation has not been recorded by discordant grains in the Coonterunah felsic volcanics although these units have also been altered and metamorphosed. This contrast may have been caused by zircons maintaining their equilibrium within a Zr-rich felsic host. It is proposed, therefore, that the zircons are metamorphically recrystallised xenocrysts in xenoliths and that modification of their U-Pb systems has been relatively complete. Xenoliths may have come from either of two possible sources and metamorphism may have spanned ~40 myr.

3.4.2 Interpretation of Coonterunah geochronology There are no greenstones exposed between the eastern and centr al Pilgangoora domains (Fig. 3.1) and as the rocks in these two domains have different strike directions and unconformity relationships, then they may not be the same succession. The felsic volcanic samples from the two domains, however, have interpreted eruption ages which are within error of each other. These data are very robust and demonstrate the lateral extent and continuity of the Coonterunah succession. In addition, the ages suggest a possible stratigraphic correlation between these felsic volcanic units, which is also consistent with their whole-rock elemental composition (Chapter 5). Pooling of these two datasets provides a depositional age of the middle part of the Coonterunah succession at ~3517 ± 3 Ma (MSWD = 0.25, n = 46). This correlation also suggests that in the central Pilgangoora domain the Coucal Formation is immediately beneath the unconformity and the entire Double Bar Formation has been removed (cf. Van CHAPTER 3: GEOCHRONOLOGY 62

Kranendonk, 2000). Therefore, the Coonterunah Group may have been >2 km thicker than indicated by the thickest (~5 km) preserved section. Isolated ènclaves within Coonterunah basalt are interpreted to be xenoliths with completely recrystallised zircons. These zircons have ages between ~3504 and ~3465 Ma which are consistent with other constraints on the timing of post-Coonterunah metamorphism and igneous intrusions. A population of 13 zircons defines a 207Pb/206Pb age of 3499 ± 4 Ma and may date the onset of metamorphism. However, further evidence is required before a significant thermal event is placed at this age because it may also be the result of incomplete equilibration during a later metamorphic event. The xenoliths may have been derived from Coonterunah subvolcanic intrusions or older granitic basement. In either case, the presence of granitic xenoliths is consistent with the ubiquitous crustal contamination found in all Coonterunah basalts (see Chapters 4-6).

3.5 CARLINDI GRANITOIDS 3.5.1 Results Seven samples have been dated from within the Carlindi Granitoid Complex. Two samples were collected from beneath the Coonterunah-Carlindi unconformity and provide valuable constraints for the timing of erosion. Two samples from within the Carlindi Granitoid Complex were dated during GSWA mapping and provide further constraints on the temporal evolution of the complex (Nelson, 1999; Van Kranendonk, 2000). The discovery of gneissic xenoliths at Coonterunah Pool led to three samples being dated in the hope that the xenoliths may represent pre-Coonterunah basement. Granitoid dating has been integrated with other geochemical studies to determine the petrogenetic evolution of the complex (Chapter 5, 6). Based on their quartz-alkali feldspar-plagioclase (QAP) ratios nearly all of the samples from the Carlindi complex have a granodiorite composition. Consequently, discriminating granitoids in the field is extremely difficult. Subdivision of the granitoid samples using their feldspar ratios shows that three are granites (94058, 090698, 100698), one is a trondhjemite (95037) and one gneissic xenolith is a tonalite (080698). The two GSWA samples could only be classified according to their mineralogical compositions and thus are granodiorites (Nelson, 1999; Van Kranendonk, 2000).

3.5.1.1 Unconformity granite, Sample 94058 CHAPTER 3: GEOCHRONOLOGY 63

This sample was collected from a massive quartz-phyric microgranite ~150 m beneath the Coonterunah-Warrawoona unconformity (Fig. 3.1; Buick et al., 1995). The microgranite comprises medium-grained (4 mm) rounded quartz (~10 %) and subhedral microcline (20 %) and plagioclase (10 %) phenocrysts in a fine-grained (<0.5 mm) quartz (20 %), feldspar (30 %) and biotite (10 %) groundmass (Plate 3.2A). Small, elongate zircon grains are included in quartz and feldspar crystals. The microgranite intrudes the Coonterunah greenstones and Carlindi trondhjemite (Sample 95037), and also extends to, and is truncated by, the unconformity. There is a possible palaeosol immediately beneath the unconformity (up to 10 m) where the microgranite somewhat decreases in grainsize and the feldspar has been completely altered to sericite. Sand- filled fractures related to the unconformity penetrate the microgranite by up to ~100 m. The above relationships indicates that the microgranite must be younger than the Coonterunah Group and Carlindi trondhjemite but older than the unconformity. As a result, the sample provides an important constraint on the timing of uplift, erosion and Warrawoona deposition. The sample yielded morphologically homogeneous, finely-zoned, euhedral zircons. Twenty-one grains were analysed of which 17 are concordant to near- concordant (Fig. 3.7). The discordance pattern suggests one main, recent Pb-loss event.

A population of 13 grains has a pooled 207Pb/206Pb age of 3468 ± 4 Ma (MSWD = 0.15), which is interpreted to be the age of crystallisation (Buick et al., 1995). Hence, the unconformity must be younger. In addition, this age is within error of widespread Pilbara felsic volcanism (~3466 Ma Duffer Formation; Pidgeon, 1978a; Thorpe et al., 1992a; McNaughton et al., 1993; Nelson, 1999, 2000), and so might be the intrusive equivalent. The microgranite zircons have significantly greater U (136-452 ppm) and Th (50-264 ppm) contents than those in the Coonterunah felsic volcanics, but U contents are similar to Duffer zircons (Nelson, 2000). This also supports a link between the microgranite and Duffer Formation, but indicates petrogenetic differences with the Coonterunah felsic volcanics. More elemental analyses are required to make this linkage robust. The two older zircon grains (~3504, 3489 Ma) are probably xenocrysts and show that the granite may have been derived from or assimilated older crust. Reaction textures in the microgranite, as demonstrated by the resorbed, rounded quartz grains (Plate 3.2A), indicate that quartz too was not stable in the magma, and may also have CHAPTER 3: GEOCHRONOLOGY 64 been assimilated. Derivation of the granite from an older crustal source is also supported by their whole-rock elemental and isotopic compositions (Chapter 5, 6). Two of the six remaining analyses are near-concordant and probably reflect hydrothermal growth or recrystallised metamict zircons.

0.8 Intercept at 3468 ± 4 Ma MSWD = 0.15 n = 13 16.1 3550 6.1 0.7 3.1 3350

3250 5.1 3150 11.1 9.1 U 8

3 3050 2 / 0.6 b P 6 U 0 8 2 3

2 3200 /

0.6 b P

6 2800 0 2 2400 8.1 0.5 0.4 2000

21.1 0.2

to origin 207 235 Pb/ U 10 20 30 0.4 18 22 26 30 34 207Pb/235U

Figure 3.7: Concordia plot of sample 94058, Carlindi granite, beneath the unconformity (Buick et al., 1995).

3.5.1.2 Unconformity trondhjemite, Sample 95037 This sample was collected from a massive trondhjemite ~30 m from where it is intruded by the microgranite (Sample 94058) and ~150 m beneath the Coonterunah- Carlindi unconformity (Fig. 3.1). It comprises coarse-grained (up to 2 cm), euhedral to subhedral microcline (20 %) and plagioclase (30 %) in a finer-grained (<1 mm) microcline (5 %), plagioclase (20 %), quartz (25 %) and biotite (5 %) groundmass. This is the most commonly exposed rock-type in the Carlindi complex and it probably represents a major intrusive event. However, exposure is very discontinuous, and, as a result, the Carlindi complex may be composed of many similar trondhjemite bodies of dissimilar age. Five grains were analysed of which 2 are near-concordant (Fig. 3.8). The discordance pattern suggests one recent Pb-loss event. All five analyses define a pooled CHAPTER 3: GEOCHRONOLOGY 65

207Pb/206Pb age of 3479 ± 11 Ma (MSWD = 0.57), which is interpreted to be the age of crystallisation. This is consistent with the trondhjemite being intruded by the ~3468 Ma Carlindi microgranite, and is also within error of other Carlindi granodiorites (Sample 153188, 153190). The zircons have low U (63-130 ppm) and Th ( 26-55 ppm) contents.

0.8 Intercept at 3479 ± 11 Ma MSWD = 0.57 n = 5 3550 3450 0.7 3350

3250

3150 U 8

3 3050 2 / 0.6 b P 6 0 2

0.5

to origin 0.4 18 22 26 30 34 207Pb/235U

Figure 3.8: Concordia plot of sample 95037, Carlindi trondhjemite, beneath the unconformity.

3.5.1.3 Wilson Well granodiorite, Sample 153188 This sample was collected during GSWA mapping in the central southern part of the Carlindi Granitoid Complex, ~12 km northeast of the above two samples (Fig. 3.1; Van Kranendonk, 2000). It was considered to be a localised rock-type and was described as an unfoliated, fine-grained biotite monzogranite (Van Kranendonk, 2000). A petrographic description by Nelson (1999), however, showed that the sample comprised 40 % plagioclase, 30 % quartz and 12 % microcline, and so is more correctly classified as a granodiorite. Acicular zircon crystals up to 0.05 mm long were included in quartz and feldspar. Separated zircons were dark brown to black, euhedral or irregular and up to 400 µm long. Internal zoning was common and many grains were metamict. Twenty-two analyses were obtained from 22 zircons, of which 18 are concordant to near-concordant (Fig. 3.9). The discordance pattern is consistent with a CHAPTER 3: GEOCHRONOLOGY 66 single recent episode of Pb loss. All 22 analyses define a single pooled 207Pb/206Pb age of 3484 ± 4 Ma (Nelson, 1999; MSWD = 0.82). Three highly discordant analyses have elevated common-Pb contents, but their removal from the population has no significant affect. Likewise, removal of a 3500 ± 7 Ma zircon, which is ~2 s from the population mean has no affect. Consequently, the age of crystallisation is interpreted to be 3484 ± 4 Ma, which is within error of the unconformity trondhjemite (Sample 95037). The zircons from this sample have a large range of U (40-500 ppm) and Th (11-190 ppm) concentrations.

0.8 Intercept at 3484 ± 4 Ma MSWD = 0.82 3550 n = 22 3450 0.7 3350

3250

3150 U

8 3050 3 2

/ 0.6 b P 6 0 2

0.5

to origin 0.4 18 22 26 30 34 207Pb/235U

Figure 3.9: Concordia plot of sample 153188, Carlindi granodiorite, Wilson Well (Van Kranendonk, 2000).

3.5.1.4 Shilliman Well granodiorite, Sample 153190 This sample was also collected during GSWA mapping, ~3 km west of the previous sample (Fig. 3.1; Van Kranendonk, 2000). It was believed to represent the dominant lithology of the southwestern part of the Carlindi complex and was described as an unfoliated to weakly foliated biotite-hornblende leucogranite (30-40 % plagioclase, 30 % K-feldspar, 25-30 % quartz, 5-8 % hornblende, ?biotite; Van Kranendonk, 2000). CHAPTER 3: GEOCHRONOLOGY 67

The analysed sample, however, comprised 50 % plagioclase, 25 % quartz and 10 % microcline, and hence was a granodiorite (Nelson, 1999). It also contained biotite but no hornblende, and thus is similar to most of the exposed Carlindi complex and not very different from the previous two samples (Samples 95037, 153188). Small (<0.1 mm), disseminated zircon grains were generally included in feldspar, quartz or biotite. Separated zircons were dark yellow, brown or black, euhedral or irregular and up to 400 µm long. Internal zoning was common and many grains were metamict. Twenty-nine analyses were obtained from 29 zircons, of which 21 are concordant to near-concordant (Fig. 3.10). The discordance pattern is consistent with several episodes of Pb loss. A population of 22 zircons, including four discordant grains, defines a pooled 207Pb/206Pb age of 3470 ± 3 Ma (MSWD = 0.36). This is interpreted to represent the age of crystallisation, and differs from that defined by Nelson (1999) by excluding a sample which is >2 s younger than the population mean. There are three possible xenocrysts with ages between 3497 and 3485 Ma, which suggests either assimilation or incongruent melting of older crust. The remaining four grains do not define any significant populations, and so their ages apparently have no geological significance. The age of the Shilliman Well granodiorite is within error of the ~3468 Ma Carlindi unconformity granite (Sample 94058) and widespread ~3466 Ma Duffer volcanism (Pidgeon, 1978a; Thorpe et al., 1992a; McNaughton et al., 1993; Nelson, 1999; 2000). CHAPTER 3: GEOCHRONOLOGY 68

0.8 Intercept at

3470 ± 3 Ma 6.1 MSWD = 0.36 3550 n = 22 3450 0.7 3350 10.1 20.1 3250 25.1 12.1 3150 U

8 3050 3

2 19.1

/ 0.6

b 12.1 P 6 0 2

0.5

to origin 0.4 18 22 26 30 34 207Pb/235U

Figure 3.10: Concordia plot of sample 153190, Carlindi granodiorite, Shilliman Well (Van Kranendonk, 2000).

3.5.1.5 Xenolith host granite, Sample 100698 The following three samples were all collected from the same outcrop at Coonterunah Pool where there is a large number of xenoliths of varying composition hosted within granite (Fig. 3.1; 3.11; Plate 3.2B-C). Sample 100698 is the host to the xenoliths and was collected ~20 m from the nearest exposed xenolith and ~2 m from the nearest wispy biotite-chlorite partings of unknown origin and thin pegmatite veins (Fig. 3.11). The granite is massive and equigranular with medium- to coarse-grained (up to 4 cm) microcline (40 %), plagioclase (30 %), quartz (30 %) and chlorite -biotite (<5 %). It is typical of granite in the area. Plagioclase cores are typically turbid due to sericite alteration. Twenty-three analyses were obtained from 23 zircons of which 15 are concordant to near-concordant (Fig. 3.12). There are two distinct populations and 8 ungrouped grains. The discordance pattern is consistent with multiple early and one recent Pb-loss events. A population of 10 zircons defines a pooled 207Pb/206Pb age of 2925 ± 6 Ma (MSWD = 0.64). These are mainly near-concordant and there are no nearby outliers. A population of 4 zircons defines a pooled 207Pb/206Pb age of 3466 ± 9 Ma (MSWD = 0.66). There is one zircon with an age of 3524 ± 7 Ma, although it is CHAPTER 3: GEOCHRONOLOGY 69 slightly discordant. The remaining 7 zircons have ages between the two pooled ages, are mainly discordant and do not define significant groups. The ~2925 Ma population has very consistent Th/U ratios (0.31 to 0.50) compared to the other zircons, although they are not mutually exclusive. The older zircons are typically more equant than the youngest grains, and all grains have luminescent rims around relatively featureless cores (Plate 3.2D-E).

0.75 Intercept at 31.1 2925 ± 6 Ma 3400 41.1 MSWD = 0.64 3200 ?xenocrysts 0.65 n = 10 3466 ± 9 Ma 45.1 MSWD = 0.66 n = 4

37.1 3000 1.1

2800 0.55 14.1 U 8 3

2 2600 / 42.1 b P 6 0

2 0.45

39.1

to origin 0.35

37.1

to origin 0.25 10 14 18 22 26 30 207Pb/235U

Figure 3.12: Concordia plot of sample 100698, Carlindi granite, host to xenoliths, Coonterunah Pool. CHAPTER 3: GEOCHRONOLOGY 70

Figure 3.11: Outcrop map of ènclave locality at Coonterunah Pool, Carlindi Granitoid Complex. AMG co-ordinate in centre of outcrop (±50 m). SHRIMP and geochemical samples marked. CHAPTER 3: GEOCHRONOLOGY 71

With bimodal zircon populations in magmatic rocks it can be difficult to assign an interpretation to each population. Here, the ~3466 Ma zircons may be xenocrysts included in ~2925 Ma granite, or alternatively, the older population may represent the magmatic age of the granite with the younger population related to metamorphic recrystallisation. However, the granite must be younger than the included xenoliths (Sample 080698, 090698), and since ~3466 Ma zircons are relatively common within the xenoliths then it is not unreasonable to expect that the granite may have incorporated such xenocrysts along with the xenoliths. Moreover, the granite shows no evidence of post-magmatic metamorphism, and so there is little additional support for a ~2925 Ma metamorphic event. Although the CL and BSE images and Th-U systematics do not discriminate between the two populations, the Th/U ratio of the younger population is >0.1, which is inconsistent with known metamorphic zircons (Schaltegger et al., 1999; Vavra et al., 1999; Hartmann et al., 2000). Therefore, the younger population is interpreted to be the crystallisation of the granite ( 2925 ± 6 Ma) and the older zircons are xenocrysts. In general, the host granite was probably derived from a mixed source, though dominantly from ~3470 Ma granitoids. This interpretation is supported by the data from the xenoliths (Sample 080698, 090698). The ~2925 Ma age correlates with other Pilbara granites, including intrusions in the Shaw and Yule Granitoid Complexes, Satirist Granite, Caines Well Granite and west Pilbara granitoids (Arndt et al., 1991; Nelson, 1997, 1998, 2000; Smith, 1999).

3.5.1.6 Gneissic tonalite xenolith, Sample 080698 This sample was collected from a ~2 x 1 m exposure of gneissic xenolith and has sharp contacts with, but no cross-cutting veins from, the host granodiorite (Fig. 3.11). An adjacent amphibolite xenolith is separated by a <10 cm-wide coarse-grained leucocratic selvedge. The xenolith consists of <1 cm thick melanocratic and leucocratic bands, generally with the melanocratic bands being slightly thicker and hence more dominant. The banding has not been folded or sheared like in some nearby xenoliths, but a schistosity has been developed parallel to the gneissosity. The xenolith comprises relatively equigranular medium-grained (~1 mm) plagioclase (50 %), biotite (15 %), quartz (15 %), microcline (10 %), chlorite (10 %) and minor FeTi-oxides. Based on its high plagioclase content it is classified as a tonalite. CHAPTER 3: GEOCHRONOLOGY 72

0.75 Intercept at 36.1 3462 ± 2 Ma 8.1 3400 MSWD = 1.2 32b.1 n = 9 9.1 29.1 6.1 0.65 3200 12.1 19.1

37.1 3000 7.1

30.1 U

8 39.1 3 2 2800 / 0.55

b 18.1

P 38.1 6 0 2 25.1

0.45 28.1 31.1

22.2

to origin 0.35 12 16 20 24 28 32 207Pb/235U

Figure 3.13: Concordia plot of sample 080698, gneissic tonalite xenolith, Carlindi complex, Coonterunah Pool.

Twenty-eight analyses were obtained from 28 zircons of which 15 are concordant to near-concordant (Fig. 3.13). A population of 9 near-concordant grains defines a pooled 207Pb/206Pb age of 3462 ± 2 Ma (MSWD = 1.2). Four grains have significantly older ages, with the oldest 3512 ± 7 Ma. There are four significantly younger near-concordant grains (3449, 3391, 3300, 3237 Ma) but they do not define groups. The tonalite is interpreted to have a crystallisation age of 3462 ± 2 Ma and have been derived from older material. It was then intensely modified after crystallisation and was incorporated in the ~2925 Ma host granodiorite. This complex history of deformation and metamorphism is recorded by the zircons and reflected by the gneissic fabric. An alternative is that the tonalite crystallised before ~3462 Ma and that these zircons reflect intense metamorphism and deformation, too. However, all of the older zircons have Th/U > 0.1 and so were probably magmatic. In contrast, a number of younger near-concordant to discordant zircons have Th/U < 0.1, consistent with their metamorphic interpretation. CHAPTER 3: GEOCHRONOLOGY 73

3.5.1.7 Gneissic granite xenolith, sample 090698 This sample was collected from a gneissic granite xenolith ~1 m east of sample 080698 (Fig. 3.11). It is similar to the sample 080698, but has a greater felsic component and more variable band widths. The xenolith comprises relatively equigranular coarse-grained (~2 mm) plagioclase (35 %), biotite (5 %), quartz (25 %), microcline (30 %) and chlorite (<5 %). It does not contain veins of the host granodiorite.

0.8 Intercept at 3476 ± 4 Ma MSWD = 1.06 n = 3 3.1

11.1 6.1 3500 0.7 Intercept at 30.1 3234 ± 2 Ma 3300 12.1 MSWD = 1.8 n = 7 23.1 31.1 29.1 14.1 20.1 7.1 33.1 3100 U 8 3 2 0.6 28.2 /

b 2900 P 6 0 2

0.5 25.1

13.1

to origin to origin 0.4 14 18 22 26 30 34 207Pb/235U

Figure 3.14: Concordia plot of sample 090698, gneissic granite xenolith, Carlindi complex, Coonterunah Pool.

Twenty-five analyses were obtained from 25 zircons of which 13 are concordant to near-concordant (Fig. 3.14). A population of 7 zircons defines a pooled 207Pb/206Pb age of 3234 ± 2 Ma (MSWD = 1.8), and 3 zircons define an age of 3476 ± 4 Ma (MSWD = 1.06). The remaining data are between these ages, including 8 slightly discordant grains between ~3467 and ~3408 Ma. These can be manipulated into significant populations, but with 3 or less grains. Since 090698 is a gneissic xenolith and has clearly been intensely modified after crystallisation then it is likely that the younger grains were associated with metamorphism. Indeed, six grains have Th/U < 0.1, consistent with such an interpretation. However, the youngest population does not CHAPTER 3: GEOCHRONOLOGY 74 include any of these six grains, and so this population may not be metamorphic. Moreover, it is unlikely that the youngest population represents the pre-gneissic crystallisation age. Instead, the youngest population may be related to partial melting of the granite to form gneissic banding. If so, the granite may have crystallised at ~3476 Ma, been periodically metamorphosed for ~200 myr and then partly remelted at ~3234 Ma.

3.5.2 Interpretation of the Carlindi granitoids The granitoid samples highlight the inherent problems in corr elating igneous events in intrusive complexes where outcrop is discontinuous but the rocks are lithologically and geochemically very similar. For example, prior to obtaining a crystallisation age, the host granite (100698) was generally thought to be equivalent to other Carlindi granites (95058, 153190), and yet it is now interpreted to have crystallised ~550 myr later. Consequently, the correlation of outcrops based solely on field criteria is misleading, and maps based on such data should be treated with scepticism (including those herein). The main outcome of the Carlindi dating is that major intrusions were emplaced between ~3485 and ~3465 Ma. This correlates chronologically with many intrusions in the northern Shaw Granitoid Complex (Bickle et al., 1983; 1993; McNaughton et al., 1988; Zegers, 1996; Nelson, 2000), and other intrusions in the Mount Edgar and Muccan complexes (Nelson, 1998). It also correlates with felsic volcanism in the Marble Bar and Coongan Belts (McNaughton et al., 1993; Nelson, 1999; 2000). Therefore, many of the Carlindi granitoids were part of a major regional felsic igneous event. Since this represents the oldest widespread granitic event preserved in the Pilbara, these rocks provide critical constraints on crustal evolution. The ~2925 Ma granodiorite at Coonterunah Pool represents a much younger major intrusive event, most clearly manifest in granitoid emplacement in the west Pilbara. Xenocrystic zircons in many of the older Carlindi granitoids indicate that they were probably derived from older crust. Their elemental and isotopic compositions support this supposition (Chapter 5), as does their compositional similarity with the ~2925 Ma granite, which was clearly derived from older crust. CHAPTER 3: GEOCHRONOLOGY 75

3.6 WARRAWOONA GROUP 3.6.1 Results In the Pilgangoora Belt, there are no felsic volcanic units in the Warrawoona Group, and, as a result, the depositional age is poorly constrained. A maximum age is provided by detrital zircons from a sandstone unit immediately above the Coonterunah- Carlindi unconformity (O’Brien, 1999), although only three grains have ages younger than the granite (94058) cut by the unconformity. Detrital zircons also provide robust evidence for the provenance of the sandstone (Chapter 2). To further constrain the depositional age, a sample was collected from stratigraphically higher in the North Pole succession (Buick et al., 1995). Numerous lines of evidence support the correlation between the Pilgangoora and North Pole successions, and so the use of these data is presumed valid here.

3.6.1.1 Strelley Pool Chert Sandstone, Sample 98OB5002 This sample was collected from the basal sandstone unit of the Strelley Pool Chert, ~230 cm above the Coonterunah-Carlindi unconformity, in the central Pilgangoora domain (Fig. 3.1; O’Brien, 1999). It consists of silicified, well-rounded, well-sorted, medium-grained quartz with prominent heavy mineral laminations. The sandstone contains abundant euhedral to rounded zircon grains that are generally concentrated in the laminations, but, in place, they remain included in quartz grains (Plate 3.2E). Zircons may be up to 250 µm long, but are generally about 100 µm. Twenty-four zircons were analysed of which 20 are concordant to near- concordant (Fig. 3.15). The discordance pattern is consistent with recent U-Pb disturbance. Twenty-one grains have 207Pb/206Pb ages between 3520 to 3473 Ma, and hence, are older than the granite (94058) cut by the unconformity. As a result, these grains do not improve the previous constraint of the age of sandstone deposition. Three zircons are significantly younger, including a near-concordant grain with an age of 3390 ± 3 Ma. These grains may reflect the depositional age. CHAPTER 3: GEOCHRONOLOGY 76

0.8

3550 3450

0.7 3350

3250 3390 Ma

3150 3411 Ma

U 3050 3473 Ma 8 3 2

/ 0.6

b 3379 Ma P 6 0 2

0.5

0.4 18 22 26 30 34 207Pb/235U

Figure 3.15: Concordia plot of sample 98OB5002, Strelley Pool Chert sandstone, detrital zircons, Warrawoona Group.

In the older population, seven zircons define a pooled age of 3516 ± 6 Ma, which is within error of the underlying Coonterunah felsic volcanics and from which they are likely to have been derived. Five zircons define a pooled age of 3498 ± 7 Ma and are within error of a zircon population obtained from granitic xenoliths hosted in Coonterunah basalt (520798). These xenolithic zircons were apparently recrystallised during post-Coonterunah/pre-Warrawoona metamorphism. Eight of the remaining zircons have ages between these two populations and probably represent varying degrees of xenocryst recrystallisation. Indeed, this spread of post-Coonterunah ages provides further evidence that metamorphic recrystallisation has not produced a single age population. As a result, the geological significance of the ~3499 Ma age obtained from the granite xenoliths is doubtful. The overlap of the detrital zircons with the two Coonterunah populations is clearly demonstrated in cumulative probability plots for each rock type, where a Gaussian distribution of each point and error have been included (Fig. 3.16). Hence, it is reasonable to conclude that the twenty older zircons were probably derived from within the Coonterunah succession. One discordant grain is within error of the crystallisation age of the Carlindi granite (94058, 153190). The U-Th abundances of the detrital zircons provide no further discriminating criteria. CHAPTER 3: GEOCHRONOLOGY 77

Detrital zircons can be used to constrain the maximum possible age of a clastic unit by attributing this to the youngest concordant grain. Accordingly, the SPC sandstone would be younger than the single concordant ~3390 Ma zircon. A discordant grain also has an age within 2 s of this and together they have a pooled age of 3388 ± 10 Ma (MSWD = 0.94). The other young grain does not fit this population. The next youngest concordant zircon is ~100 myr older, and so the above constraint is limited to two grains. Moreover, this age does not correspond with any known intrusive, volcanic or metamorphic Pilbara event (Fig. 3.19; Nelson et al., 1999; Zegers, 1996). Consequently, it is questionable whether these grains represent the age of an eroded source, and thus constrain deposition. Instead, this age may reflect analyses from recrystallised metamict grains, and if so, has no geological importance. However, recent ages obtained from the North Pole Dome (Nelson, 2000) and the Kelly Belt (Barley et al., 1998) show that the felsic volcanics beneath the SPC are significantly younger than previously thought (Thorpe et al., 1992a). In this case, these young detrital grains may indeed represent an age of geological importance. This is discussed below. Clearly, there were other available zircon sources, including Carlindi granitoids, that were either not found or very rare in the sandstone sample. For example, the SHRIMP sample was collected <7 km from where the Carlindi microgranite was visibly eroded prior to SPC deposition, but only one discordant detrital zircon within error of its age was detected. Although the detrital population may be artificially skewed by too few zircons being analysed from only a single locality, it does suggest that zircon was preferentially derived from the local Coonterunah succession. Furthermore, if the SPC is significantly younger than previously proposed, then the paucity of non-Coonterunah grains is even more remarkable. Such local sourcing may imply that the SPC sandstone was deposited on a sediment-starved coastline. If so, the absence of pre-Coonterunah zircons in this sample does not then preclude the existence of older continental basement. CHAPTER 3: GEOCHRONOLOGY 78

3498 SPC sandstone

3504 n = 20

Granitic xenoliths 3517 y t i

l n = 17 i

b 3515 a b

o Coonterunah r P felsic volcanics e v

i n = 47 t a l e R

3470 3490 3510 3530 3550 age (Ma)

Figure 3.16: Relative cumulative probability plot showing the age overlap of detrital zircons from the SPC sandstone with both the graniticxenoliths and Coonterunah felsic volcanics. Samples are plotted as Gaussian distributions with 1s errors.

3.6.1.2 North Pole Dome (#9 chert), Sample 94001 In order to constrain the duration of the Coonterunah-Warrawoona unconformity, this sample was collected from a thin Warrawoona chert horizon in the North Pole Dome (Buick et al., 1995), where the succession is remarkably similar to the Warrawoona succession in the Pilgangoora Belt. Aside from the well-documented correlation of the Strelley Pool Chert (Lowe, 1983; Buick & Barnes, 1984, DiMarco & Lowe, 1989a, b), the gross stratigraphy is also repeated (see Chapter 2). For example, the alternating order of magnesian and tholeiitic basalts is the same and the basalt units are separated by similar tuffaceous, volcaniclastic chert horizons. Trace element compositions of basalts also support this correlation (see Chapter 4), although the thickness of the basalt units varies between the greenstone domains, as they also do along strike in the Pilgangoora Belt, and stratigraphically higher units have been preserved in the North Pole Dome but not in the Pilgangoora Belt. The sample was collected from the upper part of the North Pole succession, ~3 km stratigraphically above the Strelley Pool Chert (Buick et al., 1995; Van Kranendonk, 2000), and thus was deposited after the last exposed Warrawoona unit in the Pilgangoora Belt. CHAPTER 3: GEOCHRONOLOGY 79

0.85

3415 Ma

3650 0.75 3463 Ma

3550 3450

3350 U

8 3250 3 2

/ 0.65

b 3150 P 6 0 2

0.55

0.45 20 24 28 32 36 207Pb/235U

Figure 3.17: Concordia plot of sample 94001, pumiceous arenite, North Pole Dome (Buick et al., 1995).

The sample is a rippled pumiceous arenite with minor clastic qu artz, muscovite and chert. Many zircons are rounded and clearly detrital, though the sample contains some juvenile felsic volcanic fragments. Nineteen zircons were analysed of which 15 are concordant to near-concordant (Fig. 3.17). The oldest zircon is highly discordant, and thus possibly suspect, though its 207Pb/206Pb age of 3547 ± 6 Ma suggests a pre- Coonterunah source. The 15 concordant to near-concordant zircons span a 207Pb/206Pb age range between 3520 and 3463 Ma, and do not define significant groupings. Including discordant analyses there are six possible age groups, three with only one representative. The oldest population includes five zircons and has a pooled 207Pb/206Pb age of 3514 ± 3 Ma (MSWD = 1.14), which is within error of the Coonterunah felsic volcanics, their probable source. The next nine oldest zircons span between 3511 and 3498 Ma, similar to apparently recrystallised zircons in xenoliths hosted in Coonterunah basalt. Two zircons have ages of ~3486 Ma and may have been derived from Carlindi equivalents. One angular zircon fragment has an age of 3463 ± 7 Ma, within error of the ~3466 Ma Duffer Formation (McNaughton et al., 1993), and was interpreted to record the eruption age of the juvenile felsic material (Buick et al., 1995). A superconcordant CHAPTER 3: GEOCHRONOLOGY 80

~3415 Ma zircon has not been previously regarded to be of any significance, but this may not be the case (see below).

3.6.2 Interpretation Previous studies by McNaughton et al. (1993) have shown that zircons within Pilbara metabasalts have generally been modified by post-volcanic thermal events and do not provide reliable eruption ages. Consequently, zircons from felsic volcanics are preferred for determining depositional ages. In the Pilgangoora Belt, there are no Warrawoona felsic volcanic units, and hence, the age of the succession cannot be directly dated with confidence. As a result, constraining Warrawoona ages in the Pilgangoora Belt relies on dating detrital zircons and correlations with other belts. The SPC has been reported in the Pilgangoora, North Pole, North Shaw, Coongan and Kelly Belts and is regarded as a regional stratigraphic marker that usually overlies the felsic volcanic Panorama Formation (Lowe, 1983; Buick & Barnes, 1984, DiMarco & Lowe, 1989a, b). If this correlation holds, then the age of the Panorama Formation provides an important constraint on the depositional age of the Warrawoona Group in the Pilgangoora Belt. Such a correlation has already been discussed for the upper part of the North Pole succession (see 3.6.1.2). In the North Pole Dome, the Panorama Formation underlies the SPC and was previously determined to have an eruption age of 3458 ± 2 Ma (conventional whole- grain analyses, Thorpe et al., 1992a). The SPC and overlying basalts were considered to have been deposited soon afterwards. Intrusion of the North Pole Monzogranite (née Adamellite) at 3459 ± 18 Ma and the Miralga Creek porphyry at 3449 ± 2 Ma provided consistent constraints (Thorpe et al., 1992a). However, recent microprobe dating from a felsic unit in the Panorama Formation has determined a crystallisation age of 3434 ± 5 Ma from a robust population of concordant zircons (Nelson, 2000; Van Kranendonk, 2000). This sample is described as a silicified aphanitic chert-sericite rock with slightly larger quartz grains (?microphenocrysts) and has been interpreted to be tuffaceous (Nelson, 2000). The petrographic description, however, does not include any unequivocal volcanic textures and so the unit could be intrusive, thus not dating the Panorama Formation. An intrusive interpretation is consistent with an age of 3431 ± 3 Ma obtained from a porphyritic dacite dyke that cuts the lower part of the North Pole succession (R. Buick, unpubl. data). Hence, the Panorama Formation may be ~3458 myr CHAPTER 3: GEOCHRONOLOGY 81 old and the ~3434 Ma felsic units are later intrusions or the dyke may be the intrusive part of a younger Panorama Formation. This later interpretation is supported by a ~3434 Ma population of concordant zircons from the felsic volcanics beneath the SPC in the Kelly Belt (Barley et al., 1998; Nelson, 2000). Such consistency calls into question the validity of the conventional whole-grain analyses. For instance, the North Pole Monzogranite age is defined by a chord of highly discordant analyses with a lower concordia intercept at 774 Ma. There is no explanation of this lower age and it has not been noted for other samples. Hence, it is not robust, as reflected by its large error. In contrast, the Panorama volcanic and Miralga Creek Porphyry ages are defined by more concordant zircons, and so are more robust. Indeed, the Miralga Creek Porphyry has recently been dated again using conventional zircon methods with only slightly younger results (Amelin et al., 2000), and a pooled age of ~3442 Ma is obtained from all the near-concordant zircons. Nevertheless, in light of the new microprobe data and improved understanding of zircon inheritance and recrystallisation, it is possible that the conventional whole-grain analyses have averaged the age of zoned grains, and thus their ages do not represent geological events. Taking the younger Panorama age at face value implies that the North Pole SPC must be younger than ~3434 Ma, and so, the youngest detrital grain in the overlying pumiceous arenite (sample 94001) may indeed reflect the maximum age of deposition at ~3415 Ma (cf. Buick et al., 1995). Interestingly, there is a population of ~3414 Ma zircons in the Kelly Belt felsic volcanics (Barley et al., 1998; Nelson, 2000). Extrapolating to the Pilgangoora Belt indicates that the SPC sandstone is also younger than ~3434 Ma, possibly <3414 Ma. Indeed, the youngest detrital zircons from the sandstone suggest that the SPC may be as young as ~3390 Ma. If the SPC does represent a regional marker horizon that was deposited <3415 Ma, then there must be disconformities beneath the SPC in all of the Pilbara greenstone belts. To date, a disconformity has only been recognised in the Pilgangoora Belt, perhaps because here, at least in part, the contact is angular.

3.7 DISCUSSION The use of the SHRIMP to determine zircon U-Pb contents is well-established, and there is no doubt that the analyses here were accurate. However, defining significant populations and attributing geological events to these populations is less certain. Some CHAPTER 3: GEOCHRONOLOGY 82 samples have well-defined populations with few, if any, outliers, and are readily interpreted. Other samples, however, have many discrete populations and have been interpreted according to their geological context. For example, the gneissic xenoliths and the host granite have similar bimodal populations, but contrasting interpretations. All of the samples herein have interpretations that are consistent with their context.

3.7.1 Geological history The main aim of the zircon dating was to determine the temporal relationship between the different units in the field area. This has been achieved, although new problems have arisen from the data. The geological history is depicted in Figure 3.18 and is summarised as: i) deposition of the basalt-dominated Coonterunah Group at ~3517 Ma, ii) metamorphism of the Coonterunah Group probably between ~3504 and 3480 Ma, iii) intrusion of Carlindi granitoids between ~3485 and 3465 Ma, iv) uplift and erosion prior to SPC deposition, v) deposition of basalt-dominated Warrawoona Group after ~3458 Ma, ~3434 Ma, if recent microprobe sample is volcanic, or possibly <3415 Ma or even <3388 Ma, if detrital grains are considered at face value, vi) intrusion of xenolith-host granite at ~2925 Ma. The presence of continental basement to the Coonterunah succession is discussed later and is a major proposition of this thesis. The 3480-3465 Ma Carlindi intrusive events are considered to have created the angular discordance, as opposed to invoking a separate deformation event. Folds within Coonterunah cherts are interpreted to have formed due to bedding-parallel slip within the chert, as the chert itself has not been folded. The importance of the ~2925 Ma intrusions for final doming of the Carlindi complex is unclear. A graphical representation of the reported 3300 to 3520 Ma Pilbara events is presented in Figure 3.19 and relevant dates are compiled in Table 3.3. The whole-grain conventional ages have not been included since they are possibly suspect (see 3.6.2; Thorpe et al., 1992a; Amelin et al., 2000). Each relevant greenstone belt is presented as a column and includes its immediately adjacent granitoid complex. For the Shaw Complex, most of the granitoids have been included with the North Shaw Belt, but are CHAPTER 3: GEOCHRONOLOGY 83 also applicable to Coongan Belt. In the Kelly Belt, two samples have provided two populations each (Table 3.3).

A. Deposition of the Coonterunah Group B. Intrusion of Carlindi granitoids at ~3484 at ~3517 Ma on older continental crust and ~3467 Ma cause doming and perhaps metamorphism of Coonterunah succession basalt VV V V V VV V V V V V V V V V ~3517 Ma chert V VV V V V VV V V V V V felsic volcanic ~3517 Ma V + V V ~3468 Ma basalt V x V V VV V V V V x x + + V V VV V V V VV V V V ~3484 Ma x x x + + + V x

older continental basement x x x x + + x

V V V V V V V basaltV s

V V ~34+68 Ma V V a SPC M V x x 7 x V 1 + + V V ~34+68 Ma 5 ~3484 Ma V a V 3 V x M V x ~ x x x 7 x + + + 1 + V 5 ~3484 Ma + x V 3 V x x ~ x x x x + + x x + + + + V x x x x x x x x x x x x x + + C. Uplift and erosion of Coonterunah- D. Deposition of SPC and basalts at Carlindi terrain before 3434 Ma ~3434 Ma on the Coonterunah-Carlindi unconformity

Figure 3.18: Schematic diagram showing the early evolution of the Pilgangoora Belt. Granitoid intrusions at ~2925 Ma and final tilting of terrain not shown.

The Coonterunah Group has only been identified in the Pilgangoora Belt and does not correlate with any known granitoids. It represents the oldest supracrustal rocks in the Pilbara, and thus provides crucial information about the crustal evolution. Xenoliths, xenocrysts and detrital grains aside, the Coonterunah Group is the oldest geological entity in the Pilbara. A hornblende 40Ar/39Ar plateau age of 3522 ± 13 Ma was obtained from fine-grained metabasalt in the North Shaw Belt (Zegers, 1996), but its relevance to magmatism has not been established. Notably, the hornblende contained excess argon and a duplicate sample produced a significantly younger age. However, it may represent the magmatic age, as favoured by Zegers (1996), and if so, the North Shaw Belt also contains Coonterunah basalts. Further evidence is required to verify this claim. The age of regional metamorphism is poorly constraine d, but is probably reflected by zircon recrystallisation in granitic xenoliths hosted in Coonterunah basalt. It may be related to intrusion of individual Carlindi granitoids, although a prolonged CHAPTER 3: GEOCHRONOLOGY 84 thermal anomaly may have existed between ~3505 and 3480 Ma, and thus metamorphism could have started prior to the oldest known intrusion. The Carlindi granitoids were generally coeval with similar intrusions in the Shaw, Mount Edgar and Muccan complexes. They were also contemporary with felsic and mafic volcanism in the Coongan and Marble Bar Belts. An equivalent age has also been determined in the North Shaw Belt (sample 142964), but it is unclear whether the sample is from a felsic volcanic or intrusive (Nelson, 1999; Van Kranendonk, 2000). The Carlindi granitoids thus represent part of a regionally significant magmatic event. Moreover, similar aged xenoliths and xenocrysts in younger granitoids suggest that these older granitoids were incorporated into younger melts, and thus were even more widespread. Xenocrystic zircons in the older Carlindi granitoids are also consistent with them incorporating or even melting slightly older crust. The age of the SPC and overlying basalts may be significantly younger than previous estimates and thus the timing of emergence is between 30 to 70 million years later than proposed by Buick et al. (1995). Hence, uplift and erosion of the Coonterunah-Carlindi terrain happened more than 30 myr after Carlindi granite intrusion, and not as rapidly as previously supposed. This terrain still represents the oldest known emergent landmass. The ~2925 Ma xenolith-host granite at Coonterunah Pool represents the youngest intrusive event in the area and was contemporary with granitoids in the Shaw and Yule Granitoid Complexes, Satirist Granite, Caines Well Granite and west Pilbara granitoids (Arndt et al., 1991; Nelson, 1997, 1998, 2000; Smith, 1999) . These ~2925 Ma intrusions represent some of the last known granitoids emplaced before deposition of the ~2767 Ma epicratonic basalts and sediments of the Fortescue Group (Blake, 1984, 1993; Arndt et al., 1991), and thus were probably crucial to stabilisation of the Pilbara Craton. The Carlindi complex therefore records some of the oldest and youngest intrusive events. CHAPTER 3: GEOCHRONOLOGY 85

Figure 3.19: SHRIMP zircon U-Pb ages for Pilbara magmatism between 3520 and 3300 Ma. Details are presented in Table 3.3. CHAPTER 3: GEOCHRONOLOGY 86

Table 3.3: Details of Pilbara SHRIMP zircon ages between 3520 and 3000 Ma. Bold refers to Figure 3.19. (* data recalculated from published age) Ref.: A, B– McNaughton et al., 1988; 1993; C – Buick et al., 1995; D – Zegers, 1996; E, F, G, H – Nelson, 1996; 1998; 1999; 2000; I – Barley et al., 1998; J – Barley & Pickard, 1999; K – herein, L – Buick unpubl. sample # description age ref sample # description age ref 1 70649 volcanic dacite, 3518 ± 4 K 28 K1 hyalo. rhyodacite, 3434 ± I Coonterunah Panorama 1 11* 2 70601 hyalo. rhyolite, 3515 ± 3 C 29 148502 tuff. rhyolite, 3433 ± 2 H Coonterunah Panorama 1 3 520798 Granitic enclaves, 3499 ± 4 K 30 K1 as above, 3413 ± 4* I Coont. basalt Panorama 2 4 153188 Granodiorite, 3484 ± 4 G 31 148502 as above, 3412 ± 10 H Wilson Well Panorama 2 5 090898 Gneissic xenolith, 3476 ± 4 K 32 uwa98074 porph. rhyolite, 3325 ± 4 B Coonterunah Pool Wyman Formation 6 95037 Granodiorite, 3479 ± 11 K 33 uwa98075 amyg. dacite flow, 3324 ± 4 B unconformity Wyman Formation 7 153190 Granodiorite, 3470 ± 3 G 34 Copper Hills 3321 ± 4 J Shilliman Well porphyry 8 94058 Granite, 3468 ± 4 C 35 Mondana Suite 3317 ± 2 J unconformity 9 100698 old population, 3466 ± 9 K 36 Boobina Porphyry 3315 ± 4 J xenolith host 10 080698 Tonalite xenolith, 3462 ± 2 K 37 Carbana Pool 3313 ± 9 J Coonterunah Pool Suite 11 98OB5002 SPC sandstone, 3388 ± 10 K 38 142978 Monzogranite, 3307 ± 4 H youngest detrital Carbana Pool 12 142952 Felsic volcanic, 3434 ± 5 H 39 MP1 Grey dacite, 3430 ± 3 I Panorama Form. Panorama 1 13 70969 Dacitic porphyry, 3431 ± 3 L 40 Gobbos stock 3313 ± 4 J cross-cuts > #1 14 94001 #9 tuff. arenite, 3415 ± C 41 148498 tuff. chert, 3477 ± 2 H youngest detrital McPhee Form. 13* 15 91084 leucogranite 3493 ± 4 A 42 148509 tuff. andesite, 3468 ± 2 H Duffer Formation 16 142962 Tonalite, 3469 ± 2 H 43 148500 Lapilli tuff, 3467 ± 4* G Chocolate Hill Duffer Formation 17 91085 Tonalite, 3467 ± 6 A 44 uwa98076 Felsic pyroclast., 3466 ± 4 B Coolyia Creek Duffer Formation 18 142964 Felsic gneiss, 3466 ± 3 H 45 142865 Alkali granite 3466 ± 2 F Keep-it-Dark 19 T94/222 Diorite 3462 ± 2 D 46 142828 Granodior. gneiss, 3470 ± 4 F Fred W., Muccan 20 T94/221 Grey gneiss 3451 ± 1 D 47 124755 Biotite granodior., 3443 ± 6 E Sunrise, Muccan 21 uwa98053 Granodiorite, 3431 ± 4 B 48 143807 Tonalite, Kennedy 3438 ± 4 F South Daltons Plu. Gap, Muccan 22 142966 Granodiorite, 3430 ± 4 H 49 142170 Fol. monzogranite, 3421 ± 2 G Unices Well Kangan, Muccan 23 142968 Biotite tonalite, 3425 ± 4 H 50 143803 Granodiorite, Don 3313 ± 3 F Tambourah Well Well, Muccan 24 142975 Felsic volcanic, 3474 ± 7 H 51 143809 Granodiorite, 3313 ± 6 F Duffer Formation Ngarrin, Muccan 25 T94/227 Mylonite granite, 3469 ± 3 D 52 142806 Monzogranite, 3303 ± 2 F Split Rock Shear Z Yundinna, CHAPTER 3: GEOCHRONOLOGY 87

Muccan

26 142976 Granophyre, 3469 ± 3 H 53 143871 Granodiorite, 3303 ± 5 G Mt Ada Basalt Warrawagine 27 142878 Foliated biotite 3445 ± 3 F monzogranite CHAPTER 3: GEOCHRONOLOGY 88

3.7.2 Proposed lithostratigraphy Geochronological evidence has led to the gradual unravelling of the lithostratigraphy proposed by Hickman (1983). Most of these changes have involved deconstruction of a straightforward regional-scale, layered model. For instance, the ~3517 Ma Coonterunah Group and ~3240 Ma Sulphur Springs Group have been discovered in rocks previously assigned to the ~3466 Ma Warrawoona Group. More recent results, including those presented here, necessitate further deconstruction. It seems inescapable that the current War rawoona Group comprises at least two discrete entities; one containing the ~3466 Ma Duffer and the other £3458 Ma Panorama volcanics. The present Talga Talga and Salgash Subgroups already suggest such division. However, if the ~3434 Ma age holds for the Panorama Formation, then the >30 myr age difference between the subgroups implies that some previously proposed correlations are untenable. A revised stratigraphy is proposed using this new Panorama age, current stratigraphic names have been retained (Fig. 3.20).

Wyman Formation ~3325 Ma Euro Basalt: magnesian to tholeiitic basalt & gabbro, commonly pillowed, silicified mafic volcanic horizons Strelley Pool Chert: silicified quartz-rich sandstone, laminated p h u carbonates & mafic volcanics, possibly disconformable s o a r ? ? ? g g l b a

u Panorama Formation: felsic volcanics and volcaniclastics ~3434Ma S P S

U Apex Basalt: magnesian to tholeiitic basalt & gabbro, commonly O

R pillowed, silicified mafic volcanic horizons G

A Dresser Formation: magnesian basalt, commonly pillowed, capped N by the North Pole chert-barite unit O O W

A tholeiitic and magnesian basalts, commonly pillowed, includes R

R red/white laminated Marble Bar chert A W a

p Duffer Formation: dacite to rhyolite volcanic and volcaniclastics ~3466 Ma g l u a o r T Mt Ada Basalt: tholeiitic to magnesian basalt & gabbro, typically

g a b

g massive, some pillows u l a S

T McPhee Formation: tholeiitic to magnesian basalt & gabbro, some ultramafic units North Star Basalt: tholeiitic to magnesian basalt & gabbro COONTERUNAH GROUP ~3517 Ma

Figure 3.20: Proposed lithostratigraphy the Warrawoona Group. See text for further explanation.

The Talga Talga Subgroup is re-interpreted to include all the units in the Marble Bar Belt, and thus contains the Duffer Formation. The Talga Talga Subgroup also forms CHAPTER 3: GEOCHRONOLOGY 89 the lower part of the Coongan Belt, but here the Strelley Pool Chert (Salgash Subgroup) overlies the Duffer Formation. Hence, a disconformity is inferred beneath the SPC in this area. It is possible that the basal part of the North Shaw Belt also contains Talga Talga units, including Duffer volcanics (Van Kranendonk, 2000). If so, they are also disconformably overlain by the SPC. Both Duffer and Panorama volcanics have been recently mapped here, but this has not yet been confirmed by geochronology (Van Kranendonk, 2000). It has also been proposed that this belt contains part of the Coonterunah succession (Zegers, 1996). The Salgash Subgroup is re-interpreted to include only those units in the North Pole Dome, and is little different from its original configuration but must include the Dresser Formation since this contains the North Pole chert-barite unit (previously in Talga Talga Subgroup, Hickman, 1983; Van Kranendonk, 2000). It is presently unclear whether the SPC and Euro Basalt are significantly younger and represent a further major subdivision (?unconformity in Fig. 3.20). The Dresser Formation and Apex Basalt have only been identified in the North Pole Dome. The upper part of the Salgash Subgroup is present in the Pilgangoora, North Shaw, Coongan and Kelly Belts and the McPhee Dome. The significance of younger concordant zircons in the Kelly Belt is unclear, but they may reflect a discrete event at ~3414 Ma (Fig. 3.19, Panorama 2). The new subdivision implies that the SPC represents a true marker horizon, which is supported by present evidence. However, a previous correlation with the Marble Bar Chert is now untenable (Lowe, 1983; DiMarco & Lowe, 1989a, b). Major disconformities, like that already identified in the Pilgangoora Belt, are inferred in the Coongan and North Shaw Belts. The proposed stratigraphy indicates that early Pilbara ba salt-dominated volcanism was punctuated and there does not appear to be a spatial-temporal preference to the volcanism. In other words, the basal greenstone successions do not get younger towards the margins of the craton, indicating that the east Pilbara is not composed of accreted terrains.

3.7.3 Pre-Coonterunah zircons A major proposition in this thesis is that pre-existing continental crust contributed to the earliest preserved rocks of the Pilbara, and even formed the basement to the Coonterunah greenstones. The samples analysed during this study contain no CHAPTER 3: GEOCHRONOLOGY 90 significantly older zircons, and so provide no geochronologic evidence of this precursor crust. Indeed, the oldest zircons in the Carlindi granitoids (3532 ± 9 Ma; sample 70601) and Coonterunah felsic volcanics (3527 ± 6 Ma; sample 90649) are both within 2 s of the Coonterunah magmatic age. Granitic xenoliths within Coonterunah basalt also yielded no old ages, although their zircons clearly had complicated histories (520798). There were even no old detrital grains in the SPC sandstone. There are only five Pilbara samples that have so far yielded pre-Coonterunah zircon xenocrysts (Table 3.4), despite zircon xenocrysts being common. These intrusives are widely distributed and are not a local feature. Moreover, two xenoliths have interpreted pre-Coonterunah crystallisation ages (Table 3.4). The 6 Mile Well gneissic tonalite xenolith is the best example of older crust and its broad range of >3550 Ma zircons suggests a prolonged, complex ancient history.

Table 3.4: Age and locality of pre-Coonterunah zircon xenocrysts. Last two samples are xenoliths ( A – Thorpe et al., 1992a; B, C, D – Nelson, 1998, 1999, 2000; E – McNaughton et al., 1988). Sample host age Locality Granitoid Xenocryst # (Ma) Complex age (Ma) 100507A <3458 Panorama Formation North Pole 3724 ± 1 143828B 3470 Fred Well Muccan 3538 ± 3 3574 ± 3 142884B 2933 Mount Webber Yule 3600 ± 3 142936D 2937 Yandeyarra Pool Yule 3552 ± 5 142967D 2929 Unices Well Shaw 3580 ± 5 142975D 3474 Duffer Formation, Coongan 3539 ± 9 160498D 2945 Geemas Well (intrudes Mallina) 3664 ± 4 uwa?E gabbroic ènclave, South Daltons Shaw 3578 ± 4 142879C 3244 banded tonalite gneiss, 6 Mile Warrawagine 3570 - 3658 Well CHAPTER 3: GEOCHRONOLOGY 91

PANORAMA 1 30 PANORAMA 2 sample numbers 142867, 142951, 142836, 143996, 143994, 143995, 25 142188, 142942, 142943 (Nelson, 1999; 2000) R E F F

20 U D r e b N A

m 15 M Y u H W A N N U R E

10 T N O O C 5

0 3250 3350 3450 3550 3650 3750 3850 age (M a)

Figure 3.21: Age distribution of detrital zircons from nine Pilbara clastic units. Cumulative probability plot (Gaussian with 1s error) and histogram of mean age counts of 223 grains. Ten myr bins.

A more extensive dataset of older grains has been recorded in clastic sedimentary samples. Nine samples contain >3550 Ma grains; six from the northeast Pilbara (Shay Gap, Marble Bar, Lalla Rookh belts; Williams, 1999) and three from the Mallina Basin in the central Pilbara (Fig. 1.3; Smithies, 1998, 1999). The samples were collected from <3200 Ma clastic units and probably record a relatively unbiased sample of ancient zircons. However, there may still be a tendency towards favouring local provenance, as suspected for the SPC sandstone in the Pilgangoora Belt. That aside, there does not appear to be any notable differences of age distribution between spatially separated samples. A cumulative probability plot of >3300 Ma zircons from these nine samples shows three main peaks at ~3305, 3430 and 3600 Ma (Fig. 3.21). The younger peaks correlate with the Wyman and Panorama-1 events. In contrast, there are few detrital zircons from the Coonterunah or Duffer events. The reason for this is unclear, but may reflect restricted basin development or variable uplift during the late Archaean. The oldest peak provides an estimate of the age of pre-Coonterunah Pilbara crust. Although the detrital zircons may have been sourced from a continental block that was adjacent to the Pilbara in the late Archaean, the presence of xenoliths and xenocrysts in magmas from a wide age spectrum indicates that crust of this age was integral for building the CHAPTER 3: GEOCHRONOLOGY 92

Pilbara. Therefore, there is abundant but cryptic zircon evidence for pre-Coonterunah crust which from the broad detrital zircon spectrum, probably had a prolonged, complex history. Its lack of tangible preservation suggests there has been very efficient crustal recycling.

3.8 SUMMARY To constrain the temporal evolution of the oldest par t of the Pilbara Craton, precise SHRIMP U-Pb zircon ages were obtained from the Pilgangoora Belt, Carlindi Granitoid Complex and North Pole Dome. The Coonterunah Group was deposited at ~3517 Ma and then intruded by Carlindi granitoids between ~3484 and 3468 Ma. Metamorphism of the Coonterunah succession was probably protracted, but broadly coincidental with granitoid emplacement. The Coonterunah-Carlindi terrain was uplifted and eroded to form an erosional unconformity onto which the £3458 Ma Warrawoona Group was deposited. There is now increasing evidence that the upper part of the Warrawoona Group (Salgash Subgroup) was deposited at ~3434 Ma, and the lithostratigraphy has been re-interpreted to accommodate this age. Although no pre- Coonterunah zircons were found within the Pilgangoora-Carlindi samples, an increasing number of ancient xenocrystic and detrital zircons have been found in the Pilbara Craton. They suggest that earlier continental crust played an important role in forming the Pilbara, but was efficiently recycled. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 89

Chapter 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS

4.1 INTRODUCTION The composition of a rock may be variably influenced by different processes such that unravelling its chemical evolution can be quite difficult. One major problem in ancient greenstone successions is separating the effects of metasomatic alteration from original magmatic features. However, if the affects of these processes can be recognised, then important constraints can be placed on greenstone evolution. Of particular interest here is determining the original magmatic composition of the Pilgangoora greenstones, so they can be used to constrain the petrogenesis of these ancient successions, their tectonic setting and, indeed, the composition of the Archaean mantle. For example, do the Coonterunah and Warrawoona basalts have distinct compositions, and thus, does the unconformity mark a petrogenetic and possible tectonic change? In this chapter, whole-rock elemental abundances of 39 mafic samples (24 Coonterunah Group, 15 Warrawoona Group) are presented. Analytical methods include glass-fusion and pressed-powder XRF and solution-nebulisation-Inductive-Coupled- Plasma-Mass Spectrometry (SN-ICP-MS). Prior to detailed interpretation, the quality of chemical analyses is verified, with particular emphasis on dissolution during ICP-MS sample preparation. The Coonterunah and Warrawoona basalt compositions are then presented, interpreted and compared with other Pilbara data. Details of the Coonterunah felsic volcanics and Carlindi granitoids are presented in Chapter 5. All elemental data are presented in Appendix 2.

4.2 PREVIOUS WORK Pilbara-wide geochemical studies have determined the com position of hundreds of greenstone samples (Glikson & Hickman, 1981a, b; Glikson et al., 1986a, b). However, only a few of these have come from the Pilgangoora Belt, including 53 from a stratigraphic traverse ~4 km west of Strelley Pool (Glikson et al., 1986b). Of these, 34 were from the Coonterunah Group (then Talga Talga Subgroup) and 19 from the Warrawoona Group (Euro Basalt). Major element whole-rock analyses showed that these units have variably altered magmatic compositions, but are dominantly metamorphosed tholeiitic and magnesian basalts. Possible stratigraphic variations were noted within the Warrawoona basalts (Van Kranendonk, 2000), a trend that mirrors the CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 90 mapped distribution of magnesian and tholeiitic basalts. These analyses were performed by XRF and neutron activation techniques, and concentrations of many trace elements were not significantly greater than detection limits. Consequently, these data provide little basis for rigorous interpretation. Other geochemical studies in the Pilgangoora Belt were funded by mineral exploration companies, including reconnaissance sampling of the entire belt (Sipa Resources, Lynas Gold) and detailed work around gold mines in the west (Lynas Gold; Neumayr et al., 1998). Proprietary data have not been used in this study. Moreover, since the Coonterunah Group has not been recognised elsewhere in the Pilbara, there are no comparative analyses. In contrast, there are a number of precise whole-rock elemental analyses of Warrawoona basalts in the North Pole Dome (Green et al., in prep.), Marble Bar Belt (Gruau et al., 1987; Jochum et al., 1991) and Kelly Belt and McPhee Dome (Barley, 1980; Barley et al., 1984, 1998). These studies provide a relatively complete dataset for the entire Warrawoona stratigraphy and allow regional variations to be compared. Precise geochemical studies of younger supracrustal successions include the ~3.24 Ga Sulphur Springs Group (Brauhart et al., 1998; Vearncombe & Kerrich, 1999), the ~3.0- 3.2 Ga west Pilbara Regal basalts (Ohta et al., 1996; Sun & Hickman, 1999) and the late Archaean Fortescue Group (Nelson, 1992).

4.3 ANALYTICAL METHODS Samples (~5 kg) were collected from least-altered in-situ exposures of representative rock-types, generally away from visible veins, intensely bleached domains and shear zones. They were cut into ~2 cm thick slabs and all oxidised and vein material was removed. Samples were then broken into ~1 cm fragments in a jaw crusher and any altered fragments were discarded. Most samples were powdered using an alumina mill, but some were powdered in a Cr-steel mill. No systematic compositional variations were observed from powders obtained from the different mills. All equipment was precontaminated by sacrificing a portion of each sample, thus reducing the affects of cross-contamination between samples. A thin-section was made for each sample. The concentration of over 50 elements was determined for each whole-rock sample. Three analytical techniques were used because each technique has an optimum detection range. For instance, ICP-MS is inaccurate at very high concentrations and so was only used to determine trace element abundances, whereas the opposite holds true CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 91 for glass-fusion XRF. Generally, these two techniques should have been adequate in determining all the desired elements for this study. However, due to dissolution problems during ICP-MS sample preparation, pressed-powder XRF techniques were required to determine Zr contents. Most of the basalts were analysed at the Australian National University (Research School of Earth Sciences) and the rest at Memorial University, Newfoundland.

4.3.1 ICP-MS versus pressed-powder XRF Inductively coupled plasma-mass spectrometry has found favour as an analytical tool for studying trace elements in geological samples because it offers cheap, rapid, simultaneous analysis of many elements, relatively simple sample preparation, very low detection limits, a large dynamic range and generally interference-free spectra (Jenner et al., 1990; Jarvis, 1997). Improved precision and extended element coverage have been obtained by calibrating with external standards and both elemental and enriched-isotope internal standards (Eggins et al., 1997). This method was adopted for this study because it has been successfully applied to a wide range of rock types and mineral separates and was found to be matrix tolerant (Eggins et al., 1997). For the Pilbara samples, the relative concentrations o f Zr and Hf obtained by ICP-MS were found to be unusually low, and so all samples were re-analysed using pressed-powder XRF. Hafnium was not measured by XRF, but Hf abundances of ten samples were obtained during Lu-Hf isotope analyses (Chapter 6). Random errors and detection limits aside, the ICP-MS concentrations correlate well with the XRF results except for Zr, which at low concentrations typically has much lower ICP-MS abundances. It is necessary, therefore, to decide which technique yielded the more precise Zr concentrations. To this end, the abundances obtained by ICP-MS and XRF of the Pilgangoora and twenty-eight North Pole basalts (Green et al., in prep.) are compared. These samples were chosen because the preparatory and analytical methods were the same for the ICP-MS analyses, but represent three different sample batches. For instance, two batches were analysed at the ANU (including the North Pole basalts) but months apart, whereas seven basalts were analysed at Memorial University. Furthermore, the general similarity between basalt compositions should have provided a relatively constant matrix effect. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 92

120 Rb Sr Y

40 I I C 300 C P P I C - - M M P S - S

80 M

( ( S p p p p ( m m p

200 p ) ) m ) 20

40 100

XRF (ppm) XRF (ppm) XRF (ppm) 40 80 120 100 200 300 20 40 Zr Nb Ba

200 I

C 1500 P - I M I 8.0 C C S P P - ( - M p M p S S m

(

( 1000 p ) p p p m m

100 ) ) 4.0

500

XRF (ppm) XRF (ppm) XRF (ppm) 100 200 4.0 8.0 500 1000 1500

Figure 4.1: Comparison of ICP-MS and XRF data. Pilgangoora basalts are solid triangles (ANU) and open squares (Memorial; both reported here), North Pole basalts are open circles (Green et al., in prep.). Dashed line is 1:1 correlation, solid line is best fit of data forced through the graphical origin.

Abundances of Rb, Sr, Y, Zr, Nb and Ba are generally above detection limits for both methods, and thus comparison is meaningful. For these elements, XRF detection limits are 1.0 ppm, except for Ba which is 10 ppm (Fitton, 1997), and ICP-MS limits are 0.01 ppm (Eggins et al., 1997). Other elements (e.g. Ce, Pb, Th, U) have concentrations which are generally below XRF detection limits (Fitton, 1997) and are not compared. Where both analytical techniques are accurate their abundances should define a perfectly linear 1:1 correlation. This is shown in Figure 4.1, where the dashed line represents a 1:1 correlation and the solid line represents the least-squares best fit of the data forced though the graphical origin. Clearly, Sr is the only element with a perfect 1:1 correlation, and on average Rb and Y have recorded higher and Zr, Nb and Ba have recorded lower ICP-MS abundances. However, these least-squares regressions are greatly influenced by deviations at high values, and so the concentrations are also plotted as percentage variance relative to the ICP-MS concentration (Fig. 4.2). There are three types of relationships. For Sr, the relative variances are small (<10 %), except for one sample, and there are no systematic changes of variance as the concentrations CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 93 decrease. Therefore, the Sr abundances correlate closely for both analytical methods as expected from Figure 4.1. For Rb, Y, Nb and Ba the relative variances are small for high abundances (generally <20 %, except for Ba) with a symmetrical increase of variance (around the zero percentage) as the abundances decrease (possibly except for Y which has only a few low-concentration samples). Therefore, both analytical methods correlate well for these elements, except for concentrations near the detection limits. The symmetry of variance at low concentrations indicates that these errors were generally random. For Zr, the relative variances are less than 50 % for concentrations greater than 50 ppm with the relative variance negatively increasing as the concentrations decrease. Therefore, the Zr values correlate poorly between methods, with a strong bias towards lower ICP-MS abundances, despite the concentrations being well above detection limits.

50 50 e e c 0 c 0 n n e e r r e e f f

f -50 -50 f i i d d

% -100 -% 100 -150 Rb (ppm) -150 Zr (ppm) 40 80 120 50 100 150 200

50 50 e e

0 c

c 0 n n e e r r e e f f -50 -50 f f i i d d

% % -100 -100

-150 Sr (ppm) -150 Nb (ppm) 100 200 300 2.0 4.0 6.0 8.0 10.0

50 50 e e

c 0 c 0 n n e e r r e

f e -50 f f -50 f i i d d

% -100 % -100

-150 Y (ppm) -150 Ba (ppm) 10.0 20.0 30.0 40.0 200 400 600 800 1000 1200

Figure 4.2: Percentage variance of ICP-MS and XRF abundances relative to ICP-MS abundances for individual samples plotted versus the ICP- MS abundance. Symbols used are as for Figure 4.1.

In addition to graphical representations, statistical tests were performed to determine the probability that the ICP-MS and XRF abundances were equivalent. For instance, it was expected that the differences between analyses for all the elements, except Zr, were random. The Wilcoxon matched pair test provides a straightforward CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 94 method to determine the probability that values of the same entity are equivalent. The test calculates whether the differences between measured values for each sample were drawn from a population with a mean of zero. The validity of the Wilcoxon test does not require that the population of differences was normally distributed, and so the test is very robust. The results are presented in Table 4.1.

Table 4.1: Wilcoxon matched pair tests to compare XRF and ICP-MS, and ID and ICP-MS results. p-levels were two-tailed for t-tests. Samples with concentrations below XRF detection limits were removed. Rb Sr Y Zr Nb Ba Lu Hf N 60 67 67 67 67 64 10 10 T 481 882 621 313 371.5 735 5 0.000 Z 3.1949 1.6054 3.2358 5.1597 4.7943 2.0397 2.2934 2.8031 p 0.0014 0.1084 0.0012 0.0000 0.0000 0.0414 0.2183 0.0051

The p-levels represent the probability of error in rejecting the hypothesis that the relationship exists in the population. In this case, the hypotheses are that the two methods were equivalent and that the differences were normally distributed. Low p- levels do not support these hypotheses. Therefore, Sr has the highest probability that the analytical methods were equivalent, whereas Zr and Nb have zero probability. The low p-levels for Rb, Y, Zr and Nb show that there is little correlation between analytical methods. For Rb and Nb, these low p-levels were most likely caused by increasing errors as the abundances approached the XRF detection limit. Moreover, the relative errors at higher concentrations of Rb and Nb are small (Fig. 4.2) and generally within analytical error. For Y and Zr, however, the abundances are well above the XRF detection limits, and so there must have been some bias in one of the analytical methods. At low Y abundances, the ICP-MS values are much smaller than the corresponding XRF values (negative skew in Fig. 4.2), but overall the ICP-MS abundances for Y are greater than those obtained by XRF (Fig. 4.1). Nevertheless, the relative errors of Y are consistently <20 %, and so the differences between techniques are minimal at these higher abundances. For Zr, however, the relative errors are much larger at low concentrations, and the poor correlation indicates a profound difference between ICP-MS and XRF. In addition to pressed powder XRF, ten samples were also analysed for Lu and Hf by isotope dilution-mass spectrometry (ID-MS) during Lu-Hf isotope studies (Appendix 3). Detection limits for ID-MS are even lower than for ICP-MS. Comparison CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 95 of ID-MS and ICP-MS (Table 4.1) shows a stronger correlation for Lu than Hf, although these populations are quite small. Importantly, the Hf values obtained by ICP- MS are consistently less than those by ID-MS, and, moreover, the samples with the greatest Hf deviations have correspondingly high Zr deviations. Consequently, Zr and Hf abundances have been either underestimated by ICP- MS or overestimated by XRF and ID-MS. Standards analysed during ICP-MS (MRG-1- gabbro, NIST688-basalt) and XRF (BHVO-basalt, SY2-syenite) have compositions close to their expected values (Govindaraju, 1989), and so do not indicate which methods provided incorrect Zr-Hf abundances. However, if the discrepancies are due to Zr-Hf overestimates, then there must be systematic errors between two quite dissimilar analytical methods, pressed-powder XRF and ID-MS, which is unlikely. Therefore, ICP- MS has probably underestimated Zr-Hf abundances, and it follows that XRF must have produced more accurate Zr analyses, particularly at low abundances (<70 ppm). As a result, pressed-powder XRF contents are used for Zr hereafter and reliable Hf abundances were obtained for only 10 samples.

4.3.2 Interpretation ICP-MS analyses were undertaken using the same preparatory and analytical methods; twice at the ANU and once at Memorial University. Falsely low Zr abundances were obtained during all three analytical sessions, and so Zr depletion was not a specific problem associated with a single batch. Since Sr 86, Y89, Zr91 and Nb93 have approximately equal masses, they should behave similarly in the mass spectrometer. These elements also have masses between the Sr84 and Rh103 internal standards, and Zr, Y and Nb have no significant isobaric or molecular interferences (Eggins et al., 1997). The Sr, Y and Nb abundances are comparable between techniques, and so the mass spectrometers seem to have been working correctly. This is confirmed by the small differences (<3 %) of these elements for five repeat samples at Memorial University (Appendix 2). Given its favourable mass position, there is little reason to suspect that the Zr concentrations do not reflect those in the analyte which was fed to the nebuliser. A similar argument can be advanced for Hf 178, which is positioned between the Tm 169 and Re189 internal standards, although there may be significant molecular interferences from HREE-oxides at these masses (Eggins et al., 1997). Consequently, Zr-Hf must CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 96 have failed to enter the analyte or were removed from the analyte prior to entry into the mass spectrometer. The most likely causes of Zr-Hf depletion in the analyte are failed digestion of Zr-Hf-rich minerals (e.g., zircon) or formation of Zr-Hf-bearing residues during sample preparation. Clarification of digestion issues is hampered by a poor understanding of the Zr-Hf-bearing phases in metamorphosed basalts. These highly incompatible elements would have initially concentrated in the residual basaltic glass, and not magmatic minerals, but the Zr-Hf-bearing phases formed during devitrification and subsequent metamorphism are not known. That aside, at greater Zr abundances there should be corresponding increases in the volume of Zr-rich minerals, and hence, if these minerals failed to completely dissolve, then there should be a positive correlation between the Zr concentration and the magnitude of the error. The largest errors, however, are found at the lowest Zr abundances (Fig. 4.1-2), which suggests that Zr-bearing minerals were completely dissolved. Moreover, the Carlindi granitoids contain magmatic zircon and their ICP-MS abundances were depleted in Zr-Hf but not Y and HREE (see Chapter 5), as expected for failed zircon digestion (Deer et al., 1992). The depletion for different rock-types also shows that the problem was not matrix specific. It thus seems unlikely that failed sample digestion caused the Zr-Hf depletion. It has been noted in other laboratories that HFSE, such as Zr and Hf, will adhere to container walls during and after HNO 3 digestion (Münker, 1998; Barth et al., 2000). Since the surface area of the containers is constant for all samples, then the amount of material directly adsorbed should also be relatively constant, and so the samples with the lowest concentrations should have the largest relative errors. This is consistent with the Pilbara results, and thus is probably the cause of Zr-Hf-depletion. To prevent this adsorption, dilute HF can be added to the analyte after HNO 3 digestion (Barth et al., 2000). In summary, the trace-element abundances obtained by ICP-MS are probably as precise as expected (Eggins et al., 1997), except for Zr, which was consequently determined by pressed-powder XRF, and Hf, which was determined by ID for ten samples.

4.4 MAJOR ELEMENTS CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 97

The ten major element oxides plus volatiles (water, CO 2) typically account for >98 wt% of a volcanic rock, which is true of the Pilgangoora basalts. Fresh basalts are typically composed of anhydrous minerals, such as plagioclase and pyroxene, and so have low volatile contents (<2 wt%; Le Maitre et al., 1989). The Pilgangoora basalts, however, contain abundant hydrous metamorphic minerals, such as chlorite and amphibole, and so have been hydrated. Volatile contents were determined by heating sample powders and measuring the weight of material lost at 110 ° C (moisture within the sample; H2O-) and 1100° C (water held within minerals; H 2O+). The combined weight is known as the loss-on-ignition (LOI). Most of the Coonterunah basalts have LOI <2 wt%, with four samples >6 wt%, and for the Warrawoona basalts the LOI is between 3 and 6.5 wt% (Fig. 4.3; Appendix 2). There is generally little overlap between the LOI of the two successions, suggesting a metamorphic grade control. Volatile contents may increase during subaqueous eruption, low-grade metamorphism or other thermal events (deformation, veining, igneous intrusion) or decrease as low-metamorphic-grade minerals break down (devolatilise) during high- grade metamorphism. These processes involve migrating fluids which may dissolve the country rock, deposit previously dissolved material or both. The Pilgangoora basalts were probably altered during subaqueous eruption, low-temperature metamorphism, igneous intrusion and local deformation. To determine the intensity of associated chemical modification and whether original magmatic compositions can be deduced, element pair plots (variation diagrams) are used. In comagmatic suites, geochemical trends should define linear segments that reflect the crystallising mineral assemblage, whereas altered successions should have variable or chaotic compositions. It should be noted that linear covariation of element pairs is not proof of comagmatism, as basalt suites have limited ranges of major element compositions, but it does provide strong evidence against post-magmatic alteration of those elements. To account for the change of volatile content, the major elements have been recalculated to a dry weight to allow for the constant sum effect. For example, if

12 wt% H2O was added to a basalt, then the other major elements must decrease accordingly to produce a constant sum of 100 %. If a smaller quantity of H 2O was added to the same sample, then the change to the other elements would be smaller. Plotting the compositions of the hydrous basalts would thus produce spurious results. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 98

This is not an issue for trace elements as their abundances are measured relative to external standards.

y = 173.4x R = 0.640 y = 104.5x +79.9 R = 0.869 Zr MnO (ppm) 0.25 (wt%) 200 y = 60.1x R = 0.872 0.20 150

0.25 y = 47.4x 100 R = 0.827 0.10

50 0.05

SiO MgO 2 K2O 10.0 (wt%) 2.5 60 (wt%) (wt%)

8.0 2.0 55

6.0 1.5 50 4.0 1.0

45 2.0 0.5

y = 0.1785x Al2O3 CaO r = 0.879 P2O5 18.0 12.0 0.3 (wt%) (wt%) (wt%)

10.0 16.0 y = 0.0926x 8.0 0.2 r = 0.921 14.0 6.0 12.0 4.0 0.1 y = 0.092x r = 0.808 10.0 2.0

20.0 Fe2O3 6.0 Na2O LOI (wt%) (wt%) 10.0 (wt%) 5.0 8.0 16.0 4.0 6.0 3.0 12.0 4.0 2.0

2.0 8.0 1.0

0.0 0.5 1.0 1.5 2.0 2.5 3.0 0.0 0.5 1.0 1.5 2.0 2.5 3.0 0.0 0.5 1.0 1.5 2.0 2.5 3.0

TiO2 (%) TiO2 (%) TiO2 (%) Figure 4.3: Major element variation diagrams for Coonterunah and Warrawoona basalts, recalculated to dry-wt% (except Zr, LOI). Linear correlation of TiO 2 with Zr indicates that TiO 2 was relatively immobile after deposition, and hence diagrams can be used to assess the degree of post-magmatic alteration. Arrow in SiO 2 is expected fractionation trend. Solid diamonds and lines are Coonterunah Group, those circled are eccentric Coonterunah basalts, unfilled squares and dashed lines are Warrawoona Group. To use variation diagrams successfully, a major element which can be regarded as immobile needs to be selected as an independent variable. In this study, TiO 2 has been chosen. When plotted against Zr, a typically immobile element, a general linear trend is CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 99

defined (Fig. 4.3), supporting the immobility of TiO 2. Moreover, Zr versus TiO2 has a positive correlation and intersects the graphical origin, so both elements behaved similarly during magmatism; that is, both were incompatible. Therefore, during petrogenesis of the Coonterunah and Warrawoona basalts, Zr and TiO 2 were preferentially enriched into the melt, and so the more evolved basalts have higher Zr and

TiO2 abundances. Six Coonterunah basalts have higher Zr abundances at similar TiO 2 abundances than the other Pilgangoora basalts, but these define a separate linear trend which does not intersect the graphical origin (Fig. 4.3). This still suggests that Zr and

TiO2 were immobile, but these particular basalts probably formed by different processes than the majority of the Pilgangoora basalts. In basaltic systems, silica is incompatible (Le Maitre et al., 1989), and so for the

Pilgangoora basalts, it should define a positive correlation with TiO 2. Such a correlation is observed in about half of the basalts, particularly the Warrawoona basalts, with the remaining basalts, except one, having relatively high silica abundances (Fig. 4.3). The six anomalous basalts define a negative correlation. Given that TiO 2 was immobile, then silica was probably added to many basalts during post-magmatic alteration. This implies that Harker diagrams, the standard plot of major elements versus silica, would be (total) uninformative if used for these samples. Al 2O3 and Fe2O3 versus TiO2 show moderate degrees of scatter around a general positive linear trend which suggests that (total) both were incompatible and that post-magmatic modification of Al 2O3 and Fe2O3 was variable but minor. MgO and CaO have negative correlations with TiO 2 which indicates that they were compatible and have been variably altered. For MnO, Na 2O and

K2O, there is no correlation with TiO 2 which suggests intense post-magmatic modification of these elemental abundances. Some samples have been severely depleted in CaO, Na2O and K2O. There is a strong positive correlation of P 2O5 with TiO 2. In summary, the magmatic abundances of most of the major elements, except TiO 2 and

P2O5, have been modified after eruption, with Na 2O and K2O showing the most severe alteration. During Coonterunah and Warrawoona basalt petrogenesis, TiO 2, SiO2, Al2O3,

Fe2O3 and P2O5 were incompatible, whereas MgO and CaO were compatible. It is not possible to constrain the magmatic behaviour of Na 2O and K2O, although they are typically incompatible during basalt petrogenesis (Le Maitre et al., 1989). Due to post-magmatic alteration, classifying the basalts according to their major elements is problematic. For example, the silica content of nearly half of the basalts CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 100

would indicate that they were basaltic andesites (SiO 2 > 52 %) with two andesites (SiO2 > 57 %; Le Maitre et al., 1989). As shown above, however, the high silica content of many samples was probably due to post-magmatic addition. Further subdivision of the basalts based on alkali abundances would also be suspect. Therefore, the basalts have not been reclassified according to their major element compositions and their field names have been retained because the field criteria probably provide a good proxy for the whole-rock composition. CIPW normative calculations (Cross et al., 1902) are conventionally used to approximate the modal mineral abundances of vitric basalts. They are used here to provide supporting evidence for the pre-metamorphic mineral assemblages of the Pilgangoora basalts (Appendix 2), although this approach has many shortcomings. Post- magmatic alteration of SiO 2, Na2O and K2O is the greatest problem, and so the most altered samples have been excluded. In general, the initial basalt compositions can be modelled as a mix of plagioclase, orthopyroxene and clinopyroxene, with accessory quartz or olivine, orthoclase, magnetite, ilmenite and apatite. Plagioclase compositions and orthoclase abundances are suspect due to Na2O and K2O alteration, respectively.

The presence of quartz or olivine is strongly influenced by the SiO 2 abundance. As silica has been added post-magmatically in many samples, the basalts may have initially been olivine normative. The Fe 2O3/FeO ratio is also critical in normative calculations. For the

Pilgangoora basalts, the iron content was measured as total Fe 2O3, but can be converted using the formula

Fe2O3/FeO = -0.142 + 0.0072 x SiO2 + 0.013(Na2O + K2O) [Le Maitre, 1976].

This conversion, however, is suspect due to severe alteration of all the elements used in the calculation, although the final ratios are between 0.24 and 0.36 which approximate proposed standards (Middlemost, 1989). Recalculating to an Fe 2O3/FeO ratio of 0.2, a proposed basalt standard, decreases quartz (<1.25 %) and magnetite (<1.5 %), increases olivine (<3.7 %) and variably changes pyroxene (<2.5 %) abundances. These trends also suggest that the Pilgangoora basalts tended towards olivine normative. The above problems aside, a plagioclase, pyroxene and FeTi-oxide assemblage is consistent with pseudomorphs and relict minerals found in the basalts (Chapter 2). CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 101

In the Coonterunah Group, the magnesian and pyroxene-phyric basalts have high

MgO, SiO2 and CaO and low TiO 2 and Al 2O3 concentrations, whereas the gabbros are more enriched in TiO 2 and Fe2O3 and less enriched in SiO 2 and CaO. The tholeiitic basalts have a broad compositional range probably because they represent more samples and a greater stratigraphic thickness. In general, they have the highest TiO 2 and lowest MgO and CaO abundances. In the Warrawoona Group, the relationships between the rock-types and major elements are generally similar to those in the Coonterunah Group. However, the magnesian basalts are more variable and the pyroxene-phyric basalts are more enriched total in TiO 2, Fe2O3 and MgO and depleted in CaO at equal silica contents.

As shown above, the basalts have retained their magmatic TiO 2 abundances, and since it was incompatible during petrogenesis, then the least evolved basalts should have the lowest TiO2 concentrations and more evolved basalts should have higher concentrations. Hence, magnesian basalts are interpreted to be the least evolved, with pyroxene-phyric basalts slightly more evolved and gabbros even more evolved. Tholeiitic basalts are the most evolved. Conversely, MgO was compatible during basalt petrogenesis, and so its abundance also measures basalt evolution, with higher abundances for less evolved basalts. This criterion concurs with the above TiO 2 trend. Therefore, the basalts appear to represent a compositional continuum from magnesian to tholeiitic. Moreover, the basalts also define general magmatic trends and so each basalt suite may be cogenetic. Six Coonterunah tholeiitic basalts do not plot on the general magmatic trend (gc-

09, 25, 43, 45, 48-0697, 380698) and typically have higher Zr, SiO 2, Na2O and P2O5 and lower Fe2O3, CaO, MgO abundances at any given TiO 2 (Fig. 4.3). Four of these samples have very high volatile contents (LOI >7 wt%), but the other two have volatile contents similar to the main group of Coonterunah basalts (<2 wt%). In a few variation diagrams, these eccentric basalts display discrete linear trends consistent with a magmatic origin, and so their compositional differences were probably not solely due to alteration. They also have distinct trace element compositions.

4.5 TRACE ELEMENTS Forty-six elements were analysed by ICP-MS; 45 at the ANU (not Tm) and 30 at Memorial University (not W, Ag, Cd, Sn, Sb, Be, P, Sc, Ti, V, Cr, Co, Ni, Cu, Zn, Ga). CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 102

For all samples, Cr 2O3 and V2O5 were measured by glass-fusion XRF and Sc, V, Cr, Ni, Cu, Zn and Ga were determined by pressed-powder XRF (Appendix 2). Thus, only seven (W, Ag, Cd, Sn, Sb, Be, Tm) of the 46 elements were not analysed for all samples. Trace elements have been plotted on variation diagrams to determine whether their magmatic abundances have been modified by post-magmatic alteration (Fig. 4.4). In these diagrams, only ICP-MS abundances have been used, except for Zr (pressed- powder XRF, see section 4.3.1), which is the comparator for all the plots. Anomalous samples have been removed from nine diagrams to assist in scaling; these are listed in Table 4.2. A summary of the correlation between each element and Zr is presented in Table 4.3. The eccentric Coonterunah basalts, noted from the major elements, are also differentiated in the trace element plots. In general, each element behaved similarly in bo th basalt suites, except for V, which has a negative Coonterunah and a positive Warrawoona correlation. The Coonterunah correlation for V, however, is strongly influenced by the eccentric tholeiitic basalts. Elements with poor correlations are interpreted to have been modified during post-magmatic processes, and thus have not preserved their magmatic abundances. These include large ion lithophile elements (LILE) such as Li, Ba, Sr and Rb, which are known to be mobile during sea-floor alteration (Shiano et al., 1993). In contrast, elements with well-defined correlations, for example, the high field strength elements (HSFE), such as Th, Ta, Nb, Zr and LREE (La to Sm), are interpreted to have preserved their original magmatic abundances. Elements with positive correlations were incompatible during basalt petrogenesis, and where the trends intersect the graphical origin, as shown for most of the HFSE, their behaviour was comparable to Zr (highly incompatible). This implies that in these basaltic melts the HFSE had partition coefficients much less than unity, and that no crystallising phases concentrated these elements. This is borne out by the absence of petrographic evidence for zircon crystallisation. The Ga trend is displaced for all samples towards relatively higher abundances and does not intersect the graphical origin. This shows that Ga was more compatible than Zr, but still slightly incompatible. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 103

Li 300 Zn 0.12 Cd 30

200 0.08 20

10 100 0.04

Be 2.00 Ga 1.60 Sn

20 1.20

1.00 0.80 10

0.40

Sc Rb 1.60 Sb 60 60 1.20

40 40 0.80

20 20 0.40

V 400 Sr Cs 0.60 400 300

0.40 200 200 100 0.20

1000 Cr 40 Y 400 Ba

30 300 600 20 200

200 10 100

Co 10.0 Nb La 20.0 60 8.0

40 6.0 10.0 4.0 20 2.0

Mo Ce 200 Ni 1.2 40.0

0.8 100 20.0 0.4

Cu Ag 6.0 Pr

200 0.15 4.0

100 2.0 0.05

0 50 100 150 200 0 50 100 150 200 0 50 100 150 200 Zr Zr Zr Figure 4.4: continued overleaf. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 104

Nd 1.60 Ho 0.80 W

20.00 1.20 0.60

0.80 0.40 10.00 0.40 0.20

Sm 4.00 Er 0.40 Tl

4.00 3.00 0.30

2.00 0.20 2.00 1.00 0.10

Eu 0.60 Tm 4.00 Pb 1.60

3.00 1.20 0.40

2.00 0.80 0.20 0.40 1.00

Tb 4.00 Yb Bi

0.80 3.00 0.08

2.00 0.40 0.04 1.00

6.00 Gd 0.60 Lu 4.00 Th

3.00 4.00 0.40

2.00

2.00 0.20 1.00

Dy Ta 0.80 U 6.00 0.60 0.60

4.00 0.40 0.40

2.00 0.20 0.20

0 50 100 150 200 0 50 100 150 200 0 50 100 150 200 Zr Zr Zr Figure 4.4: Trace element variation diagrams of Pilgangoora basalts show intensity of post-magmatic alteration and magmatic compatibility of unmodified elements. Abundance in ppm, elements in order of atomic mass. Some highly anomalous samples have been deleted, see main text for details. Solid diamonds are main Coonterunah basalts, circled diamonds are eccentric Coonterunah basalts, open squares are Warrawoona basalts. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 105

Table 4.2: Anomalous samples removed from variation diagrams (Fig. 4.4). Sample # Conc. Sample # Conc. (ppm) (ppm) Li gc110697 (Warr.) 118 W gc210697 (Coont.) 2.62 Cd gc120697 (Warr.) 0.82 Tl gc210697 (Coont.) 0.719 Sn gc360697 (Warr.) 3.70 Pb 390798 (Coont.) 10.5 Ba 210898 (Warr.) 1094 Bi gc210697 (Coont.) 0.198 Cs gc210697 (Coont.) 1.53

Table 4.3: Summary of trace element correlations versus Zr (Fig. 4.4). Correlation Poor Positive Negative elements Li, Cu, Zn, Rb, Sr, Be, Ga, Y, Nb, Sc, Cr, Co, Ni, Mo, Ag, Cd, Sb, Sn, La to Lu V (Coont.) Cs, Ba, W, Tl, Bi (REE), Ta, Pb, Th, U, V (Warr.)

Elements with negative correlations, such as Cr, Ni, Co and Sc, were compatible, and provide evidence for the crystallising phases. Both chromium and nickel have inflections in their variation diagrams which reflect changes to the crystallising assemblage. For the least evolved basalts (low Zr content), there were probably minerals that preferentially and rapidly incorporated Cr, but when the magmas evolved to Zr abundances of ~50 ppm, Cr became undersaturated in the magma and this Cr-rich phase stopped crystallising. However, the more gradual decrease of Cr contents for even more evolved basalts indicates that Cr remained compatible with the remaining assemblage. This inflection suggests that a Cr-rich spinel crystallised in the least evolved basalts. Reconnaissance work by laser-ablation ICP-MS has found that the magnetite in the Coonterunah magnesian basalts is relatively enriched in chromium, consistent with this inference. The Ni data also define a negatively correlated curve which suggests that this element was also more readily incorporated into crystallising phases at less evolved compositions, but was compatible over the entire compositional range. The compatibility of Sc indicates that crystallisation of pyroxene was more important than olivine, and the totality of compatible element data can be attributed to early magnetite- pyroxene crystallisation evolving to a pyroxene-rich assemblage (Arth, 1976; Schock, 1979). This is consistent with relict minerals and normative calculations. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 106

The eccentric Coonterunah basalts lie on the compatible trace element trends, and hence were composed of similar crystallising phases. However, for incompatible elements, they lie along the main trend for Nb, La, Ce, Pr, Ta, Th and U, but are displaced towards higher Zr abundances for Ga, Y and the remaining REE (Nd to Lu). This displacement is most pronounced for Y and the HREE. An intensely altered Warrawoona tholeiitic basalt (gc110697) shows similar behaviour, but since its other HFSE plot along the general magmatic trends and the eccentric basalts define separate linear trends, the eccentric basalts are probably not the result of post-magmatic alteration. The Ga, Y and M-HREE abundances of the eccentric basalts are comparable with the main suite, and thus the eccentric basalts have probably been enriched in Zr,

Nb, La, Ce, Pr, Ta, Th, U and P 2O5. This enrichment is also shown in variation diagrams where the eccentric basalts define the high abundance end of the magmatic trend (e.g. Zr versus La, Ce). Eccentric basalts aside, linear trends are well defined for La to Sm, but the trends become progressively less coherent from Eu to Lu, such that the Yb and Lu samples define fields bounded by two lines which intersect the graphical origin. For Y, there is a similar bounded field. Therefore, the elements which best distinguish the eccentric basalts also show variable behaviour for the main trend. These variations are shown by both Coonterunah and Warrawoona basalts.

To reveal any stratigraphic trends, TiO 2, Zr, Nb, Lu, La and Cr abundances are plotted against datum height (Fig. 4.5). In the Coonterunah Group, stratigraphic height was measured from the lowermost pyroxene-phyric basalt (gc020697, gc180697) which is exposed in several fault blocks. In the easternmost Pilgangoora Belt, the felsic volcanics are interpreted to correlate with those immediately below the Strelley Pool Chert in the central part of the belt, a link supported by petrographic, geochemical and age similarities. There are pronounced stratigraphic trends of basalt composition in the lower part of the Coonterunah succession, where the least-evolved basalts are confined to the base of the succession and successively overlying basalts are progressively more enriched in incompatible elements (Fig. 4.5). The TiO 2 trend has a pronounced inflection defined by the eccentric basalts, but this is not reflected by the trace elements. Importantly, the eccentric tholeiitic basalts are not stratigraphically confined. The stratigraphic base of the Warrawoona Group is measured from the base of the Strelley Pool Chert, and, in contrast to the Coonterunah basalts, they show no large-scale CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 107 stratigraphic trends, other than the mapped changes between magnesian and tholeiitic compositions. Gabbros have not been included in the stratigraphic height plots.

2.0 1.0

SPC SPC SPC

) 6.0 m (

t 5.0 FV FV FV h g i

e 4.0 h

. t 3.0 a r t

s 2.0

1.0 0.0

0.0 0.5 1.0 1.5 2.0 2.5 50 100 150 200 2.0 4.0 6.0 8.0 10.0

(%) (ppm) (ppm) TiO2 Zr Nb

2.0

1.0 SPC SPC SPC ) 6.0 m (

t

h 5.0 FV FV FV g i e

h 4.0

. t

a 3.0 r t s 2.0

1.0 0.0

0.0 0.1 0.2 0.3 0.4 0.5 0.6 5.0 10.0 15.0 20.0 200 400 600 800 1000

(ppm) (ppm) (ppm) Lu La Cr

Figure 4.5: Stratigraphic height versus TiO 2, Zr, Nb, Lu, La and Cr abundances for Pilgangoora basalts. Height of Coonterunah Group was measured from the lowermost pyroxene-phyric basalt, for Warrawoona Group from the Strelley Pool Chert (SPC). FV indicates probable stratigraphic position of felsic volcanics in eastern domain. Plots do not include gabbros. Symbols as for Figure 4.3.

In summar y, the HFSE have not been significantly modified by post-magmatic processes, and so for these elements the Pilgangoora basalts have preserved their magmatic compositions. All elements, except V, behaved similarly in both the Coonterunah and Warrawoona basalts. Compositional trends of compatible elements suggest early magnetite crystallisation in an overall pyroxene-dominated system. The eccentric Coonterunah basalts are enriched in some HFSE but this does not reflect post- magmatic alteration. There is also a breakdown of magmatic trends for these elements within the main basalt suite. There are profound stratigraphic trends of basalt composition in the lower part of the Coonterunah succession, but these are not repeated in the Warrawoona succession.

4.6 NORMALISED DIAGRAMS CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 108

Normalised diagrams show elemental variations of a sample relative to a standard composition. Widely used standards include C1-carbonaceous chondrites, primitive mantle and normal mid-oceanic ridge basalts (N-MORB). Elements are generally arranged in order of compatibility such that the more incompatible elements are to the left. As a result, typical magmatic processes such as partial melting and fractional crystallisation should produce straightforward changes to the plots. For example, evolved samples should be more enriched in incompatible elements and so have pronounced negative slopes in normalised diagrams. Moreover, magmatic processes which are not ‘typical’ may be reflected as anomalies. Normalising values used here are from Sun & McDonough (1989). The Pilgangoora basalts are plotted on N-MORB-normalised diagrams (Fig. 4.6; Sun & McDonough; 1989). The general element order of Pearce (1983) has been adopted, although Sr and K 2O may in fact be more compatible than Rb (Sun & McDonough, 1989). Plotting these more mobile elements to the left of Rb produces greater coherence between the immobile elements. Plots are arranged in stratigraphic order. N-MORB-normalised patterns are all characterised by the same shape: high Th, U and LREE and lower Ta, Nb, P, Zr, Ti, Y and M-HREE. Although most elements have preserved their original magmatic abundances, the prominent irregularities in Sr,

K2O, Rb and Ba in many diagrams reflect post-magmatic alteration of these elements. There are prominent negative P anomalies in all and pronounced negative Ta-Nb anomalies in most of the Pilgangoora basalts. There are also subtle negative Eu, Nd and

TiO2 and positive TiO 2 and Yb anomalies in some basalts. The background signature after removal of these anomalies shows an overall negative slope, produced by relatively greater concentration of incompatible elements. Therefore, the Pilgangoora basalts are enriched in incompatible elements relative to N-MORB. The eccentric Coonterunah basalts have steeper negative slopes and so are more enriched in the more incompatible elements. They also have the most extreme Ta, Nb and P anomalies. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 109

100 100 gc020697 (MBK) A E 390798 (MBT) gc180697 (MBK) 280698 (MBT) gc010697 (MBM) 380798 #(MBT) gc170697 (MBM)

B gc100697 (MG) gc160697 (MBM) 10 R 10 O M - N

/

E L P

M 1 1 A S

100 100 B gc210697 (MBT) F gc360697 (MBT) gc040697 (MBT) gc120697 (MG) gc200697 (MBT) gc370697 (MBM) gc030697 (MG) gc270697 (MBM) 10 B 10 gc190697 (MBT) R gc110697 (MBT) O M - N

/

E L 1 P 1 M A S

100 100 C gc250697 #(MBT) G gc290697 (MBT) gc070697 (MBT) 770798 (MBT) gc060697 (MBT) gc140697 (MBM) gc050697 (MBT) gc280697 (MG)

10 B 10 gc480697 #(MBT) R gc130697 (MBT) O M - N

/

E L P M A S 1 1

100 100 D gc090697 #(MBT) 750798 (MBM) H gc300697 (MBT) gc430697 #(MBT) gc450697 #(MBT) gc150697 (MBT) gc080697 (MG) 760798 (MBT)

10 B 10

gc260697 (MG) R 210898 (MBK) O M - N

/

E L P M A S 1 1

0.1 0.1 Sr Rb Th Ta La P Zr Ti Yb Sr Rb Th Ta La P Zr Ti Yb K Ba U Nb Ce Nd Eu Y K Ba U Nb Ce Nd Eu Y

Figure 4.6: N-MORB-normalised diagrams for Coonterunah (A-E) and Warrawoona (F-H) basalts (Sun & McDonough, 1989). Plots are in stratigraphic order (A is lowest) and bottom sample in each legend is lowest for that plot. MBT = tholeiitic, MBM = magnesian, MBK = pyroxene-phyric, MG = gabbro, # indicates eccentric basalts. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 110

In general, the tholeiitic basalts have the most enriched but variable incompatible element abundances followed by the gabbros, pyroxene-phyric and magnesian basalts. The Coonterunah and Warrawoona basalts are generally similar, except that the Warrawoona basalts are more enriched in Ba. This may be related to Ba mobility during hydrothermal alteration. Elsewhere within the Warrawoona Group, the North Pole chert-barite units contains thick beds and veins of barite (Buick & Dunlop, 1990), and such alteration may have been widespread. The composition of the Warrawoona pyroxene-phyric basalt is comparable with the tholeiitic basalts. For REE, the Pilgangoora basalts are also p lotted on C1-chondrite-normalised (Sun & McDonough, 1989) diagrams (Fig. 4.7). Patterns are generally flat for all Pilgangoora magnesian and pyroxene-phyric basalts and gabbros at approximately ten times chondritic. There are small LREE enrichments and depletions in a few samples. Gabbros have more enriched but less uniform REE contents than the magnesian basalts, which may reflect their greater stratigraphic extent and multiple emplacement events. Tholeiitic basalts have variable REE compositions, often with strong LREE enrichment. Greater LREE enrichment typically accompanies smaller increases of M-HREE contents. However, there are a few samples with slight La depletions relative to Nd. There are small negative and positive Eu anomalies, which were probably controlled by plagioclase removal and accumulation, respectively. There are no Ce anomalies. The eccentric basalts are more enriched in LREE and have steep L-MREE and gentle M- HREE negative slopes. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 111

100 100 A gc020697 (MBK) gc180697 (MBK) E gc010697 (MBM) 1 C

gc170697 (MBM) /

gc160697 (MBM) E L P 10 10 M A

S 390798 (MBT) 280698 (MBT) 380798 #(MBT) gc100697 (MG)

100 100 B F 1 C

/

10 E 10 L P gc360697 (MBT) gc210697 (MBT) M A

gc040697 (MBT) S gc120697 (MG) gc200697 (MBT) gc370697 (MBM) gc030697 (MG) gc270697 (MBM) gc190697 (MBT) gc110697 (MBT) 100 100 C G 1 C

/

E 10 L 10 gc290697 (MBT) gc250697 #(MBT) P M 770798 (MBT) A

gc070697 (MBT) S gc140697 (MBM) gc060697 (MBT) gc280697 (MG) gc050697 (MBT) gc130697 (MBT) gc480697 #(MBT)

100 100 D H 1 C

/

E 10 L 10

gc090697 #(MBT) P M

gc430697 #(MBT) A 750798 (MBM) S gc450697 #(MBT) gc300697 (MBT) gc080697 (MG) gc150697 (MBT) gc260697 (MG) 760798 (MBT) 210898 (MBK) 1 1 La Ce Nd Sm Eu Gd Dy Er Yb Lu La Ce Nd Sm Eu Gd Dy Er Yb Lu

Figure 4.7: C1-chondrite-normalised diagrams for Coonterunah (A-E) and Warrawoona (F-H) basalts (Sun & McDonough, 1989). Plots are arranged and symbols the same as Figure 4.6.

4.7 INTERPRETATION Whole-rock elemental abundances were determined by XRF and ICP-MS. Absorption of Zr and Hf to containers during ICP-MS sample preparation led to underestimates of their abundances, and so pressed-powder XRF and isotope dilution were used for these elements instead. Comparison of repeats, blanks and standards show that the whole-rock elemental compositions of all samples used here were accurate and precise. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 112

The Pilgangoora basalts have been metamorphosed at low temperatures such that their original mineral assemblages have been almost completely replaced. Their compositions have also been modified by post-magmatic processes, although the abundances of only some elements, particularly the major alkalis and LILE, have been significantly modified. Most elements have preserved their magmatic abundances and the following interpretations are based on them. Results from previous studies in the Pilgangoora and other Pilbara greenstone belts are compared in the discussion.

4.7.1 Relationships within and between basalt suites The ~3517 Ma Coonterunah basalts generally display magmatic trends on variation diagrams, with magnesian and pyroxene-phyric basalts least evolved, gabbros more evolved and tholeiitic basalts most evolved (Fig. 4.3, 4.4). This is consistent with the N-MORB- and chondrite-normalised diagrams where there are no major changes in pattern shapes with lithology, but the magnesian basalts are the least enriched and the tholeiitic basalts are the most enriched in incompatible elements (Fig. 4.6, 4.7). The overall enrichment differences between the basalts probably reflect partial melting or fractional crystallisation processes, where magnesian basalts most closely reflect the composition of the source. The profound stratigraphic trends in composition in the lower part of the Coonterunah succession indicate that basalt petrogenesis also involved a process whereby successive magmas, at least briefly, became more enriched in incompatible elements. The Warrawoona basalts were deposited >60 million years later and are separated from the Coonterunah basalts by an erosional unconformity. They also display magmatic trends with least-evolved magnesian basalts, more-evolved gabbros and most- evolved tholeiitic basalts (Fig. 4.3 to 4.7). There are no large-scale stratigraphic trends, other than those reflecting gross changes between magnesian and tholeiitic compositions. This agrees with the general interpretation of Glikson et al. (1986b), although their data suggest small-scale fractionation trends which were not observed here (Van Kranendonk, 2000). Thus, the Warrawoona basalts also represent a comagmatic suite related by straightforward magmatic processes. The Coonterunah and Warrawoona basalts have very similar geochemical characteristics. Indeed, their compositions overlap on elemental variation diagrams and their N-MORB- and chondrite-normalised plots have equivalent shapes and ranges. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 113

Therefore, the petrogenetic processes responsible for forming these two suites must have been similar, even though they formed >60 myr apart. It follows then that constraining the petrogenetic processes that formed the Warrawoona basalts should also constrain those for the Coonterunah basalts. Given that these are both comagmatic suites, then the compatible elements (Cr, Ni, Co, Sc) can be used to constrain the crystallising assemblage. Chromium and nickel trends suggest early magnetite crystallisation, whereas Sc and Co indicate that the system was controlled by pyroxene more than by olivine. The normative calculations, although not rigorous, show that the basalts have compositions consistent with an original plagioclase, pyroxene and magnetite composition. This is also supported by relict and pseudomorphed magmatic minerals. Therefore, within the suites, the basalts seem to be related by simple fractionation of the least evolved magma.

4.7.2 Crustal component All of the basalts have a crustal signature compared with N-MORB: that is, Ta- Nb depletion associated with Th, U and LREE enrichment. These characteristics are most readily explained by addition of a crustal component to a mantle-derived magma either by modification of the mantle source region, for example during subduction- related metasomatic processes (Brenan et al., 1994; Pearce et al., 1995), or by assimilation of crustal material during upward passage of the magma (Barley, 1986). There are no general geochemical criteria to discriminate between these two processes. To demonstrate this limitation, La/Nb ratios ar e compared with Nb to show the magnitude of the negative Nb anomaly and the fractionation between these elements (Fig. 4.8). Both elements have been normalised with N-MORB to simplify the interpretation. During typical basalt petrogenesis, La is slightly more compatible than Nb (Fig. 4.6), and so more evolved basalts should develop smaller La/Nb ratios. This is reflected by the general negative trend of the Pilgangoora basalt suites. However, subduction-related mantle metasomatism preferentially adds La to the mantle as it is more soluble (Perfit et al., 1980; Davidson, 1996; MacDonald et al., 2000), and as a result the La/Nb ratio of the mantle wedge increases markedly and derived basalts inherit a negative Nb anomaly (La/Nb > 1). For example, the South Sandwich island-arc basalts have much higher La/Nb ratios associated with a steep, negative trend (Pearce et al., 1995) and large, negative Nb anomalies which decrease CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 114 with fractionation. The Pilbara basalts show only slight La enrichment, and thus are more similar to enriched-MORB like that from the Pacific-Antarctic Ridge (Vlast élic et al., 2000). This shows that Nb anomalies are not as profound in the Pilbara basalts as those from some modern subduction systems, though their negative trend on the La/Nb versus Nb diagram indicates that they still reflect typical fractionation processes. Thus, the Pilbara basalts may indeed be just the result of partial melting and crystallisation of an enriched mantle source.

5.0 South Sandwich island arc

4.0

Carlindi average continental crust granitoids ) m

r 3.0 o n ( b N /

a 2.0 L * 1.0

PAR Hawaii fractionation vector

0.0 0.0 2.0 4.0 6.0 8.0 10.0 Nb(norm)

Figure 4.8: La/Nb versus Nb shows a general negative correlation caused by fractionation, but this breaks down at higher Nb values for the Pilbara basalts. Symbols as for previous figures. All elements normalised with N-MORB (Sun & McDonough, 1989). Unity is the average N-MORB; PAR are E-MORB from Pacific-Antarctic Ridge (Vlastélic et al., 2000); South Sandwich island arc (Pearce et al., 1995); high-MgO Hawaiian basalts (Wagner et al., 1998); continental crust (Barth et al., 2000); Carlindi granitoids (Chapter 5).

Average continental crust, however, has greater La/Nb ratios than N-MORB (La/Nb = 2.1, Nb = 3.4 - N-MORB normalised, from Barth et al., 2000) and so crustal addition also produces enriched basalt. Since such addition would precede or accompany crystallisation, then a trend like that of the Pilbara basalts may also result from crustal contamination. Moreover, local Archaean continental crust, as represented by Carlindi granitoids (see Chapter 5), has a wide compositional range, including components with large La/Nb ratios at lower Nb abundances, and thus greater Nb anomalies can be produced with quite small crustal additions. Importantly, elemental geochemistry cannot discriminate between subduction-related mantle metasomatism and CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 115 crustal contamination because they both involve similar changes and these are masked by fractionation. Perhaps significantly, the eccentric basalts have trace element abundances comparable to average continental crust, and hence, their negative Nb anomalies are associated with much greater La and Nb abundances. Whatever the cause, it is likely the Ta and P anomalies formed by equivalent processes. Given that the Warrawoona basalts were erupted onto continental basement, then they had ample opportunity to assimilate continental crust. Thus, the crustal component evident in the basalt geochemistry may reflect such assimilation. The geochemical similarities between the Warrawoona and Coonterunah basalts indicate similar petrogenetic processes, and so the latter could have also assimilated continental basement. The existence of a pre-Coonterunah continental basement is supported by the frequent occurrence of more ancient zircons in magmatic and clastic rocks elsewhere in the Pilbara (Fig. 3.20, Table 3.4). So, it is indeed possible that crustal assimilation was responsible for the geochemical signature of both suites of Pilgangoora basalts.

4.7.3 REE models Trace element modelling provides further evidence for crustal contamination of the Pilgangoora basalts. These models suggest that the REE contents of the most evolved basalts are reproduced more precisely by combined assimilation-fractional crystallisation of the least evolved basalts than by fractional crystallisation alone. In the models, plagioclase and clinopyroxene are assumed to be the main crystallising phases because they are the dominant minerals in the basalts, not only as relics and pseudomorphs, but also in normative calculations. Orthopyroxene has been included because it is required to form the compatible trace element trends and is also an important normative mode. Although no relict olivine has been found, it is included in the models because it features in many of the normative calculations and probably crystallised from the more primitive melt compositions. However, the predominance of pyroxene is consistent with the compatible trace element trends. Conversely, relict igneous hornblende is found in places but has not been included in the models due to its low modal abundance (<5 %) and the complexity of assigning partition coefficients to the hornblende series. Magnetite is also not included, but it has a REE partition coefficient near unity and so would not substantially influence these models (Schock, 1979). The models assume a 40 % plagioclase, 30 % clinopyroxene, 20 % CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 116 orthopyroxene and 10 % olivine assemblage. Partition coefficients are taken from Fujimaki et al. (1984) and Schwandt & McKay (1998). Varying the modal abundance of the minerals by up to 20 % has no profound influence on the REE models, but does control whether the final magma has a basaltic composition. This is especially noticeable for Al2O3, which is controlled solely by plagioclase fractionation. At greater than 60 % fractional crystallisation, the Al 2O3 concentration increases unrealistically to over 20 %

(assuming An 80), although lower anorthite/albite ratios limit this increase. More complex models where the modal mineralogy varies during fractional crystallisation have not been considered. The models have been assigned a starting composition based on local igneous compositions, for example a melt with the composition of the least evolved magnesian basalt (gc170697) modified by addition of 10 % granodiorite (gc410697). They are then numerically evolved using Rayleigh fractional crystallisation with the aforementioned partition coefficients. The aim is to produce REE patterns similar to the most evolved Coonterunah-Warrawoona tholeiitic basalts (not eccentric) within a plausible percentage of mass fractionation. The local granodiorite is used as the additive in the models, although it did not crystallise until after the eruption of the Coonterunah Group. It is the most likely crustal contaminant for the Warrawoona Group, and its composition probably reflects that of the putative pre-Coonterunah basement as it is a typical Archaean TTG (cf: Martin, 1994). Possible crustal contaminants are further discussed in Chapter 5. Three models are presented in Figure 4.9. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 117

A. Fractional crystallisation only 100 95%

90%

70% E T I

R 50% D N

O 10 H gc170697 C

/

K C O R

1 La Ce Nd Sm Eu Dy Er Yb Lu

100 B. 10% assimilation and fractional crystallisation

70%

50% E T I

R 0% D

N 10 O

H gc170697 C

/

K C O R

1 La Ce Nd Sm Eu Dy Er Yb Lu C. 20% assimilation and fractional crystallisation 100 70%

50%

E 0% T I R D N

O 10

H gc170697 C

/

K C O R

1 La Ce Nd Sm Eu Dy Er Yb Lu

Figure 4.8: Fractional crystallisation models with a starting composition of the least evolved basalt (gc170697) and assimilation of local granodiorite (gc410697) An evolved tholeiitic (gc290697) and eccentric LREE-enriched (gc430697) basalt are plotted (dashed lines) to assess the ability of the models. (A) No crustal assimilation with 50, 70, 90 and 95 % fractional crystallisation. (B) 10 % assimilation with 0, 50 and 70 % fractional crystallisation. (C) 20 % assimilation with 0, 50 and 70 % fractional crystallisation CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 118

The first model shows straightforward fractional crystallisation of a melt with a magnesian basalt (gc170697) composition and no assimilation. Two problems arise from this model: the patterns produced are flatter than those of the Pilbara basalts and the amount of fractionation required to explain the most evolved basalts is unrealistically high (~90 %). Therefore, given the assumptions used here, the most evolved tholeiitic basalts cannot be simply explained by varying amounts of fractional crystallisation of the least evolved magnesian basalts. The other two models show 10 % and 20 % assimilation, respectively, of the local granodiorite (gc410697) into a melt with a magnesian basalt composition (gc170697) and then fractional crystallisation. The assimilation models produce patterns with a shape consistent with the evolved Pilbara basalts, except for some MREE depletion. Adding hornblende to the fractionating assemblage would only enhance this depletion as, for basaltic melts, the MREE are more compatible in hornblende than the LREE and HREE (Green & Pearson, 1985). The amount of mass fractionation required to attain the LREE and HREE concentrations of the most evolved basalt is much lower than for the assimilation-free model, although still high (up to 75 %). In addition, the basalts have higher concentrations of many incompatible elements than the granodiorite, which suggests that the actual contaminants were more enriched in incompatible elements than the granodiorite or that there was selective assimilation of certain elements. However, despite the number of assumptions required for the models, particularly contaminant composition, mineral modes and partition coefficients, it is apparent that the crustal assimilation-fractional crystallisation models match the observed compositions of the Coonterunah and Warrawoona basalts far better than the simple fractionation model.

4.7.4 Eccentric basalts Compared with most Pilgangoora basalts, the eccentric basalts are more enriched in Zr, Nb, La, Ce, Pr, Ta, Th, U and P 2O5 than in Y and the M-HREE. This may reflect addition of contaminants with distinctive compositions. However, such enrichment may also reflect partial melting of a mantle source with a similar bulk composition but different mineralogy to the main basalt suite. For instance, the presence of garnet during melting produces Y- and HREE-depleted magmas (Herzberg et al., 1988; Ohtani et al., 1989; McCuaig et al., 1994). The eccentric basalts, however, have Y and HREE abundances comparable to the main group and so were probably enriched in highly CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 119 incompatible elements rather than depleted in Y and the HREE. Hence, differences in mineral phases in the basalt sources were probably not responsible for the eccentric basalts. Instead, the eccentric basalts probably assimilated material that had been relatively depleted in Y and the M-HREE compared to the other incompatible elements. A contaminant derived from a garnet-bearing source is most likely. One option that is not supported is the assimilation of Coonterunah felsic volcanic magmas, as these have less pronounced LREE enrichment than the Carlindi granitoids (see Chapter 5). The divergence of Y and the M-HREE from the trend defined by the main basalt suite suggests that this contaminant probably affected many of the basalts, and was probably an important component of the basement. In other words, the Pilgangoora basalts have probably assimilated two distinct contaminants, one resembling granodiorite and the other derived from a garnet-rich source.

4.7.5 Contamination estimate Assuming that the Pilgangoora basalts were produced by crustal assimilation and fractional crystallisation, then the amount of contamination can be estimated from mixing curves of HFSE, given that the elements have similar partition coefficients during typical basaltic petrogenesis and that the elemental ratio in the contaminant differs markedly from that in the mantle source. This approach works because the elements should not fractionate during source melting and basalt crystallisation (Sylvester et al., 1997), and hence, any variations should be caused by either crustal addition or source heterogeneity, provided that they are immobile during post-magmatic metamorphism and alteration. As shown previously, post-magmatic mobility of HFSE was minor in the Coonterunah and Warrawoona basalts. Two elements particularly useful for this approach are Nb and U; they have similar partition coefficients at basaltic compositions and much greater Nb/U ratios in the mantle than the crust (Hofmann et al., 1986; Sims & DePaolo, 1997). As already shown, the Pilgangoora basalts have generally retained their magmatic Nb and U compositions (Fig. 4.5). In Figure 4.10, Nb/U is plotted against Nb/Th and La/Sm as both Th and the LREE are enriched in continental crust relative to the mantle (Rudnick & Fountain, 1995), and thus mixing produces a wide range of values. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 120

10

8 gc200697

1 % 6 h T / b 10 % N 4 granodiorite (gc410697)

2

25 %

granodiorite 6 (gc410697)

5 n ] 4 m

S 25 % /

a 3 L [

1 % 2 gc200697

1 10 %

0 0 10 20 30 40 50 Nb/U Figure 4.10: Mixing models of least contaminated Pilgangoora basalt (gc200697) and local granodiorite (gc410697) show estimates of the percentage contamination. La and Sm have been normalised with C1 chondrite (Sun & McDonough, 1989). Symbols as for previous figures.

The basalt with the highest Nb/U (gc200697 = 40.6) is interpreted to be the least contaminated, which is borne out by its high Nb/Th and low LREE values. Assuming that the mantle source also had a Nb/U ratio of 40.6 and the local granodiorite is the contaminant (Nb/U = 4.3), it can be shown that the maximum level of crustal contamination for the main suite of basalts was ~10 %. As expected, the most contaminated sample (gc290697) has enriched trace element patterns, but the degree of contamination does not necessarily mirror the evolved status of the basalts. For example, all of the magnesian basalts are more contaminated than many of the tholeiitic basalts. In other words, some tholeiitic basalts are less contaminated, but more fractionated, than the magnesian basalts. The eccentric basalts are the most contaminated, up to ~25 %, under the above assumptions. Alternatively, the eccentric basalts may have assimilated less of a contaminant with smaller Nb/U ratios. CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 121

Thermal considerations suggest that crustal assimilation would have been greatest in primitive magmas, with up to 30 % and 10 % assimilation possible for komatiites and tholeiites, respectively (Huppert & Sparks, 1985). Contamination would also have been enhanced at shallower crustal levels and for contaminants with low fusion temperatures (Reiners et al., 1995). Therefore, the estimates of ~10 % crustal assimilation for the main suite of Pilgangoora basalts seem reasonable. The least contaminated and least evolved basalts (for instance, magnesian basalt 750798, Nb/U = 40.4) can provide constraints on the composition of the mantle source. These have flat to LREE depleted, approximately 10 times chondritic REE patterns which are similar to modern ocean-floor basalts (Sun & McDonough, 1989). However, they are still enriched in K, Rb, Ba, Th and U and depleted in Ta, Nb, P, Nd, Eu, Ti, Y and Yb relative to average N-MORB, the characteristics expected from contamination. Thus, these differences are probably due to the small degrees of contamination even in these basalts. Without this contamination, the N-MORB normalised patterns of the Coonterunah and Warrawoona basalts would most likely be similar to modern ocean- floor basalts derived from a depleted mantle.

4.8 DISCUSSION The Pilgangoora basalts have crustal signatures compared to N-MORB which may be due to subduction-related metasomatism of the mantle source or crustal assimilation and fractional crystallisation of magmas derived from a depleted mantle source. Distinguishing between these two models using elemental abundances is not possible, but geological and geochemical arguments have shown the viability of assimilation-fractional crystallisation models. Unfortunately, the Coonterunah Group has not been identified elsewhere in the Pilbara and so there are no comparative analyses. However, a subduction-related setting has been proposed for the Warrawoona Group (Barley, 1980, 1993; Barley et al., 1984, 1998). Hence, elemental abundances from all early Archaean Pilbara basalts are reviewed here to determine whether they are significantly different from the Pilgangoora basalts. It should be noted, however, that the review here is limited to elemental abundances of basalts and a few ultramafic rocks, and that these results are integrated with data from Pilbara felsic volcanics and granitoids and Nd-Hf isotope systematics in the following chapters. Prior to this review, the CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 122 previous geochemical survey in the Pilgangoora Belt is compared with the data presented here.

4.8.1 Previous Pilgangoora survey The only prior geochemica l survey in the Pilgangoora Belt involved 53 whole- rock analyses from a stratigraphic traverse ~4 km west of Strelley Pool (Glikson et al., 1986b). Of these, 34 were from the Coonterunah Group and 19 from the Warrawoona Group. Major element abundances show that these rocks have variably altered magmatic compositions, but are dominantly metamorphosed tholeiitic and magnesian basalts. Most trace element abundances were rather imprecise and so provide few additional petrogenetic constraints. However, Zr and Y abundances were well above detection limits and show no evidence of post-magmatic alteration, and so probably reflect magmatic compositions. Systematic variations of Zr, Y, TiO 2 and P2O5 abundance with stratigraphic height confirm temporal trends of Coonterunah basalt composition, and may even define some smaller-scale fluctuations. Moreover, similar trends can be shown for the Warrawoona basalts, although stratigraphic heights were poorly defined by Glikson et al. (1986b) and so these trends should be treated cautiously. These analyses also confirm the existence of eccentric basalts, as defined by enrichment of Zr and P 2O5 relative to Y and TiO 2.

4.8.2 Other Pilbara basalts Previous Pilbara-wide geochemical studies of greenstone belts (Glikson & Hickman, 1981a, b; Glikson et al., 1986a, b) have rather imprecise trace element abundances, as seen above, and thus provide little basis for rigorous interpretation. They are not considered further. Twenty-eight basalts from the North Pole Dome have crustal signatures compared with N-MORB: that is, Ta-Nb depletion associated with Th, U and LREE enrichment (Green et al., in prep.). They also have pronounced negative P anomalies and thus are very similar to the Pilgangoora basalts. Hence, the North Pole basalts probably formed by equivalent processes as the Pilgangoora basalts. Importantly, the least evolved North Pole basalts are more depleted in LREE than those from the Pilgangoora Belt, perhaps providing a better constraint on the composition of the mantle source. The North Pole basalts also display systematic compositional variations with CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 123 stratigraphic height, but these are the reverse to those in the Coonterunah succession, with more evolved basalts at the base becoming progressively less evolved towards the top. This trend continues for half of the succession and is then repeated. Four plagioclase-phyric basalts are restricted to the base of the succession and have compositions similar to the Pilgangoora eccentric basalts, whereas the remaining North Pole basalts have similar Y and M-HREE dispersions as the main Pilgangoora basalt suite. Interestingly, one of the eccentric Coonterunah basalts is also plagioclase-phyric, but it is difficult to see how plagioclase accumulation could be associated with Zr, Nb,

La, Ce, Pr, Ta, Th, U and P 2O5 enrichment relative to Y and the M-HREE. The North Pole basalts (Euro Basalt) above the Strelley Pool Chert have equivalent compositions to those from the same stratigraphic positions in the Pilgangoora Belt, supporting the correlation between belts. In the Marble Bar Belt, 14 samples from the Pilbara-wide surveys (Glikson & Hickman, 1981b) were analysed for REE by ID-MS (Gruau et al., 1987), and 3 of these were re-analysed by spark source-mass spectrometry (SSMS, Jochum et al., 1991). The SSMS results verify the poor precision of many trace element analyses from the Pilbara- wide surveys. The results show that the samples have retained magmatic compositions of many HFSE and so valid interpretations can be made from these elements. Five

Al2O3-depleted magnesian rocks were analysed, including three ultramafic rocks (two putative komatiites, one pyroxenite), and so the Marble Bar samples represent a broader compositional range than the Pilgangoora data. The Marble Bar samples do not display any systematic stratigraphic trends or have eccentric compositions. The Marble Bar basalts have similar REE compositions to the Pilgangoora basalts: generally enriched in LREE relative to HREE, and the least evolved basalts have lower REE abundances with slight LREE depletion. The most enriched Marble Bar basalts are similar to the most enriched non-eccentric Pilgangoora basalts, but the least evolved Marble Bar basalts are Al2O3-depleted which is not mirrored by the Pilgangoora basalts. The three ultramafic samples have very depleted LREE abundances, but are collinear with the Marble Bar basalts on many HFSE variation diagrams (Zr, Y, REE). Moreover, N-MORB-normalised trace element patterns of the three SSMS samples, including the pyroxenite, are remarkably similar with Nb-P depletion associated with Th, U and LREE enrichment. Hence, there is no geochemical evidence to suggest that the

Marble Bar samples were not comagmatic. Aside for the Al 2O3 depletion in five CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 124 samples, the Marble Bar basalts and ultramafics have remarkably similar compositions to the Pilgangoora basalts, and probably formed by equivalent processes. Indeed, assimilation of older crust is supported by their Nd-isotope compositions (see Chapter 6). Importantly, the ultramafic samples may provide better estimates of the mantle source composition. A number of samples, including some basalts, have been analysed from the Kelly Belt and McPhee Dome (Barley, 1980; Barley et al., 1984, 1998). Only a few HFSE (Y, Zr, REE) were determined for the Kelly basalts, whereas a more comprehensive suite of elements was obtained from the McPhee basalts. Each succession defines a discrete linear trend on some HFSE variation diagrams (Zr, LREE versus Y, HREE) which mirrors the eccentric deviations of the Pilgangoora basalts (e.g., the high Zr/Y samples of Barley et al., 1998). Interestingly, two of the eccentric McPhee Dome basalts were originally described as plagioclase-phyric (Barley et al., 1984), although they were subsequently reclassified as andesites (Barley et al., 1998). Nevertheless, this further suggests a correlation between eccentric compositions and the presence of plagioclase phenocrysts. The linear trends on HFSE variation diagrams indicate that the Kelly and McPhee basalts have retained their magmatic HFSE abundances. No stratigraphic trends could be established. The basalts from both successions are highly enriched in LREE relative to the HREE, and the McPhee basalts have N-MORB-normalised trace element patterns with Nb-P depletion associated with Th, U and LREE enrichment. Thus, they are also remarkably similar to the Pilgangoora basalts. Importantly, these successions have been interpreted to be subduction-related due to their compositional similarities with modern Andean-type convergent margins (Barley, 1980, 1993; Barley et al., 1984, 1998). However, crustal contamination alone can account for these compositions, and so this model is not necessarily supported. Furthermore, even if a ~3434 Ma magmatic ages is accepted for both the Kelly Belt and McPhee Formation (see Chapter 3), then there are still xenocrystic zircons in both successions (Barley et al., 1998; Nelson, 2000), consistent with crustal assimilation. In summary, all the North Pole, Marble Bar, Kelly and McPhee basalts have compositions remarkably similar to the Pilgangoora basalts, with crustal signatures relative to N-MORB. Thus, it is likely that they all formed by similar processes. Moreover, the least-evolved samples from the North Pole Dome and Marble Bar Belt suggest that the mantle source for the Pilbara basalts was extremely depleted in CHAPTER 4: ELEMENTAL GEOCHEMISTRY – MAFIC ROCKS 125 incompatible elements, particularly LREE. Furthermore, eccentric basalts are present in many Pilbara successions and may be associated with plagioclase accumulation.

4.9 SUMMARY The Coonterunah and Warrawoona basalts are separ ated by an unconformity and were erupted >60 million years apart. Compositional variations within each succession are consistent with straightforward magmatic processes, and the basalts have crustal signatures compared to N-MORB. Marked geochemical similarities between the two suites imply that they formed by similar petrogenetic processes. Geological considerations indicate that the Warrawoona basalts were erupted onto continental basement, and thus they may have assimilated continental crust. It follows that the Coonterunah basalts probably formed likewise. Trace element modelling shows that the range of basalt compositions can be easily explained by small degrees of crustal assimilation and then fractional crystallisation. The eccentric basalts have compositions consistent with assimilation of a distinct contaminant, and this contaminant probably also affected the other basalts. Estimates of contamination suggest a maximum of ~10 % for the main basalt suite. The least contaminated Pilgangoora basalts have trace element compositions similar to modern ocean-floor basalts derived from depleted mantle. Therefore, the Coonterunah and Warrawoona basalts are interpreted to have formed by eruption from a depleted mantle source through continental crust, and not by subduction-related mantle enrichment. Basalts from other Pilbara greenstone successions have remarkably similar compositions, and so probably formed likewise. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 126

Chapter 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS

5.1 INTRODUCTION Whole-rock elemental abundances of 10 Coonterunah felsic volcanics, 16 Carlindi granitoids and 3 xenoliths were determined. These samples represent at least four discrete magmatic events: ~3517 Ma Coonterunah volcanism, ~3480 Ma Carlindi granodiorites and ~3468 and 2925 Ma Carlindi granites. The main objectives in this chapter are to describe and compare their compositions, determine their relationships and constrain their petrogenesis. Of special interest is establishing whether the Carlindi granitoids represent intrusive relatives of the Coonterunah volcanics, although they formed at different times. A further aim is to determine whether any of the range of felsic rocks could have represented the contaminants of the Pilgangoora basalts. Previous studies of Pilbara felsic volcanic geochemistry concentrated on craton- wide surveys, but generally their trace element data are limited and not very robust, with many elements below or near detection limits (Davy & Lewis, 1981, 1986; Glikson & Hickman, 1981a, b; Glikson et al., 1986a, b). A subset of the above samples from various greenstone belts and stratigraphic units have yielded more precise Rb, Sr and REE abundances, along with limited Sr-Nd isotopic data (Jahn et al., 1981). Such studies highlight the problem of discriminating silicified mafic volcanics from felsic volcanics. For example, three supposedly felsic samples from the North Shaw Belt were re-interpreted, after precise analysis, as intensely altered pillowed basalts (Jahn et al., 1981). Studies of specific stratigraphic units have concentrated on the ~3430 Ma felsic volcanics from the Kelly Belt and McPhee Dome (Barley, 1980; Barley et al., 1984, 1998; Barley & Pickard, 1999) and the Panorama Formation from the North Pole Dome (Cullers et al., 1993). Geochemical studies on Pilbara granitoids have been performed on the Mount Edgar (Davy & Lewis, 1981, 1986; Collins, 1993), Shaw (Bickle et al., 1983, 1989, 1993) and Corunna Downs Granitoid Complexes (Barley, 1980; Davy, 1988; Barley and Pickard, 1999) and some west Pilbara granitoids (Smith et al., 1998; Smithies & Champion, 1999).

5.2 WHOLE-ROCK GEOCHEMISTRY The analytical methods used to determine the whole-rock compositions were the same as those used for the basalts and are detailed in Chapter 4. In summary, major CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 127 elements were obtained by fusion-disc XRF, trace elements minus Zr by ICP-MS and Zr by pressed-powder XRF. The Coonterunah felsic volcanics have experienced the same metamorphic conditions as the surrounding basalts (upper greenschist- to lowermost amphibolite- facies), although felsic mineral assemblages are not overly sensitive to these changes. Chlorite- and feldspar-bearing breccias and associated quartz veins are common in the felsic volcanics and indicate that they experienced significant hydrothermal fluxes. Some of the analysed samples may have been volcaniclastic, as primary fabrics have been overprinted by alteration and metamorphism in some instances. Such samples could have contained mixtures from numerous volcanic events, possibly causing spurious geochemical trends. These issues aside, lavas have been preferentially sampled and the least altered and most homogenous samples have been analysed. Hence, the geochemical sampling was the best possible, although caution in interpretation is still warranted. The Carlindi Granitoid Complex is generally undeformed, except in the northwest of the mapping area where tectonometamorphic fabrics have been developed and subsequently folded (Chapter 2). The samples cover a large area of the complex, but were obtained prior to resolving zircon U-Pb ages. As a result, no preference was given to sampling the different age suites as they could not be differentiated based on field criteria. Indeed, it is still unclear which age category many of the samples represent. The Carlindi granitoids show some replacement of feldspar by fine-grained white mica (sericite, paragonite) and of mafic minerals by chlorite. Alteration was most intense for samples proximal to the unconformity or intruding the Coonterunah Group, where >90 % of the feldspar has been replaced and chlorite is relatively common. These samples may represent the upper levels of the intrusions. The Carlindi xenoliths have melting textures and were intensely deformed. Their degree of hydrothermal alteration is unclear, but zircon ages suggest they experienced many significant events. Detailed lithological descriptions are presented in Chapter 2 and 3.

5.2.1 Major elements Major element Harker diagrams show a wide compositional range for both the Coonterunah and Carlindi suites (Fig. 5.1). In general, silica abundances are between 65 and 76 wt%, although six samples are outside this range. Three felsic volcanic samples have silica abundances that are greater than those of typical magmatic rocks (up to CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 128

82 wt%; gc330697, gc350697, gc440697; Le Maitre et al., 1989), and this probably reflects pervasive silica alteration. Assuming that TiO 2 was immobile, as shown for the enclosing basalts, then an increase of up to 10 wt% silica is likely. Three samples have lower silica abundances than the main sample suite (070698, 210697, 470798), and they may represent primitive compositions. Loss-on-ignition for all felsic samples is between 1 and 3 wt%.

TiO2 MnO Na2O 0.20 1.20 5.0

4.0 0.80 3.0 0.10

2.0 0.40

1.0

K2O Al2O3 MgO 4.0 15.0

4.0 3.0

10.0 2.0 2.0

5.0 1.0

Fe2O3 CaO P2O5

9.0 6.0 0.20 6.0

3.0 0.10 3.0

50 60 70 80 50 60 70 80 50 60 70 80 SiO2 SiO2 SiO2

Figure 5.1: Major element Harker diagrams for Coonterunah felsic volcanics (solid diamonds) and Carlindi intrusives (triangles; granodiorite - solid, granite - open, anomalous - cross, xenolith – circled). Dry wt%.

The Carlindi granitoids define linear trends for most elements, and the Coonterunah felsic volcanics generally follow sub-parallel trends but with greater silica abundances. Major element abundances in any felsic rock, however, are typically very similar as they are controlled by the equilibrium between quartz, feldspar and melt at low temperatures (Tuttle & Bowen, 1958; Chappell et al., 1998). Most of the granitoid samples have probably not been significantly altered, whereas the felsic volcanics have CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 129

been chemically modified. For both suites, TiO 2, Al2O3, Fe2O3, MnO, MgO, CaO and

P2O5 were compatible and SiO2 and K2O were incompatible. There is considerable scatter for Na2O, much of it due to the felsic volcanic samples, but for the granitoids

Na2O was also incompatible. Sample 210698 plots off the linear trend and has lower

TiO2 and P2O5 and higher MgO contents at equivalent silica abundances than the other granitoids.

Q An

Coonterunah T basalts

granite 210698 1 2 3 070698 b a tonalite 190698 A c granodiorite

4 5 trondjhemite granite 150698 A P Ab Or Figure 5.2: Modal quartz-alkali feldspar-plagioclase (QAP; Streickeisen, 1976; Le Maitre et al., 1989) and feldspar ternary classification (Barker, 1979) for Carlindi intrusives. Fields in QAP diagram are: 1- monzogranite, 2-granodiorite, 3-tonalite, 4-quartz monzodiorite, 5- quartz diorite. Differentiation series: T=tholeiitic, A=alkaline, a=low- K, b=intermediate-K, c= high-K calc-alkaline (Lameyre & Bowden, 1982). Mineral abundances as CIPW norms. Symbols as for Figure 5.1.

Field names have been retained for the felsic volcanics because post-magmatic alteration was too great to rely on their present composition as a guide for nomenclature. This is not the case for the granitoids, and thus, major element abundances have been used to classify them. The quartz-alkali feldspar-plagioclase (QAP) diagram (Streckeisen, 1976) uses the same criteria as for field identification, except that normative compositions have been plotted to avoid misidentification of sericitised feldspar (Fig. 5.2). The Carlindi samples are predominantly granodiorite, with three monzogranites (050697, 090698, 150698), one quartz monzodiorite (210698) and one tonalite (070698). The normative feldspar ternary diagram (Fig. 5.2; Barker, 1979) further classifies these rocks as mainly trondhjemites with five granites (gc400697, 050698, 090698, 100698, 150698), one tonalite (080698) and one granodiorite CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 130

(360698). Three samples (210698, 070698, 190698) have anomalously high anorthite contents, perhaps due to loss of Na2O during alteration. One granite sample (150698) has much greater K-feldspar than the other granites. Thus, the Carlindi samples can be subdivided into trondhjemites (10 samples including the granodiorite), granites (5 samples) and anomalous (4 samples including the tonalite). Compositional differences are also mirrored by age variations, with the trondhjemites older than the granites, although this is based on only a few constrained ages.

5.2.2 Trace elements The use of trace elements to study felsic igneous rocks is more complex than for basalts because they are commonly concentrated in many common accessory minerals (cf. Cr in basalts). Hence, their abundances are typically controlled by the physicochemical parameters that determine the mineral assemblage. For example, apatite solubility increases with increasing Al 2O3 contents (Wolf & London, 1994). Trace element abundances have been plotted on variation diagrams to determine the extent of post-magmatic alteration and to show relative element compatibility for each suite (Fig 5.3). ICP-MS abundances were generally used, but since most of the felsic rocks were analysed at Memorial University for a limited suite of elements (see Chapter 4.5), pressed-powder XRF abundances were used for V, Cu, Ga, Zn and Cr. In the variation diagrams, Zr was not used as the comparator because two felsic volcanic samples have very high concentrations and cause scaling problems (Table 5.1). Instead, Ta was used as the comparator because it is typically immobile during post-magmatic processes (cf: Fig 4.5) and has similar abundance ranges for the Coonterunah and Carlindi suites. It has a strong affinity for Ti-bearing minerals (rutile, titanite) and so its compatibility may also be controlled by accessory minerals. However, to determine whether the trace elements were immobile during post-magmatic processes and whether each rock suite behaved similarly, Ta is a suitable comparator. Samples not included on the diagrams are detailed in Table 5.1.

Table 5.1: Anomalous samples removed from trace element variation diagrams (Fig. 5.3). CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 131

Sample # Conc. Sample # Conc. (ppm) (ppm) Y gc470697 (Coont.) 358 Zr gc470697 (Coont.) 786 gc490697 (Coont.) 179 gc490697 (Coont.) 572 Nb gc470697 (Coont.) 56.0 Mo gc400697 (gran.) 5.56

Variation diagrams with HFSE (Y, Zr, Nb, REE) are more variable for the felsic suites (Fig. 5.3) than the well-defined linear trends of the Pilgangoora basalts (Chapter 4). However, the linear trend between Nb and Ta show that these elements were not significantly modified by post-magmatic processes. Thus, poor correlations between HFSE and other elements must be due to magmatic processes or post-depositional mobility of the latter. Correlations with Ta are summarised in Table 5.2, but cannot be used as a general index of fractionation because correlations vary between suites or are poorly defined. For example, the correlation between TiO 2, which was probably incompatible for all suites (Fig. 5.1), and Ta is positive for the trondhjemites, negative for the granites and neutral for the felsic volcanics. In other words, Ta was probably compatible for the trondhjemites, incompatible for the granites and neutral for the volcanics. The Coonterunah felsic volcanics define linear trends for many elements, particularly the HFSE, and demonstrate that abundances of these important elements have not been significantly altered. Only V and Sr have negative correlation and Ga, Cs and W have neutral correlations. The Sr trend has a pronounced inflection at Ta = ~0.8 ppm. The Carlindi trondhjemites and granites differ markedly from the felsic volcanics, with fewer poor correlations, and thus less post-magmatic modification. For Y, Zr and the HREE, the felsic volcanics have strong positive correlations, but the granitoids have flatter trends with possible inflections at Ta = ~1 ppm marking the change from a positive trondhjemite to a negative granite trend (see trends on Fig. 5.3). CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 132

Li Zn Ag (ppm) 150(ppm) 0.15 (ppm)

100 100 0.10

50 50 0.05

Be Ga Cd 4.0 40 0.06 (ppm) (ppm) (ppm)

3.0 30 0.04 2.0 20

0.02 1.0 10

Sc 400 Rb Sn 60 (ppm) (ppm) 4.0(ppm) 300 40 3.0 200 2.0

20 100 1.0

V Sr 0.8 Sb 200(ppm) (ppm) (ppm)

200 0.6

0.4 100 400 0.2

Cr Y Cs (ppm) (ppm) 8.0(ppm) 600 80 6.0 400 4.0 40 200 2.0

15.0 Co Zr Ba (ppm) (ppm) (ppm)

400 800 10.0

5.0 200 400

Ni 40 Nb La (ppm) (ppm) (ppm)

20.0 30 40

20

10.0 20 10

Cu M o Ce (ppm) (ppm) (ppm) 60 1.50 80

40 1.00

40 20 0.50

0 1.0 Ta 2.0 3.0 1.0 Ta 2.0 3.0 1.0 Ta 2.0 3.0 (ppm) (ppm) (ppm)

Figure 5.3: continued overleaf..... CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 133

Pr Dy 15.0 3.0 (ppm) (ppm) W 20.0 (ppm) 10.0 2.0

10.0 5.0 1.0

Ho 80 Nd 6.0 Tl (ppm) (ppm) (ppm)

60 2.0 4.0

40 1.0 2.0 20

Sm Er Pb 15.0 20.0 (ppm) (ppm) (ppm) 30

10.0 20 10.0 5.0 10

Eu Tm Bi (ppm) 1.50(ppm) (ppm)

4.0 0.40 1.00

2.0 0.20 0.50

Tb Yb Th 15.0 4.0 (ppm) (ppm) 20.0 (ppm)

3.0 10.0

2.0 10.0 5.0 1.0

Gd Lu U (ppm) (ppm) (ppm) 6.0 20 2.0

4.0

10 1.0 2.0

0 1.0 Ta 2.0 3.0 1.0 Ta 2.0 3.0 1.0 Ta 2.0 3.0 (ppm) (ppm) (ppm)

Figure 5.3: Trace element variation diagrams of Coonterunah felsic volcanics (solid diamonds) and Carlindi felsic intrusives (triangles; trondhjemite - solid, granite - open, others - cross, xenolith - circle) show intensity of post-magmatic alteration and magmatic compatibility of unmodified elements. Some highly anomalous samples have been deleted, see text for details. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 134

Table 5.2: Summary of trace element correlations versus Ta for Coonterunah felsic volcanics and Carlindi intrusives (Fig. 5.3). Correlation Poor Positive Negative Neutral Inflection Coonterunah Ti, Li, Cr, Be, Sc, Ni, Rb, Y, V Ga, Cs, W Sr felsic Cu, Zn, Zr, Nb, Mo, Ag, volcanics Cd, Sn, La, Ce, Pr, Nd, Sm, Sb, Ba, Tl, Eu, Tb, Gd, Dy, Pb, Bi, U Ho, Er, Tm, Yb, Lu, Th Carlindi Li, Be, Cr, Ti, V, Rb, Nb, Ag, Sc, Eu Ni, Ga, Y, Sr, Zr, Ba, trondhjemite Cu, Zn, Cd, Sn, Sb, Cs, Tl, Tb, Gd, Dy, La, Ce, Pr, and granite Mo, W, Bi Pb, Th, U Ho, Er, Tm, Nd, Sm Yb, Lu

5.2.3 Normalised diagrams Source compositions and element partitioning within felsic systems are different than for basalts, and so the samples are normalised using the composition of average upper continental crust (Taylor & McLennan, 1985; Nb from Barth et al., 2000) and a suitable element compatibility order (Sylvester, 1994). Zn and Cu have not been included on these diagrams as they were below detection limits for many samples and probably mobile (Fig. 5.3). In general, the crustal normalised variation diagrams should have a negative slope as the elements are ordered with increasing compatibility to the right. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 135

10 10 A lower Coonterunah S E - Carlindi trondhjemite A M P L

1.0 E 1.0

/

C R U S T gc330697 0.1 0.1 gc470697 270898 gc490697 gc240697

050798 260898 B - eastern Coonterunah F - Carlindi trondhjemite S A 1.0 M 1.0 P L E

/

C R

340798 U 0.1 S 0.1 140698 500798 T 190898 470798 360698 (gd) 450798 C - sub-unconformity Coonterunah G - Carlindi granite S A 1.0 M 1.0 P L E

/

C R

U 050698 gc350697 0.1 S 0.1 gc400697 T gc440697 100698 gc060697 090698 (xeno) gc430697 150698 D - Carlindi trondhjemite H - Carlindi anomalous S A

1.0 M 1.0 P L E

/

C gc420697 R 080698 (ton, xeno) U

0.1 S 0.1

gc220697 T 210698*

gc230697 190698*

gc410697 070698 (xeno) 0.01 0.01 Th Rb Y La Ce Ba Zr Ti Nb P Sr Th Rb Y La Ce Ba Zr Ti Nb P Sr

Figure 5.4: Upper-continental-crust-normalised variation diagrams for felsic rocks and two Coonterunah basalts. See text for normalising values. gd - granodiorite, ton - tonalite, xeno - xenolith, * - outside feldspar ternary categories.

The Coonterunah felsic volc anics have extremely irregular patterns (Fig. 5.4A- C), although this may be due, at least in part, to post-magmatic alteration (eg., Ba). However, the HFSE have not been significantly altered (Fig. 5.3), and so these irregularities must reflect magmatic processes. Most of the samples have pronounced positive Y, Zr and Nb, negative Rb, Ba and Sr, and variable Ti and P anomalies, and those with the least variable patterns have abundances approximating upper continental crust. Samples have been stratigraphically grouped; lower Coonterunah are the thin, moderately continuous units generally from the central domain (Fig. 5.4A), east- CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 136

Coonterunah are the volcanics in the eastern domain (Fig. 5.4B) and sub-unconformity are two samples from immediately beneath the unconformity (Fig. 5.4C). In general, the pattern shapes are consistent between the groups, although the lower Coonterunah volcanics have much larger anomalies and may represent a distinct suite. One sample (050798) was collected from the deformed western domain and its composition supports it being volcanic and from the lower Coonterunah suite. Similarities between the east- Coonterunah and sub-unconformity felsic volcanics support their stratigraphic equivalence, as suggested by their zircon U-Pb ages (Chapter 3.4.2). Typical Coonterunah and eccentric tholeiitic basalts (gc060697, gc430697) have similar patterns to the felsic volcanics. The Carlindi granitoids (Fig. 5.4D-H) have markedly different patterns than the felsic volcanics, but they are similar to each other, except for two samples (150698, 070698). Elemental abundances are typically less than those in average upper continental crust. With the elements arranged in compatibility order, most of the trondhjemites have positive Sr and negative Th, Rb and Y anomalies relative to La, Ce,

Ba, Zr, Ti, Nb and P (eg., [Th/La] norm < 1). The elements in the middle of the diagram

(La to P) define a slight negative ([La/P] norm > 1), as expected for their relative compatibilities. The granites have positive Ba, Nb and Sr and negative Y anomalies imposed on an overall negative trend. In general, the granites have slightly steeper negative slopes (as defined by the La/P ratio) caused by lower abundances of the more compatible elements, such as P and Ti. Granite 150698 differs markedly from the others with large Rb, Y, Zr, Nb and P abundances relative to the other elements. Xenolith 070698 also differs markedly from the adjacent granitoids, but it is a greenstone relict and predominantly contains chlorite and biotite. Chondrite-normalised REE diagrams (Fig. 5.5; Sun & McDonough, 1989) for the lower Coonterunah felsic volcanics are flat, whereas the other felsic volcanics have modest negative L-MREE slopes with generally flat HREE slopes. Concentrations are generally between 10 and 100 times chondritic and they all have negative Eu anomalies. The felsic volcanics are more enriched than the basalts (gc060697, gc430697), although their patterns are similar. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 137

1000 1000 A - lower Coonterunah E - Carlindi 270898 gc240697

trondhjemite 260898

100 1 100 C

/

E L P M A

gc330697 10 S 10 gc470697 gc490697 050798

B - eastern Coonterunah F - Carlindi 140698 190898

trondhjemite 360698 (gd)

100 1 100 C

/

E L P

340798 M A

10 S 10 500798

470798

450798

C - sub-unconformity Coonterunah G - Carlindi granites 050698 gc400697

100698 100 1 100 090698 (xeno) C

/ 150698 E L P M A

gc350697 10 S 10 gc440697 gc060697 (tholeiitic basalt) gc430697 (eccentric basalt)

080698 (ton, xeno) D - Carlindi gc420697 H - Carlindi 210698* gc220697 anomalous trondhjemite 190698*

gc230697 100 1 100

C 070698* (xeno)

/ gc410697 E L P M A

10 S 10

1 1 La Ce Nd Sm Eu Gd Dy Er Yb Lu La Ce Nd Sm Eu Gd Dy Er Yb Lu

Figure 5.5: C1-chondrite-normalised diagrams for felsic rocks and two Coonterunah basalts. Normalising values from Sun & McDonough (1989). gd - granodiorite, ton - tonalite, xeno - xenolith, * - outside feldspar ternary categories.

In general, the Carlindi granitoids have consistent patterns with steep negative L- MREE and flatter, but still negatively sloping, M-HREE trends. Three samples have profoundly different patterns. Granodiorite 360698 has a steep negative L-MREE and a flat M-HREE trend, with significant enrichment of HREE relative to the trondhjemites. The greater HREE content reflects the presence of ~10 % modal hornblende. Granite 150698 has a steep negative L-MREE and positive M-HREE trend, reflecting the presence of <1 % modal garnet. Xenolith 070698 has a gentle monotonic negative CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 138 slope, consistent with the greenstones. In general, the trondhjemites have positive and the granites have negative Eu anomalies.

5.3 INTERPRETATION It is evident that the whole-rock composit ions of the Coonterunah felsic volcanics have been modified by post-magmatic processes, whereas the Carlindi intrusives have generally retained their magmatic compositions. As with the Coonterunah basalts, the HFSE elements have not been significantly modified, and so interpretations can be confidently based on these elements.

5.3.1 Petrogenesis : Coonterunah felsic volcanics The felsic volcanics were deposited in a basalt-dominated succession and could represent continued differentiation of these basaltic magmas. Alternatively, they may not be petrogenetically linked to the basalts, perhaps representing crustally-derived magmas or small degrees of partial melting of the mantle. The felsic volcanics could also be extrusive expressions of Carlindi-type granitoids, and if so, the petrogenesis of the Carlindi granitoids could provide further constraints on the tectonic setting of the supracrustals. Although many of the differences between the Coonterunah felsic volcanics and the Carlindi granitoids can be dismissed as alteration, the contrasting HFSE abundances indicate that they formed by dissimilar processes. This is clearly shown by the Y, Zr and REE differences in the normalised variation diagrams (Fig. 5.4-5), but is also expressed as separate trends in the trace-element variation diagrams (Fig. 5.3). Such variations indicate contrasting petrogenetic histories and so the felsic volcanics cannot be the extrusive equivalents of Carlindi-type magmatism. Furthermore, Carlindi-type granitoids cannot be restite bodies to the volcanics because the trondhjemites, and particularly the granites, have more evolved compositions than some of the felsic volcanics. Discriminating between late-stage fractionation of basaltic magmas and small volumes of mantle melting is more difficult. In general, the composition of small-volume partial melts may reflect the presence of minor residual high-pressure minerals, such as garnet, which would not be retained during large degrees of mantle melting, as required for basalts. Therefore, felsic volcanics formed by partial melting may have distinctive compositions relative to those formed by late-stage basalt differentiation during low- CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 139 pressure fractional crystallisation. For instance, mantle melting with residual garnet should produce melts with low Y and HREE abundances as these elements are preferentially retained by garnet. However, in the absence of distinct mineral assemblages, discriminating between the two processes may still be almost impossible. Nevertheless, there are no profound trace element variations between Coonterunah basalts and felsic volcanics (Fig. 5.4,5), perhaps indicating that the felsic volcanics were not formed by small-volume partial melting of the mantle. The most tenable model is that the Coonterunah felsic v olcanics represent continued segregation of basaltic magma. This model readily accounts for pattern similarities between the normalised variation diagrams for felsic volcanics and basalts (Fig. 5.4; 5.5) and the higher abundances of incompatible elements in the felsic volcanics. Moreover, many of the observed chemical variations can be produced by straightforward changes to the crystallising assemblage. For instance, during basalt crystallisation SiO 2, Al2O3, Fe2O3 and P2O5 were incompatible and MgO and CaO were compatible (Fig. 4.4), whereas for the felsic volcanics SiO 2, Na2O and K2O were incompatible and TiO 2, Al2O3, Fe2O3, MgO, CaO and P2O5 were compatible. If this model is correct, then the evolving magmas must have crystallised TiO 2-, Al2O3-, Fe2O3- and P2O5-bearing minerals in different modal abundances than for basaltic compositions. Moreover, such minerals must not have been too enriched in Y, Zr or the REE as these remained relatively incompatible over the entire compositional range. These relationships can be shown on variation diagrams with a comparator that maintained constant compatibility, such as SiO 2 or MgO, and is thus an index of fractionation. Abundances of both of these elements have been modified by post-magmatic alteration, but MgO was the least modified and so is used here. Less evolved rocks have higher MgO contents and so evolution trends are from right to left (Fig. 5.6).

At MgO contents of ~4.0 wt% (SiO2 contents of ~55 to 60 wt%) there are marked inflections in the TiO 2, Al2O3 and P2O5 plots, although there are too few samples with MgO abundances between 2.5 and 4.0 wt% to specify the inflection points. Hence,

TiO2, Al2O3 and P2O5 must have been more readily incorporated into the crystallising assemblage when the basaltic melts reached ~4 wt% MgO. Some of these changes probably reflect the crystallisation of new mineral phases, such as ilmenite, sphene or rutile for TiO 2, K-feldspar, plagioclase or mica for Al 2O3 and apatite or xenotime for

P2O5. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 140

TiO2 Fe2O3 Zr (wt%) (wt%) (ppm) 1.5 12.0 600

1.0

6.0 300 0.5

trondhjemite trondhjemite trondhjemite granite granite granite

Al2O3 P2O5 CaO (wt%) (wt%) (wt%) 12.0

0.20 10.0 8.0

0.10 5.0 4.0 trondhjemite trondhjemite granite granite

0.0 4.0 8.0 12.0 0.0 4.0 8.0 12.0 0.0 4.0 8.0 12.0 MgO (wt%) MgO (wt%) MgO (wt%) Figure 5.6: Major element variation diagrams for Coonterunah basalts and felsic volcanics. Arrows indicate general evolution trends. Carlindi trondhjemite and granite field are also shown. Major element oxides recalculated to dry wt%.

Xenotime preferentially incorporates Y, Th and the HREE, and so its crystallisation should result in inflected trends for these elements, too. These are not observed and hence xenotime was probably not the crystallising phosphate phase. Likewise, potassic minerals do not satisfactorily account for the alumina inflection because K2O was incompatible in the felsic volcanic magmas (Fig. 5.1). Since plagioclase was already prevalent in the basaltic magmas (see Chapter 4.7.3), the Al 2O3 inflection may indicate a marked change of plagioclase composition, the onset of greater plagioclase crystallisation or both. A straightforward change of plagioclase composition is unlikely because magmas tend to evolve to more sodic compositions, and thus plagioclase develops a lower alumina content (paired substitution of NaSi with CaAl). However, CaO was compatible at all compositions, and so if the evolving magmas crystallised less pyroxene, as expected, then another Ca-rich mineral, perhaps anorthite, must have become more prevalent. It thus seems likely that the Al 2O3 inflection resulted from increased modal plagioclase. The ubiquitous negative Eu anomaly supports a plagioclase-rich crystallising assemblage.

For Fe2O3, an inflection between 7 and 8 wt% MgO indicates that an Fe-rich mineral, or an Fe-rich member of a solid-solution series, started to crystallise. Since the CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 141

Fe2O3 and TiO 2 inflections do not coincide then this change cannot reflect significant ilmenite crystallisation, although it does not preclude ilmenite from the assemblage when MgO reached ~4 wt%. The iron inflection probably indicates the onset of magnetite crystallisation or a change in pyroxene composition. In general, Zr was incompatible, although there may be an inflection at very low MgO contents. The extremely high Zr concentrations in a couple of samples suggest that conditions were unfavourable for zircon crystallisation until extremely evolved compositions were attained. Under such conditions the preservation of xenocrystic zircon is unlikely because Zr is undersaturated in the melt. This may account for a lack of pre-magmatic zircons derived from the inferred crustal contaminants (see Chapter 3.4). In summary, the general composition of the felsic volcanics can be accounted for by continued fractionation of Coonterunah basaltic magmas through reasonable changes in the crystallising assemblage. The six eccentric Coonterunah basalts are transitional between th e main basalt suite and the felsic volcanics (Fig. 5.6), and it has already been shown that they are enriched in Zr, Nb, La, Ce, Pr, Ta, Th, U and P 2O5 relative to Y and the M-HREE and the main Coonterunah basalts suite (Chapter 4.7.4). Therefore, as the felsic volcanics are evidently products of basalt fractionation, then they may be related to either suite. To discriminate between the two suites, the elements which show the greatest contrasts between them are plotted (Fig. 5.7). Praseodymium (Pr) is used as the comparator because it has not been too greatly enriched in the eccentric basalts, defines the most robust linear trend between both suites and has been unaffected by post-magmatic processes (Fig. 4.5). The diagrams show two oblique trends, each defined by the same samples. The main basalt suite has four related felsic samples (gc330697, gc470697, gc490697, 0500798), whereas the eccentric suite has six samples (gc350697, gc440697, 340798, 450798, 470798, 500798) at more evolved compositions. Significantly, both groups are stratigraphically restricted. The four main suite samples are from the lower felsic units in the central and western domains, and the eccentric felsic units are from immediately beneath the unconformity in the central domain and from the thick felsic succession in the eastern Pilgangoora domain. This further supports the stratigraphic correlation between the felsic volcanics immediately beneath the unconformity with those in the eastern domain. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 142

Ce 16.0 Yb (ppm) (ppm)

80.0 12.0

8.00

40.0

4.00

Tb Y (ppm) 4.00 (ppm) 300.0

3.00

200.0 2.00

100.0 1.00

0.00 5.00 10.0 15.0 20.0 0.00 5.00 10.0 15.0 20.0 Pr (ppm) Pr (ppm)

Figure 5.7: Trace-element variation diagrams showing the contrasting Coonterunah trends. Filled diamond - basalt, open diamond - felsic volcanic, circled - eccentric, dashed line - eccentric, ellipse - Carlindi trondhjemites and granites.

These trends are consistent with the felsic volcanics forming by continued fractionation of the Coonterunah basalts, although such correlations may also be due to mixing of unrelated felsic volcanic magmas with the Coonterunah basalts. Indeed, the two distinct felsic volcanic suites could be the two proposed contaminants of the basalts. However, the chemical relationships used to discriminate between the eccentric and non- eccentric basalts are also the mixing lines for the different felsic volcanic suites, and so reconciling such differences would not be easy. Nevertheless, mixing seems highly unlikely. In general, the felsic volcanics have similar compositions, as evident in their normalised diagrams, and this must reflect similar petrogenetic processes. Their likeness is especially profound when compared with the granitoids. Hence, one trend cannot be the late-stage fractionation of basaltic magma and the other the partial melting of a granitoid. In other words, if the trends are mixing lines then there must be two distinct but similar sources, and the magmas must have evolved by the same processes. Small- volume partial melts of the mantle would be unlikely to be so enriched in incompatible elements (eg., >750 ppm Zr) and the retention of garnet is precluded by the felsic volcanics still having greater abundances of HREE than the least evolved basalts. It will also be shown that none of the Carlindi granitoids has the necessary composition to be a CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 143 restite to the felsic volcanics. Thus, there is no obvious source for one, let alone two partial melts to form highly enriched felsic volcanic magmas. The most tenable model is that the Coonterunah felsic volcanics represent late-stage fractionated melts of crustally contaminated basalts, and that both the main and eccentric basalt suites are represented.

5.3.2 Petrogenesis: Carlindi granitoids The Carlindi granitoids represent at least three magmatic events (~3480, ~3468, ~2925 Ma), but identifying them in the field has proven difficult as discontinuous exposure obscures their contacts. Moreover, they have very similar bulk compositions, but can be divided into trondhjemites and granites based on proportions of normative feldspar (Fig. 5.2), amongst other criteria (Fig. 5.8). The differences between these groups are significant and so they are treated separately. Two xenoliths (070698, 080698) and two samples from immediately adjacent to and within the greenstones (210698, 190698) have less evolved compositions than the trondhjemites or granites and may reflect less fractionated trondhjemitic-granitic magmas or contamination by the greenstones (Fig. 5.1). These anomalous samples are discussed separately.

5.3.2.1 Trondhjemite In the field area, the majority of the exposed Carlind i complex is probably composed of trondhjemite, although field criteria are not especially diagnostic of this. However, radiometric counts from airborne geophysical surveys support the predominance of low-K granitoids, consistent with abundant trondhjemites (Mackey & Richardson, 1997b). Two trondhjemites have been dated: 270898 at ~3484 Ma and gc420697 at ~3479 Ma, with neither sample yielding xenocrystic zircons (270898 = 153188, gc420697 = 95037 in Chapter 3.5). These are the oldest-known Carlindi intrusions, and among the oldest in the Pilbara Carton. They were emplaced ~35 myr after the Coonterunah greenstones, and precede granite intrusion and Duffer volcanism by ~14 myr (Fig. 3.19). Therefore, they played a pivotal role in crustal evolution. Although they were collected from throughout the field area, the trondhjemites have surprisingly limited compositional ranges that define relatively tight clusters on variation diagrams (Fig. 5.1-3). Recalculated to dry weight, they have SiO 2 = 70.2 to total 73.9 wt%, Al2O3 = ~15.4 wt%, FeO /MgO > 3, Na2O = ~4.8 wt% and Na2O/K2O > 1.7. Their restricted and highly evolved compositions indicate that the trondhjemites CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 144 represent entire melts rather than just restite or cumulate fractions, although there are obvious fractionation trends between samples. Moreover, their positive Eu anomalies probably reflect minor accumulation of plagioclase. Elements with highly compatible behaviour include TiO 2, Al2O3, Fe2O3, MgO, CaO and P2O5, whereas Zr, Y, Nb, Ba and

LREE were only slightly compatible. Incompatible elements include SiO 2, Rb and Sr, whereas Na2O, K2O, U, Th and the HREE were essentially neutral. Crystallisation of plagioclase, sphene, zircon, apatite, allanite and xenotime can account for most of the compatible elements. Potassic minerals such as K-feldspar and biotite control the compatibility of Ba (Arth, 1976; Blundy & Wood, 1991), although the neutral behaviour of K2O and incompatibility of Rb suggests such minerals were not overly abundant. The trondhjemites have trace elements abundances that are typically lower than average upper continental crust, with negative Th, Rb and Y and positive Sr anomalies relative to an overall negative slope (Fig. 5.4). They have very steep chondrite- normalised REE patterns (Fig. 5.5), generally with La n > 60, Ybn < 5 and [La/Yb]n > 25. They are also enriched in Cr, with abundances greater than in many of the Pilgangoora basalts, but they are not correspondingly enriched in Ni and V. These samples are chemically classified as trondhjemites, although half of them are slightly more mafic than the defined criteria (FeO (total) + MgO < 3.4 wt%; Barker,

1979). They are relatively enriched in K 2O, their compositions overlap the proposed high-low alumina boundary (15 wt% Al 203 at 70 wt% SiO2) and they are not particularly calcic. They have low-K calc-alkaline affinities (Fig. 5.2; Lameyre & Bowden, 1982), typical of most trondhjemites, but their greater Na/K ratio and lower normative K- feldspar abundances distinguish them from classical calc-alkaline granitoids (Fig. 5.8; Barker & Arth, 1976). They are all mildly peraluminous, as measured by apatite- corrected aluminium saturation indexes (ASI = 1.02 to 1.17; Zen, 1986) and normative corundum contents (0.36 to 2.13 %). The trondhjemites can be classified as I-type granites, although some samples are slightly more aluminous (Chappell & White, 1974; Chappell et al., 1998). Their compositional similarity with the infamous Archaean grey gneiss is striking (Martin, 1994), and they may represent undeformed examples. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 145

K Q

trondhjemite

calk-alkaline

calk-alkaline

trondhjemite

basalt basalt Na Ca Ab Or

Figure 5.8: K-Na-Ca diagram and quartz-albite-orthoclase (Q-Ab-Or) triangle (Barker & Arth, 1976) show significant differences between the Carlindi trondhjemites (solid) and granites (open).

Petrogenetic models for trondhjemites have generally focussed on their Na 2O enrichment and extreme fractionation between the LREE and HREE, and can be divided into three general classes based to the number of required melting generations: i) First generation melt models propose that the mantle-like isotopic compositions of many trondhjemites indicate that they were derived directly from the mantle either by fractional crystallisation of basaltic magmas (Arth et al., 1978; Barker, 1979) or by very low degrees of partial melting (Moorbath, 1975; Stern & Hanson, 1991). However, the Carlindi trondhjemites are too depleted in Y, Zr and the HREE to have formed by continued fractionation of the Pilgangoora basalts. Assuming residual garnet in the melting region, very low degrees of partial melting can account for Y and HREE depletion. However, it cannot explain the depletion of other incompatible elements, such as Zr, or the very large volumes of magma necessary to form the Carlindi complex. Moreover, the Sm-Nd isotopic composition of the Carlindi trondhjemite is not as depleted as the Coonterunah basalt source (Chapter 6). Thus, the Carlindi trondhjemites were probably not derived directly from the mantle. ii) Second generation melt models involve partial melting of basaltic crust with either hornblende, garnet or both in the residue (Hanson & Goldich, 1972; Arth & Hanson, 1972; Arth & Barker, 1976; Jahn et al., 1981; Martin, 1987, 1993; Drummond & Defant, 1990; Rapp et al., 1991; Cullers et al., 1993; Winther, 1996; Springer & Seck, 1997). A large volume of experimental and theoretical work has demonstrated the CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 146 general viability of these models for producing tonalitic and trondhjemitic melts with extremely fractionated REE. However, both the REE and major element abundances of the Carlindi trondhjemites cannot be satisfactorily reproduced by straightforward partial melting of typical Archaean basalts, especially the Coonterunah basalts, according to the experimental constraints. Given a batch melting system, where all the magma equilibrates with and then separates from the source, bulk partition coefficients and melt fractions can be constrained for trace elements. The simplified batch melting system has the following relationship:

CL 1 = CO DR + F(1- DR)

where CL and CO are the concentrations of the trace element in the melt and source, respectively, DR is the bulk partition coefficient of the residue and F is the fraction of melt. The bulk partition coefficient is calculated for the residual minerals when the melt separates from the source, not for the pre-melting composition (Hanson, 1978). This simple model is satisfactory for the discussion here as it most closely reflects the typical experimental method (Winther, 1996; Springer & Seck, 1997). The equation can be represented graphically as the concentration ratio ( CL/CO) versus the fraction of melt ( F) with various bulk coefficient trends ( DR, Fig. 5.9). Therefore, by plotting the trace element ratio (CL/CO) of the trondhjemite (melt) versus the basalt (source) the relationship between F and DR can be constrained. Fields for La and Yb are shown as they represent the end members of REE behaviour. The La field is relatively small and shows that the necessary enrichment of 2.0 to 9.1 times the Coonterunah basalt abundances can only be attained under limited conditions, essentially low DLa and low F. The potential field shrinks as the source becomes more primitive (higher CL/CO ratios) as shown by the large difference between magnesian and tholeiitic basalts. In other words, as the source becomes more primitive the conditions for possible LREE enrichment become correspondingly tighter as greater enrichment is required. In contrast, the Yb field is much larger and shows that the necessary depletion of 2.9 to 11.1 times the Coonterunah abundances can only be attained for high DYb, but for a larger range of melt fractions. Here, the tholeiitic basalts CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 147 have the smaller potential field as greater depletion is required from the more evolved basalts. Therefore, the LREE and HREE provide conflicting constraints, and show that neither very enriched nor very primitive basalts are likely to form a suitable source for the Carlindi trondhjemites. This problem has been illustrated, at least in part, in experiments where melts from evolved basalts have failed to produce suitable HREE depletion (Springer & Seck, 1997). Further insights can be gained by combining this model with established partition coefficients for likely source minerals (Table 5.3). For instance, the bulk DYb is greater than 1.5 for all melt fractions in the above model (Fig. 5.9), and hence, the source must contain garnet, as no other probable source mineral has a sufficiently large partition coefficient. The presence of garnet, however, adversely affects the bulk DLa, which needs to be low to achieve LREE enrichment, but for garnet is relatively high. The opposite is true for plagioclase and pyroxene. Consequently, there must be a balance between these minerals to produce the required bulk partition coefficients. Moreover, the Carlindi trondhjemites are relatively voluminous and so they probably do not represent very small melt fractions. At melt fractions of 5 to 20 %, the necessary LREE enrichments and HREE depletions are not viable for any combination of residual minerals assuming melting of a Coonterunah basalt. Although fractional crystallisation will enhance REE fractionation, without a suitable crystallising assemblage, perhaps including hornblende, the HREE will also become enriched. Indeed, it appears that producing an extremely REE fractionated melt with very low HREE abundances may be extremely difficult, without a partly fractionated precursor. Moreover, producing extreme LREE enrichment from a very primitive melt, such as a komatiite, also seems implausible without an intervening step. Melting an amphibolite with >20 % garnet gets closest to producing the desired REE abundances, consistent with previous studies (Arth & Barker, 1976; Rapp et al., 1991; Winther, 1996; Springer & Seck, 1997).

Table 5.3: REE partition coefficients for important minerals in the melting of garnet-amphibolite mafic crust (compilation from Springer & Seck, 1997). Plag OPx CPx Hbl Ga CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 148

La 0.13 0.03 0.10 0.20 0.37 Ce 0.24 0.03 0.21 0.30 0.53 Nd 0.17 0.04 0.43 0.80 0.81 Sm 0.13 0.04 0.68 1.10 5.50 Eu 2.11 0.04 0.64 1.30 1.37 Tb 0.05 0.08 0.68 2.00 19.60 Yb 0.08 0.11 0.90 1.70 26.00 Lu 0.06 0.11 0.69 1.50 23.50

Experimental studies by Winther (1996) showed the major elemental abundances of melts produced by partially melting average Archaean tholeiitic basalts over a wide range of temperatures, pressures and water contents. The experiments showed that trondhjemites form at higher pressures, lower temperatures and lesser water contents than tonalites. The experimental results have also been used to derive equations relating the abundances of seven major elements to the melting conditions. As most of the Coonterunah tholeiitic basalts have similar compositions to those used in the experiments, they can be realistically substituted. Since the Carlindi trondhjemites are within the compositional limits defined for the equations, and assuming the experimental conditions prevailed, then the melting conditions for each Carlindi sample can be estimated. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 149

100

D = 0.01 not possible

10 magnesian D = 0.03

La D = 0.1 tholeiite O C

D = 1.0 / 1

L D = 1.5 C

D = 5 Yb magnesian D = 10 0.1 tholeiite

D = 26 not possible 0.01 0 0.2 0.4 0.6 0.8 1 Melt Fraction (F)

Figure 5.9: Graphical representation between degree of melting ( F), enrichment-depletion factor ( CL/CO) and bulk partition coefficients ( D) for a batch melting system. Fields show the required La and Yb relationships for the Carlindi trondhjemites to be partial melts of Coonterunah basalts.

30 30 2.0 wt% 4.5 wt% SiO2 25 TiO2 25 CaO FeO K2O Na2O K2O

20 SiO2 Al2O3 20 FeO ) ) r r a a

b b CaO

k 15 k 15

( ( Al2O3

P 10 P 10 Na2O 5 5

TiO2 0 0 650 750 850 950 1050 1150 650 750 850 950 1050 1150 T (°C) T (°C)

Figure 5.10: Range of Carlindi trondhjemite major elements during partial melting of average Archaean tholeiitic basalts at 2.0 and 4.5 wt% H 2O and various temperature and pressure conditions. Based on experimental work by Winther (1996). CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 150

There are seven linear equations with only three variables, and thus it is an overdetermined case with no unique solution. Hence, least squares solutions have been calculated (Table 5.4), with coefficient errors defined for each equation not accounted for in the calculation. Estimated melting conditions for each Carlindi trondhjemite are varied, but within experimental range. Recalculating these conditions back into the equations, however, shows that the estimates produce melts with significantly lower

K2O and higher TiO 2 abundances than the Carlindi trondhjemites. In other words, the

Carlindi trondhjemites are significantly enriched in K 2O and depleted in TiO 2 relative to the experiments. This inconsistency between the experiments and Carlindi trondhjemites can be shown graphically by plotting pressure and temperature fields for fixed water contents (Fig. 5.10). Importantly, these diagrams are only valid for melt compositions equivalent to the Carlindi trondhjemites. The water contents shown cover the range of estimates

(2.0 and 4.5 wt% H2O). In the plots, each element defines a series of parallel lines and so only the upper and lower limits are shown. Critical relationships can be seen between

K2O and the other elements. At H 2O = 2.0 wt%, K2O and SiO2 are mutually exclusive and CaO and FeO intersect the K 2O stability field under vastly different conditions. As the water content increases, the K 2O stability field moves to lower temperatures and progressively separates from the other elements, such that at 4.5 wt% H 2O the K2O stability field does not coincide with any of the other six elements. Hence, melts with compositions similar to the Carlindi trondhjemites are not formed according to the experimental model, as they are too enriched in K 2O. The opposite trend is observed for

TiO2, but the difference is more subtle. Therefore, according to melting experiments of average Archaean tholeiites (Winther, 1996), which cover the compositional range of

Coonterunah basalts, the Carlindi trondhjemites are too enriched in K 2O to be derived from such a source. The general model for forming trondhjemites by partial melting of basaltic crust with either hornblende, garnet or both in the residue is well-founded, and can account for many of the features unique to these rocks. However, the extreme REE fractionation with HREE depletion and enriched K 2O of the Carlindi trondhjemites are difficult to reconcile with such a model. Indeed, it seems impossible that the trondhjemites could have been derived from partial melting of the Coonterunah basalts, or even average Archaean tholeiites, without some further chemical processing. However, the overall CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 151 applicability of this model suggests that the genesis of the Carlindi trondhjemites probably involved such a process as an initial step.

Table 5.4: Estimated temperature (T), pressure (P) and water content

(YH2O) for Carlindi trondhjemites based on least squares solution of equations derived from melting experiments of average Archaean tholeiites (Winther, 1996). Errors = ~10 %.

Sno. T (°C) P (kbar) YH2O Sno. T (°C) P (kbar) YH2O (%) (%) gc220697 948 20.6 3.6 140698 833 16.6 3.2 gc230697 890 20.6 4.4 360698 1019 31.5 5.2 gc240697 937 19.1 3.7 190898 843 12.2 3.6 gc410697 885 15.4 2.7 260898 809 16.5 3.8 gc420697 869 13.2 1.8 270898 869 18.0 4.2

iii) Third generation melts are derived from fusion of second generation sialic crust, and have attracted little attention in the petrogenesis of trondhjemites (Johnston & Wyllie, 1988; Puffer & Volkert, 1991; Collins, 1993). Such limited interest is mainly due to many trondhjemites having mantle-like isotopic signatures and little or no enrichment of many incompatible elements. Fusion of sialic crust should produce melts enriched in many incompatible elements, such as K 2O, Nb, Zr and the REE, because the crust is predominantly composed of low-pressure minerals, and bulk partition coefficients for these elements are generally low (Table 5.3). As a result, crustal melting is commonly considered to be the source of post-tectonic, minimum melt K-rich, alkaline granites. In addition, the time-integrated difference between extraction from the mantle and trondhjemite crystallisation is commonly considered to be too short for an intervening crustal episode, especially when the crust has significantly fractionated REE. With large degrees of partial melting, however, melts progressively approach the source composition (Fig. 5.9), and consequently the Carlindi trondhjemites may reflect complete melting of a source with a similar composition, possibly tonalite. In this scenario, the fractionated REE composition reflects an earlier magmatic event, probably like that described for second generation melts. Nevertheless, for the Carlindi trondhjemites, enrichment of HREE must still be suppressed during crustal fusion, perhaps by hornblende or garnet in the residue. A few of the Carlindi trondhjemite CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 152

samples contain hornblende and such a residue is consistent with the compatible Fe 2O3 and MgO trends (Fig. 5.1). Crustal fusion can also account for the slightly higher K 2O contents relative to the experimental models and the restricted, but highly evolved compositions of the trondhjemites. In a crustal fusion model, however, relics of earlier crust are expected, but have not been noted in the Carlindi trondhjemites. The absence of such relics, particularly zircon xenocrysts, may reflect high melting temperatures causing complete consumption of the earlier crust, as proposed for more recent granitoids (Chappell et al., 1998). The general viability of crustal fusion to form many of the chemical features of the trondhjemites is considered further in the discussion of the Carlindi granites, where there is analogous chemical behaviour but the additional evidence is compelling. In summary, the favoured model is that the Carlindi trondhje mites formed by fusion of pre-existing crust, but the evidence is not conclusive. Partial melting of basaltic crust cannot readily account for the extreme REE fractionation, HREE depletion and

K2O enrichment of the trondhjemites, although this process probably formed the precursor crust.

5.3.2.2 Granites Elemental abundances have been determined for five granite samples, three of which have been dated; gc400697 at ~3468 Ma, gneissic xenolith 090698 at ~3476 Ma and xenolith host 100698 at ~2925 Ma (gc400697 = 94058 in Chapter 3.5). Samples gc400697 and 100698 contain xenocrystic zircon, and 100698 also hosts tonalitic, granitic and greenstone xenoliths. The crystallisation age of sample 090698 has been defined from only three zircon grains and has a wide range of younger, concordant zircons. Hence, the old zircons may actually be xenocrysts. Irrespective of this, the Carlindi granites represent at least two magmatic events separated by ~540 myr. The five samples have been defined as granites according to th e QAP and feldspar ternary diagrams (Fig. 5.2), and plot along the calc-alkaline trend, distinct from the Carlindi trondhjemites (Fig. 5.8). Four of the granites have surprisingly similar compositions although they crystallised at different times and one has been intensely deformed. Relative to the trondhjemites they are enriched in Ba, Rb, Th, Nb, Ta, U, Pb,

SiO2 and K2O and depleted in Sr, Zr, TiO 2, Al2O3, Fe2O3, CaO, MgO and Na2O. They have pronounced negative Y and positive Ba, Nb and Sr anomalies relative to a CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 153 moderate negative slope (Fig. 5.4) and have extremely fractionated REE patterns (Fig. 5.5). Importantly, their HREE abundances are equivalent to the trondhjemites, and four of the granites have pronounced negative Eu anomalies. Compatibility trends cannot be defined because there are too few samples and the granites are clearly from different magmatic events. The fifth granite (150698) contains <1 % fine-grained, anhedral garnet, which has contributed to a markedly different composition compared to the other Carlindi granitoids. Fine-grained quartz-feldspar inclusions and the rounded, embayed morphology strongly suggest that the garnet is xenocrystic (Plate 2.3E).

Although it has an evolved composition (SiO 2 = 75.6 dry wt%), the presence of xenocrystic garnet indicates that this granite was not an equilibrated melt. Compositional similarities between the granites are remarkable and suggest a common petrogenesis. Moreover, fractionated REE patterns could indicate an ultimate derivation from partial melting of garnet amphibolite, as outlined for second generation trondhjemitic melts. Nevertheless, enrichment of many incompatible elements, particularly K 2O, suggests that the granites were derived by partial melting of sialic crust. The negative Eu anomalies indicate separation from a plagioclase-rich source. Direct evidence of a crustal source is provided by xenocrystic zircon in both the ~3468 and 2925 Ma granites and the inclusion of tonalitic to granitic xenoliths. The inclusion of xenocrystic garnet also constrains the composition of the source region. It is unlikely that these features are due to crustal contamination of a second generation melt because the high SiO 2 contents indicate that only evolved crust could have been assimilated, but the K2O enrichment would require very large amounts of such assimilation. The observed xenoliths and xenocrysts are too primitive to be suitable contaminants. Sm-Nd isotopic ratios support the derivation of the granites from a crustal source (Chapter 6). It is proposed, therefore, that the Carlindi granites were derived from partial melting of older tonalitic-trondhjemitic crust. Zircon xenocrysts in the ~3468 Ma granite suggest melting of relatively young crust, perhaps <40 myr old, whereas the ~2925 Ma granite was derived from older, heterogeneous crust. The resultant melts were enriched in many incompatible elements because the degree of melting was not very large, whereas the HREE were not enriched because garnet and possibly hornblende were present in the source region. Direct evidence of this residual garnet is provided as xenocrysts in sample 150698. This shows that HREE depletion does not preclude derivation of melts by crustal fusion, and that extreme REE fractionation can be CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 154 preserved. The similarities with the third generation melt model for trondhjemites are marked. The presence of xenocrysts and xenoliths and the greater abundance of many incompatible elements suggests that the degree of melting was low, possibly due to low fusion temperatures, in contrast to the high temperatures proposed for trondhjemite formation.

5.3.2.3 Anomalous samples There are four samples that are neither trondhjemites nor granites. Sample 070698 is a greenstone xenolith within the ~2925 Ma granite and its composition is broadly similar to the Pilgangoora basalts. It probably represents a raft of metamorphosed basalt, although it is over ten kilometres from the nearest greenstones. Sample 210698 is from a massive granitoid outcrop adjacent to the eastern Pilgangoora greenstones where there are abundant greenstone xenoliths. This sample is the least evolved of the granitoids (SiO 2 = ~57 dry wt%) and has no obvious affinities with any of the local felsic suites. Its broad Zr, Ti and Nb anomaly suggests a relationship with the greenstones (Fig. 5.4) and it may be appreciably contaminated. Both of these samples plot in the Coonterunah basalt field in the feldspar ternary diagram (Fig. 5.2). Sample 190698 is from a ~10 m wide coarse- to medium-grained porphyritic dyke within the eastern Pilgangoora greenstones. Its REE composition most closely matches the Coonterunah felsic volcanics (Fig. 5.5), but it does not have associated enrichment of Y, Zr and Nb (Fig. 5.4). It also does not plot within the defined fields on the feldspar ternary diagram (Fig. 5.2). In general, these three samples have no obvious affinities with the other Carlindi granitoids and are not discussed further. By contrast, gneissic xenolith 080698 is defined as a tonalite in both the QAP and feldspar ternary diagrams (Fig. 5.2). The xenolith is hosted within the ~2925 Ma granite and has an interpreted crystallisation age of ~3462 Ma (Chapter 3.5.1.6). It contains four older zircon grains, but they are not much older (<50 myr). The tonalite does not plot on the major element fractionation trends defined by the trondhjemites and granites (Fig. 5.1), and so there is no evidence to suggest that this is a less evolved part of those suites, although the sample is gneissic and has been modified. However, the tonalite has a similar trace element composition to the Carlindi trondhjemites, except for lower LREE and greater Y and HREE abundances. It also has a positive Eu anomaly, as do most of the trondhjemites, indicating plagioclase accumulation. Hence, the tonalite CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 155 has two important features: the less fractionated REE composition is more consistent with the experimental results for partial melting of garnet amphibolite, and its general composition approximates the proposed crustal source required to produce Carlindi trondhjemites and granites by crustal fusion. This sample, therefore, may be the link between second and third generation melts; that is, having the composition of the crust from which the Carlindi trondhjemites were derived. Irrespectively, the tonalite demonstrates that there were less evolved granitoids within the Carlindi domain.

5.4 DISCUSSION The felsic units in the field area represent at least four magmatic events, and thus they record crustal evolution over a significant part of the Pilbara Craton’s history. This evolution is discussed in relation to compositional differences between the felsic suites described here and those elsewhere in the Pilbara. The compositions of proposed crustal contaminants for the Pilgangoora basalts are also discussed.

5.4.1 Pilbara correlations The composition of Pilbara felsic volcanic units are here compared with those from the Coonterunah succession to determine whether they all formed by similar petrogenetic processes. Likewise, the Carlindi granitoids are compared with other coeval Pilbara granitoids to determine whether they are representatives of regional magmatic events caused by common processes. In the Marble Bar Belt, felsic volcanics from the ~3468 Ma Duffer Formation and the underlying North Star Basalt show the same variability as the Coonterunah felsic volcanics, mainly due to post-magmatic mobility of the LILE (Jahn et al., 1981). They have steep L-MREE and flatter M-HREE trends ([La/Yb] n = 7.4 to 20.0) with HREE abundances about 10 times chondritic. There are no obvious stratigraphic differences, although the samples come from two discrete parts of the greenstone succession. They have profound trace element differences with the Carlindi granitoids, precluding them from being extrusive equivalents. Indeed, they closely resemble the Coonterunah felsic volcanics from the eastern Pilgangoora domain, but typically have smaller negative Eu anomalies. However, Marble Bar felsic volcanics typically have lower Y and HREE abundances than their associated tholeiitic basalts (Gruau et al., 1987), which preclude them from forming by lower-pressure fractional crystallisation of these basalts. CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 156

Therefore, the Marble Bar felsic volcanics must have formed by a process distinctly different from that presumed to be responsible for the Coonterunah felsic volcanics. One possible model is that they represent similar assimilation processes as the Coonterunah eccentric suite, although no eccentric basalts have been sampled in the Marble Bar Belt (see Chapter 4.8.2). However, it is still difficult to reconcile this process with such profound HREE depletion. A more likely alternative is that the Marble Bar felsic volcanics formed by small-degrees of mantle melting with residual garnet. Hence, they are broadly similar to the adjacent basalts, but are relatively more enriched in many incompatible elements and depleted in Y and HREE. Sm-Nd isotope data show that the felsic and mafic units in the North Star Basalt were derived from an isotopically similar source, consistent with partial melting of the same mantle region (Hamilton et al., 1981). In contrast, Jahn et al. (1981) considered that the Marble Bar felsic volcanics and adjacent granitoids formed by partial melting of amphibolitic crust, even though they highlighted compositional differences between these rock-types. In the North Pole Dome, whole-rock compositions have been reported for felsic volcaniclastic and volcanic units from the Panorama Formation (Cullers et al., 1993).

These units generally have extremely fractionated REE patterns ([La/Yb] n = 20 to 40) with very low HREE abundances (Ybn = 1.5 to 4.0), and are profoundly different from the adjacent basalts (Green et al., in prep.). These REE compositions are very similar to those from the Carlindi granites which suggests that they were formed by similar processes; that is, by partial melting of pre-existing sialic crust with residual garnet or amphibole. The discovery of a ~3724 Ma xenocrystic zircon in the Panorama Formation may provide direct evidence of this crust (Thorpe et al., 1992a). In the Kelly Belt, the ~3430 Ma dacites have very similar REE patterns to the

Carlindi trondhjemites, but with greater abundances ([La/Yb] n = 15 to 30, Ybn = 7 to 10; Barley et al., 1984, 1998). Adjacent tholeiitic basalts have significantly greater Y and HREE abundances than the felsic volcanics, which indicates that the felsic volcanics cannot be the result of basalt fractionation (see Chapter 4.8.2). Instead, the Kelly Belt felsic volcanics may have formed by partial melting of sialic crust without the residual HREE sinks. In contrast, the ~3430 Ma McPhee Dome andesites and dacites have very similar compositions to the adjacent basalts, and do not show significant depletion of Y and the HREE ([La/Yb]n = 5 to 15, Ybn = 7 to 15; Barley et al., 1984, 1998). Notably, the McPhee Dome basalts have compositions similar to eccentric Coonterunah basalts; CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 157 that is, enriched in incompatible elements. Hence, the felsic volcanics in the McPhee Dome were probably formed by continued fractionation of basaltic magmas. In summary, the felsic volcanics from each Pilbara greenstone succession have distinct compositions and so they probably formed by different processes. Trace element abundances suggest that the felsic volcanics formed by fractionation of basaltic magmas in the Pilgangoora Belt and McPhee Dome, partial melting of sialic crust in the North Pole and Kelly Belt (but with differing effects of residual phases), and low-degrees of mantle melting with residual garnet in the Marble Bar Belt. The Pilbara granitoid complexes are composed of many distinct intrusive suites, which are commonly distinguished by textural and compositional criteria. For example, Blockley (1980) outlined compositional differences between pre- and syn-tectonic and more evolved post-tectonic intrusions, whereas Hickman (1983) noted the importance of the late tin granites. Many of the previous geochemical studies were not robust as there were few precise trace element data or age constraints. However, the Shaw and Mount Edgar Granitoid Complexes provide informative analogues to the Carlindi granitoids. In the Shaw Granitoid Complex, discrete phases were distinguished by cross- cutting and textural relationships (Bettenay et al., 1981) and subsequently confirmed by isotopic dating (Bickle et al., 1983, 1989, 1993; Zegers, 1996; Nelson, 1998, 2000). The older, pre- to syn-tectonic granitoids in the Shaw Complex are tonalites, trondhjemites and granodiorites (Bickle et al., 1983, 1993) and have extremely fractionated REE patterns ([La/Yb] n = 25 to 55) with very low HREE abundances

(Ybn = 1.5 to 5). Their trace element abundances are similar to the Carlindi trondhjemites with negative Y, Rb and Th and positive Sr anomalies relative to average continental crust. Isotopic systematics indicate that these older granitoids were derived from a source more enriched than the contemporaneous depleted mantle, suggesting they were mixtures of depleted mantle and older sialic crust (Bickle et al., 1993). Such a mixing model is only viable if both the mantle and crustal sources had appreciable garnet and hornblende to retain Y and the HREE. Alternatively, the older Shaw granitoids may have formed by whole-scale melting of older heterogeneous crust with a garnet- hornblende residue, as suggested for the Carlindi trondhjemites. The younger, post-tectonic Shaw granitoids are generally monzogranites and have variable but typically very fractionated REE patterns with pronounced negative Eu CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 158 anomalies (Bickle et al., 1989). Their REE abundances are greater than in the Carlindi granites, and they do not have the same HREE depletion. They are also more enriched in incompatible elements (Th, LREE) than the older Shaw granitoids. Isotopic systematics provide compelling evidence that these post-tectonic granitoids were derived from fusion of older heterogeneous sialic crust. The fractionated REE signatures were inherited from the source and the lack of HREE depletion indicates that garnet and hornblende were not significant in the melting residue. In summary, the Shaw granitoids show similar chemical evolution to the Carlindi granitoids with early tonalites and trondhjemites and later monzogranites. Varying degrees of partial melting of sialic crust can account for these temporal changes. Likewise, cross-cutting and textural relationships show that the Mount Edgar Granitoid Complex is also composed of various phases (Davy & Lewis, 1986; Collins, 1989, 1993) which generally intruded about 3300 Ma, although recent work has shown some age variations (Collins & Gray, 1990; Williams & Collins, 1990; Nelson, 2000). Whole-rock compositions of ~400 granitoid samples have been determined, although these analyses included few trace elements and many of these were commonly present below detection limits (Davy & Lewis, 1986). A lack of restite minerals, calcic plagioclase cores and mafic granitoid ènclaves led Collins (1993) to suggest that the Mount Edgar granitoids were derived from chemically evolved sources. Moreover, the earlier data (Davy & Lewis, 1986) were modelled to show that the granitoids were probably derived by partial melting of tonalitic-dacitic crust in the presence of abundant water (Collins, 1993). High degrees of partial melting were required to allow complete plagioclase dissolution in the tonalitic-dacitic precursors. The presence of xenocrystic zircons in many samples provides support for this crustal fusion model (Nelson, 2000). Thus, the Pilbara granitoid complexes are composed of multiple intrusions, most of which can be explained by partial melting of earlier sialic crust. The older Pilbara phases probably reflect high temperatures and large degrees of partial melting. The melts generally mirror their source with only limited enrichment of incompatible elements. Later granitoids formed at lower temperatures and lower degrees of partial melting and so they are greatly enriched in incompatible elements. The variable behaviour of Y and the HREE suggests that the precursor crust was profoundly heterogeneous.

5.4.2 Basalt contaminant CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 159

It was outlined in Chapter 4 that the Pilgangoora basalts show evidence of appreciable crustal contamination superimposed on an overall depleted mantle signature. Geological considerations and trace element models show the plausibility of crustal assimilation and fractional crystallisation producing basalts with these compositions. Modelling also shows that <10 % assimilation of the local trondhjemites can account for the contamination of most of the basalts. The eccentric basalts, however, are more enriched in many incompatible elements and probably reflect assimilation of different crustal material. The Coonterunah felsic volcanics also show two distinct compositional trends, consistent with fractionation of these basaltic suites. Thus, two distinct crustal contaminants are required to explain the range of Coonterunah magmatic compositions. As the main contaminant probably had a similar composition to the Carlindi trondhjemites, then another local rock-type may provide clues to the source of the eccentric contaminant. The eccentric suite is enriched in Zr, Nb, La, Ce, Pr, Ta, Th, U and P2O5 relative to Y and the M-HREE, and so a contaminant derived from a garnet- hornblende-bearing source is most likely. The Carlindi granites are enriched in Ba, Rb, Th, Nb, Ta, U and Pb and depleted in Sr and Zr relative to the trondhjemites, and so are broadly similar to the required contaminant, except for Zr and LREE differences. In detail, however, the differences between the granites and the trondhjemites are not great enough to account for the eccentric contamination. There are also other significant differences. For instance, both Nb and U are enriched in the granites but they have lower Nb/U ratios than the trondhjemites. Consequently, the granite trend is further from the eccentric basalts on mixing diagrams than the trondhjemite trend (Fig. 4.10). Hence, direct assimilation of rocks like the Carlindi granites cannot produce the eccentric contamination. Instead, the processes that formed the granites may be analogous to the processes that formed the contaminant; that is, fusion of garnet-hornblende-bearing sialic crust. Thus, during transit through the crust, the Coonterunah-Warrawoona basaltic magmas could have induced partial melting of trondhjemitic basement to produce small, highly enriched melts. These melts separated from a garnet-hornblende- bearing source and so were not enriched in Y or the M-HREE, but were highly enriched in many other trace elements. The basaltic magmas may then have incorporated variable amounts of these melts on top of their general assimilation of trondhjemitic basement. Hence, assimilation involved two distinct processes: whole-rock incorporation, and CHAPTER 5: ELEMENTAL GEOCHEMISTRY – FELSIC ROCKS 160 induced melting and mixing of partial melts. This double contamination suggests that the basement to the early Pilbara basalts was complex, comprising domains with discrete mineralogical compositions. The variety of xenoliths supports this supposition.

5.5 SUMMARY The Coonterunah felsic volcanics have compositions suggesting that they formed by late-stage fractionation of the Coonterunah basalts, a process that may have also formed some other felsic volcanic successions in the Pilbara. The eccentric contamination of some Coonterunah magmas was probably caused by mixing with small degrees of partial melt induced by the basaltic volcanism. These melts must have separated from a garnet-hornblende-bearing source. The Carlindi granitoids represent two discrete phases: early trondhjemites and later granites. These probably formed by fusion of older sialic crust under different conditions. The trondhjemites represent large degrees of crustal melting at high temperatures with residual garnet or hornblende, and so are enriched in some incompatible elements, but not Y and the HREE. The granites represent smaller degrees of partial melting at lower temperatures, also with a garnet-hornblende residue, and so are greatly enriched in many incompatible elements, but not Y and the HREE. Similar compositional trends have been noted in some other Pilbara granitoid complexes, and so crustal fusion may represent an important process in the Pilbara Craton. Regional variations may be explained by differing mineralogy in the source regions. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 161

Chapter 6: RADIOGENIC ISOTOPE GEOCHEMISTRY

6.1 INTRODUCTION Aside from determining geological ages, radiogenic isotopes can also be used to constrain petrogenetic processes. Since the mass difference between the isotopes of large mass elements is relatively small, their isotopes are not fractionated during many geological processes, particularly those involving crystal-melt equilibria. Most radiogenic elements have large masses, and so for these elements, the melt inherits the isotopic composition of the source and this remains constant during crystallisation. This provides an important foundation for the study herein because the source of both basaltic and granitic melts can thus be characterised and mixing between distinct sources can be identified. Hence, the objective of this chapter is to determine the isotopic composition of the Pilgangoora greenstones and the Carlindi granitoids. Parent and daughter elements, however, are c ommonly fractionated during geological processes, and so each part of a geological system may evolve contrasting isotopic compositions. Using few assumptions, the history of a magmatic source can therefore be determined. Such methods can be applied to the Pilgangoora rocks, and their implications are also discussed in this chapter. The interpretation of isotopic ratios is hotly debated, with argument generally converging on one issue: the influence of specific events on isotopic ratios. The problem is highlighted in studies of the intensely deformed and metamorphosed early Archaean Isua Greenstone Belt, where whole-rock Nd isotopes are variously interpreted to be magmatic (Hamilton et al., 1978, 1983; Bennett et al., 1993) or reset by later metamorphic events (Gruau et al., 1996; Vervoort et al., 1996; Moorbath et al., 1997; Vervoort & Blichert-Toft, 1999; Blichert-Toft et al., 1999a). Even when a ratio can be attributed to a certain event, conflicting explanations are sometimes presented. This is probably best illustrated by the use of isotopic data to unravel the geochemical evolution of the mantle and the extraction of continental crust. This is a major theme of this thesis and is discussed in Chapter 9. Two isotopic systems have been used in this study: 147Sm-143Nd and 176Lu-176Hf. Significantly, these isotope pairs have covariant behaviour during many geological processes. For instance, during typical melting and crystallisation processes the parent isotopes are more compatible than their daughters, and thus the Nd-Hf isotopic systems CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 162 of primitive domains evolve faster than for chemically evolved domains. In contrast, the Rb-Sr and U-Pb systems have more compatible daughters. Furthermore, Sm, Nd, Lu and Hf are refractory and lithophilic, and so were not fractionated in the solar nebula or during planetary core formation (DePaolo, 1988). Hence, fractionation between Sm-Nd and Lu-Hf probably reflects magmatic processes, and thus they record the differentiation of the terrestrial mantle and the formation of crust. Therefore, they are crucial to constraining the processes that formed continental crust.

6.2 PREVIOUS WORK There have been surprisingly few studies in the Pilbara Craton that have used isotopes for constraining petrogenetic models, with most of the early work concentrating on determining magmatic ages. Many of these early studies used whole- rock Rb-Sr and Pb-Pb which have been variably disturbed by later thermal events, especially for the older units (Oversby, 1976; Pidgeon, 1978a; Cooper et al., 1982; Bickle et al., 1993). Three studies in the Marble Bar Belt reported the Sm-Nd isotopic ratios of 28 samples that were used to define an isochron and estimate the magmatic age of the greenstones (Hamilton et al., 1981; Jahn et al., 1981; Gruau et al., 1987). More precise zircon U-Pb magmatic ages have since been determined for this succession (McNaughton et al., 1993; Nelson, 1999, 2000), and so the Sm-Nd isotopic data can instead be used for constraining the petrogenesis of the Marble Bar greenstones. In the Shaw Granitoid Complex, the Sm-Nd isotopic ratios of 7 early granodiorites and 8 post- tectonic monzogranites were obtained to characterise their sources (Bickle et al., 1989, 1993). In the west Pilbara, the Sm-Nd isotopic ratios of 8 granitoids have been determined, along with precise zircon U-Pb ages of 6 of these samples (Smith et al., 1998). The Sm-Nd isotopic ratios of 19 samples from the <2.76 Ma Fortescue Group have also been obtained to characterise their source (Nelson et al., 1992). Only one Lu- Hf study has so far been reported from the Pilbara Craton, with analyses of zircons from three felsic units as part of a global survey of ancient terrains (Amelin et al., 2000). A number of whole-rock Lu-Hf studies are currently underway (Vervoort et al., 1999b). CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 163

6.3 ANALYTICAL METHODS The Sm-Nd isotopic ratios and elemental abundances of twenty-six whole-rock samples were determined at the Department of Earth, Atmospheric and Planetary Science, Massachusetts Institute of Technology. Analytical techniques followed Bowring et al. (1989) and are summarised here. Samples were crushed and milled to a very fine powder, and between 200 and 400 mg of each powdered sample was spiked 149 150 with ~0.05 g of Sm- Nd tracer. The spiked samples were dissolved with HF-HNO3 in steel-jacketed Teflon dissolution vessels at 220° C for five days followed by conversion to chlorides through addition of HCl and fluxing for 48 hours at 120 ° C. The separation and purification of Sm and Nd followed a standard two-stage cation exchange and hydrogen-di-2-ethyl-hexyl-phosphate-(HDEHP)-on-teflon column procedure. Isotopic ratios were determined using a VG Sector 54 multicollector thermal ionisation mass spectrometer (TIMS). Samarium was loaded on single Ta filaments with 152 H3PO4 and analysed in static multicollector mode with a Sm ion beam of 0.5 V, whereas Nd was loaded on triple Re filaments with H 3PO4 and analysed in dynamic multicollector mode with a 144Nd ion beam of 1.5 V. The Sm and Nd data were normalised to 152Sm/147Sm = 1.783 and 146Nd/144Nd = 0.7219, respectively. The La Jolla standard at MIT produced an average 143Nd/144Nd value of 0.511845 ± 10 (Bowring & Housh, 1995; accepted value 0.511860), and during the measurement period the USGS standard BCR-1 gave average values of 143Nd/144Nd = 0.512642 ± 5 and 147Sm/144Nd = 0.1380 ± 16 (2s, n = 3, accepted values 0.512640 ± 30 and 0.1371 ± 13, respectively). These values are a measure of the external reproducibility of the standards and provide a reasonable estimate of the analytical precision of the reported ratios. The 2 s errors given for the 143Nd/144Nd ratios are in-run errors and show the relative quality of the analyses (Appendix 3). The ratios obtained for both standards were within analytical error of the accepted values, and thus reported ratios have not been bias-corrected. An external error of ± 0.16 % was estimated for 147Sm/144Nd. Six replicate samples were also run. The Lu-Hf isotopic ratios and elemental abundances of ten whole-rock samples were performed at Ecole Normale Supérieure de Lyon, France. Analytical techniques followed Vervoort & Blichert-Toft (1999) and are summarised here. Samples were crushed and milled to a very fine powder, and between 200 and 250 mg of each powdered sample was spiked with a 176Lu-180Hf tracer. The spiked samples were CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 164

dissolved with HF-HNO3 in steel-jacketed Teflon dissolution vessels at 160° C for 5-7 days. The separation and purification of Lu and Hf followed a three-stage ion-exchange separation procedure; 1) Hf columns with AG50W-X12 cation resin, 2) Hf columns with AG1-X8 anion resin, and 3) removal of Ti with AG50W-X8 cation resin. Analyses were performed using a Fisons Instruments VG Plasma 54 Magnetic Sector-Multiple Collector-Inductively Coupled Plasma-Mass Spectrometer (MS-MC-ICP-MS). Lu and Hf were measured statically and Lu analyses used natural Yb in the sample to correct for mass fractionation (Blichert-Toft et al., 1997). Results were normalised for mass fractionation to 179Hf/177Hf = 0.7325, bias factors were verified by regular measurement of the JMC 475 Hf standard (every fourth sample on average), and machine drift was minimal during the runs. Samples have been normalised to the JMC 475 Hf standard (176Hf/177Hf = 0.282160), and total procedural Hf and Lu blanks were less than 25 and 20 pg, respectively, or at least 1000 times lower than the processed Hf and Lu. The 2 s errors for the 176Hf/177Hf ratios are in-run errors and are given to show the relative quality of the analyses (Appendix 3). An external error of <1 % was estimated for 176Lu/177Hf. No replicates were run.

6.4 RESULTS Below, the reproducibility of standard and replicate samples is checked to verify the analytical precision, and elemental abundances are also compared with SN-ICP-MS results to provide further support. The post-magmatic immobility of the REE has already been established, but since there were problems with Zr-Hf dissolution during SN-ICP-MS sample preparation (Chapter 4, 5), this is confirmed using the isotopic results.

6.4.1 Sm-Nd system The 147Sm/144Nd and 143Nd/144Nd ratios of 13 Coonterunah and 7 Warrawoona basalts, 1 Carlindi trondhjemite, 2 Carlindi monzogranites and 3 xenoliths hosted in ~2925 Ma Carlindi granite have been determined. The basalts include tholeiitic, magnesian and gabbroic compositions, and there are two basalts from the eccentric Coonterunah suite (Appendix 3). As outlined above, the standards analysed with the samples are within error of their expected values, and so the overall analytical procedure was reliable. In addition, CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 165 six samples were replicated, some in separate dissolutions, to determine the sample variability. Replicate differences for 147Sm/144Nd are generally less than 0.25 % (one sample at 0.48 %) and for 143Nd/144Nd mostly less than 0.003 % (one sample at 0.007 %). This variability is slightly greater than for the standards, and hence the stated error, so there are probably small heterogeneities in the powders. Elemental abundances typically vary by <10 % from those obtained by SN-ICP-MS (Appendix 2), with isotope dilution consistently producing lower values and suggesting a minor procedural bias. Differences are greater for the granitoids and this may reflect increased powder heterogeneity. However, the Sm and Nd variations are strongly correlated, and differential Sm/Nd ratios are typically <5 %. Replicate variations have a similar magnitude to those between analytical methods. In summary, the sample powders have small heterogeneities, but these have had only limited affects on the Sm-Nd isotopic results, and so the isotopic ratios are interpreted to reflect those of the powders. The Sm-Nd elemental abundances obtained during i sotopic analyses are plotted versus Zr contents (pressed-powder XRF; Appendix 2) to confirm that these elements have not been modified by post-magmatic processes (Fig. 6.1). This follows the procedure outlined in Chapter 4 where the regression line for highly incompatible elements is forced through the graphical origin. Both Sm and Nd show strong positive correlations that are consistent with magmatic trends, and hence they have not been significantly modified by post-magmatic processes. The two eccentric basalts have higher Zr abundances relative to Sm, consistent with previous determinations (Chapter 4.7.4). Unfortunately, there are too few granitoid samples to define magmatic trends, although the SN-ICP-MS results are consistent with Sm-Nd immobility. However, this may not be the case for the gneissic xenoliths. Nevertheless, for most of the Pilbara rocks the Sm-Nd isotopic ratios can be confidently interpreted to be primary magmatic features. The samples have variable 147Sm/144Nd and 143Nd/144Nd ratios due to the time- integrated decay of 147Sm in broadly similar aged rocks with diverse Sm/Nd ratios. The ratios show a positive linear correlation, similar to an ancient isochron, and are slightly displaced above the present-day bulk silicate Earth (CHUR) value (Fig. 6.2). The basalts have the largest isotopic ratios reflecting their less evolved compositions, and the eccentric basalts have relatively low ratios reflecting their profound LREE enrichment. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 166

The Coonterunah basalts have much greater variability than the Warrawoona basalts, possibly due to their larger sample population.

0.8 Sm Lu 2 2 6.0 r = 0.873 r = 0.937 0.6

4.0 0.4

2.0 0.2

0.0 Zr 0.0 Y 0 50 100 150 200 250 0 10 20 30 40 50

25 6.0 Nd Hf 2 5.0 2 20 r = 0.954 r = 0.993

4.0 15

3.0

10 2.0

5 1.0

0 Zr 0.0 Zr 0 50 100 150 200 250 0 50 100 150 200 250

Figure 6.1: Elemental abundances for Coonterunah and Warrawoona basalts obtained during isotopic analyses have well-defined positive correlation with Zr (pressed-powder XRF) and Y (SN-ICP-MS). Regression lines have been forced through the graphical origin, abundances in ppm. Coonterunah and Warrawoona results have been regressed together to increase population statistics, although they are discrete successions. Diamonds- Coonterunah, circled – eccentric, squares - Warrawoona.

The basalts can be used to define imprecise isochrons (or errorchrons). The Coonterunah basalts define an isochron with an age of 3378 ± 160 Ma (MSWD = 53), which is greatly influenced by sample gc200697. Its removal redefines the isochron to 3470 ± 100 Ma (MSWD = 17). Both of these estimates are well below, but within error of, the established zircon U-Pb age of ~3517 Ma. The Warrawoona basalts define an isochron with an age of 3148 ± 550 Ma (MSWD = 25), which is also well below, but also within error of, its established U-Pb age. These age estimates are probably poor because they are derived from basalts with relatively restricted 147Sm/144Nd ranges. Moreover, the basalts define errorchrons rather than isochrons (large MSWD; Brooks et al., 1972), and so they do not satisfy all the required isochron assumptions; that is, they probably have diverse source components. However, the consistently low age estimates are unusual for magmatic rocks that have supposedly assimilated older continental crust. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 167

For instance, binary mixing has been well documented for producing falsely old isochrons (Cattell et al., 1984; Chauvel et al., 1985; Gruau et al., 1990b).

0.513 T = 0

0.2824 0.512 T = 0 0.2820 d

0.511 f N H 4 0.2816 CHUR 7

4 CHUR 7 1 / 1 / f d 0.510 0.2812 H N 6 3 7 4 1 1 T = 2925 T = 3434 0.509 0.2808 T = 3434

0.2804 0.508 T = 3517 T = 3517 0.10 0.12 0.14 0.16 0.18 0.20 0.015 0.020 0.025 0.030 0.035 147Sm/144Nd 176Lu/177Hf

Figure 6.2: Measured isotopic ratios for Pilbara samples define negative sloping trend whereas calculated ratios at time of formation define flat trend. Evolution of CHUR is shown with relevant ages marked. Replicate analyses not included. Solid diamonds - Coonterunah basalt, circled - eccentric, open square - Warrawoona basalt, open triangle Carlindi granitoid, circled triangle - xenolith. T-values refer to position of CHUR in millions of years.

6.4.2 Lu-Hf system The 176Lu/177Hf and 176Hf/177Hf ratios of 5 Coonterunah and 5 Warrawoona basalts have been determined. These include magnesian and tholeiitic basalts and two eccentric Coonterunah basalts. The isotopic standards analysed with the samples are within error of the expected values, and so the analytical procedure is reliable. No replicates were analysed, so the Lu-Hf heterogeneity within the powders is not known. Hafnium abundances were significantly underestimated by SN-ICP-MS (Chapter 4.3.2), and so its post-magmatic immobility has not been established. However, its generally analogous behaviour with Zr would suggest that it was highly refractory. This is confirmed by the very strong correlation between Hf and Zr abundances, which indicates that Hf has remained immobile since magmatism (Fig. 6.1). Due to the variations between the HREE and many incompatible elements during eccentric contamination, Lu is compared with Y instead of Zr. They are strongly correlated, and so Lu has also been unmodified by post-magmatic processes. Therefore, the Lu-Hf isotopic compositions can be confidently interpreted to represent magmatic abundances. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 168

The 176Lu/177Hf and 176Hf/177Hf ratios are variable due to the time-integrated decay of 176Lu in rocks with diverse Lu/Hf ratios. They define a positive linear correlation, similar to an ancient isochron, that is displaced well below the present-day bulk silicate Earth (CHUR) value (Fig. 6.2). Both basalt suites have large isotopic ranges with the eccentric basalts having the lowest ratios, again reflecting their profound compositional differences from the main suite. The Coonterunah basalts define an isochron with an age of 3188 ± 460 Ma (MSWD = 24) and the Warrawoona basalts define an isochron with an age of 3073 ± 420 Ma (MSWD = 5.9). As for the Sm-Nd isochrons, these values are within error of, but generally much lower than, the established zircon U-Pb ages.

6.5 INTERPRETATION The measured Sm-Nd and Lu-Hf isotopic ratios provide only limited information about the Pilbara rocks because they result from prolonged decay in rocks of differing original composition. More information can be obtained from their initial magmatic ratios: the 143Nd/144Nd and 176Hf/177Hf ratios at the time of crystallisation. This is easily determined by subtracting the time-integrated growth of the daughter from the measured present-day value, but is only valid where the rocks have retained their magmatic compositions and have well-constrained ages. It has already been established that the Pilbara samples have preserved their magmatic Sm, Nd, Lu and Hf abundances, and most of the samples have well- constrained ages. The Coonterunah basalts have an eruption age of 3517 ± 3 Ma, as determined by SHRIMP U-Pb zircon dating of two felsic volcanic units interlayered with the basalts. Geological considerations suggest an age of ~3434 Ma for the Warrawoona basalts, although they have not been directly dated and may be as old as ~3458 Ma. The younger age is used here and the implications for an older age are discussed where appropriate. Precise SHRIMP U-Pb zircon ages have been obtained for the three Carlindi granitoids. The two gneissic xenoliths have ages based on the oldest concordant zircon populations, although these samples may be somewhat younger as they have prolonged tectonothermal histories and are included in a ~2925 Ma granite. The mafic xenolith has not been dated, but it is grouped with the other xenoliths and assigned a notional age of ~3470 Ma, equivalent to Talga Talga Subgroup volcanism. In summary, CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 169 the initial magmatic Nd-Hf isotopic ratios can be confidently determined for most of the Pilbara samples. By themselves, initial ratios are rather uninformative because comparison between rocks of different ages is not straightforward. A time-integrated reference is thus used to help conceptualise ancient initial isotopic ratios. The Chondritic Uniform Reservoir (CHUR) reference is used for Sm-Nd and Lu-Hf (DePaolo & Wasserburg, 1976), and essentially provides a time-integrated reference to compare the growth of 143Nd/144Nd and 176Hf/177Hf throughout Earth history. CHUR is based on the assumptions that the Sm-Nd and Lu-Hf pairs were not fractionated during nebular condensation or planetary core formation and that chondritic meteorites have sampled the early nebula. It follows that the bulk silicate Earth (BSE or primitive mantle), solar nebula and chondritic meteorites have equivalent Nd-Hf isotopic compositions. Therefore, isotopic measurements of chondrites provide the present-day CHUR values, and these can be calculated back through time. Importantly, the value of CHUR is independent of the age of the Earth and any Sm-Nd or Lu-Hf fractionations are interpreted to be due to magmatic processes within the silicate portion of the Earth. The CHUR concept thus provides a reference for both Nd and Hf isotopes through time and against which ancient samples can be compared. The relative deviation from CHUR is measured with the epsilon ( e) parameter:

143 Nd 143 Nd 144 (T ) - 144 (CHUR,T) e(Nd) =10 4 [ Nd Nd ] 143 Nd (CHUR,T) 144 Nd where T is the age of the sample. The e(Hf) relationship has an identical form. It should be noted that the denominator decreases with age and so there must be a corresponding decrease in the numerator to maintain parity of e-values. In other words, a modern e- unit has a greater absolute difference between the sample and CHUR than an ancient example. However, this change is approximately 1 % for both Nd and Hf over 4.5 byr and so is relatively insignificant, particularly when comparing samples with similar ages as in the Pilbara. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 170

There is presently considerable debate about both the Nd and Hf values of the bulk silicate Earth (CHUR), and so commonly accepted values are used herein (Appendix 3). However, it is noted that these values have been arbitrarily chosen from the mode of chondritic 147Sm/144Nd and 176Lu/177Hf values, and that only one chondrite is common to both datasets (Jacobsen & Wasserburg, 1980; Blichert-Toft & Albarède, 1997; Salters & White, 1998). An improved understanding of the Nd-Hf isotopic character of chondrites may well change the accepted values for the bulk silicate Earth, which may profoundly change the present interpretation of magmatic processes (see Chapter 9). The 176Lu decay constant used herein is also an accepted standard (Sguigna et al., 1982), but using the other commonly accepted value (Tatsumoto et al., 1981) does not change initial e(Hf) by more than the stated analytical error. However, a much lower value has been proposed by Nir-el & Lavi (1998), which, if adopted, would significantly change initial e(Hf) for the Pilbara samples, and hence would necessitate a profoundly different interpretation. However, this value has been hotly contested from various standpoints (Amelin et al., 2000; Begemann et al., 2001), and so it has not been adopted here. All isotopic values from previous studies have been recalculated with the values used here.

6.5.1 Sm-Nd system The samples are plotted on an initial e(Nd) versus age diagram (Fig. 6.3). The Coonterunah basalts cluster in a 1.3 e range (e(Nd) = 1.16 to 2.47), except for one outlier (gc200697; e(Nd) = –0.73), and the Warrawoona basalts have a 1.7 e range (e(Nd) = –0.06 to 1.65). Changing the Warrawoona age to 3458 Ma increases initial e(Nd) by ~0.06, less than half the reported error, and so is rather insignificant. The Coonterunah basalts generally have greater initial e(Nd) than the Warrawoona basalts, although there are ten samples from the combined population with overlapping values. There is no correlation between initial e(Nd) and the stratigraphic height of the Coonterunah basalts, but initial e(Nd) may decrease up the Warrawoona succession, although this trend is defined by only a few samples (Fig. 6.4). Magnesian basalts have the largest initial e(Nd) in both the Coonterunah and Warrawoona successions, and the eccentric basalts also have high initial e(Nd). Excluding the single outlier, each suite has remarkably little variability, particularly as the samples represent a wide range of CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 171 compositions and stratigraphic positions. In contrast, the initial e(Nd) of modern oceanic basalts span more than 10 e-units (Ito et al., 1987), the ~800 m thick Miocene Picture Gorge Basalt spans ~2.8 e-units (Brandon et al., 1993), mafic and ultramafic rocks from the Cretaceous Gorgona Island span 4.7 e-units (Arndt et al., 1997) and the <2.76 Ga Fortescue Group and ~2.70 Ga Ventersdorp Supergroup span more than 3 e-units (Nelson et al., 1992). The tighter initial e(Nd) range of the Pilgangoora basalts may reflect their antiquity, when there was less time for radiogenic ingrowth.

3.0 4.0

2.0 2.0

1.0 e(Hf) e(Nd) 0.0 0.0

-2.0 -1.0

2900 3100 3300 3500 3400 3450 3500 3550 3600 Age (Ma) Age (Ma)

Figure 6.3: Initial Nd-Hf compositions ( e-values) of the Pilgangoora basalts and Carlindi granitoids. CHUR parameters given in Appendix 3 and discussed in text. Symbols as for Figure 6.2.

The older Carlindi granitoids and xenoliths also have positive initial e(Nd): the ~3479 Ma Carlindi trondhjemite has initial e(Nd) = 1.68 and the cross-cutting ~3468 Ma granite has initial e(Nd) = 1.26. These are within the overlap range of the two greenstone successions. The ~2925 Ma granite has initial e(Nd) = -1.04, whereas the xenoliths it hosts have relatively high e(Nd): the gneisses have initial e(Nd) = 2.36 and 2.43 and the greenstone has initial e(Nd) = 2.32 at their interpreted ages. In general, most of the Pilgangoora and Carlindi samples were derived from sources with greater initial 143Nd/144Nd ratios than CHUR, and hence, from sources with time-integrated superchondritic Sm/Nd ratios. In other words, their sources had a prior history of LREE depletion and Sm-Nd differentiation, and such processing preceded Pilbara magmatism by a sufficient time interval to allow distinct 143Nd/144Nd ratios to develop. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 172

6.5.2 Lu-Hf system The Pilgangoora basalts are shown in an initial e(Hf) versus age plot (Fig. 6.3). The Coonterunah basalts have a <1.3 e range (e(Hf) = 1.59 to 2.87) with one outlier (gc010697; e(Hf) = -2.29), whereas the Warrawoona basalts have a <0.8 e range (e(Hf) = -1.04 to –0.27) with one outlier (gc300697; e(Hf) = 2.14). The main groups have little variability, although this possibly reflects their small populations. The Coonterunah basalts generally define a discrete population with its outlier close to the Warrawoona population, and vice versa. There are no stratigraphic trends, although there are too few samples to be confident about this (Fig. 6.5). The eccentric basalts have positive initial e(Hf), similar to the other basalts. In contrast to e(Nd), the magnesian basalts do not have the largest initial e(Hf). The Pilgangoora basalts were derived from variable sources that were both depleted and enriched in 176Hf/177Hf relative to CHUR, and hence their sources had time-integrated superchondritic and subchondritic Lu/Hf ratios. This reflects Lu-Hf fractionation of the sources prior to magmatism over a period sufficient to develop distinct 176Hf/177Hf ratios.

6.5.3 Basalt components The initial Nd-Hf isotopic compositions of the Coonterunah and Warrawoona basalts are diverse and may reflect either heterogeneity within the mantle, mixing between isotopically distinct sources or a combination of both. It has already been shown that the basalts have trace-element abundances consistent with addition of a crustal component to N-MORB, that is, Ta-Nb depletion associated with Th, U and LREE enrichment (Chapter 4.7.2). As Nd and Hf isotopes are not fractionated during partial melting or crystallisation, it should then be possible to characterise the mantle and crustal components if they were isotopically distinct. This is achieved by plotting elemental ratios that vary greatly between crustal and mantle sources against initial e(Nd) and e(Hf) (Fig. 6.4, 5). Compared with their Sm/Nd, Nb/Ta, Nb/La and Nb/U ratios, both the initial e(Nd) and initial e(Hf) of the Pilgangoora basalts have negative correlations. Since these elemental ratios are typically lower in crust-derived rocks than in mantle-derived rocks, as shown by the relative position of the granitoids and basalts in the diagrams, then greater crustal contributions should be associated with lower elemental ratios. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 173

Therefore, increasing initial e(Nd, Hf) in the basalts should correspond with a greater crustal contribution. The same relationship is also demonstrated by the positive correlation between [Sm/Yb]n and initial e(Nd, Hf). Importantly, Nb/Ta and Nb/U are not significantly fractionated during basalt petrogenesis, and so these variations can be attributed solely to the crustal component. In contrast, the Sm/Nd, Nb/La and [Sm/Yb] n ratios are fractionated during typical magmatic processes, and thus a relatively small crustal contribution may have been amplified, although this will not affect initial isotopic ratios. Nevertheless, the relationship between the trace-element ratios and initial e(Nd, Hf) suggests that the isotopic diversity is predominantly due to mixing between a mantle source and a crustal component with greater initial 143Nd/144Nd and 176Hf/177Hf.

Sm/Nd 6.0 Strat. height 0.3

4.0 0.2

2.0

0.1

0.0

Nb/Ta Nb/La 1.2 16

12 0.8

8

0.4 4

40 Nb/U [Sm/Yb]n 2.0

30

20 1.0

10

-1.0 0.0 1.0 2.0 3.0 -1.0 0.0 1.0 2.0 3.0 e(Nd) e(Nd)

Figure 6.4: Relationships between e(Nd) and elemental ratios show the influence of crustal components. Stratigraphic height is measured from a komatiitic basalt marker horizon in the Coonterunah succession. Symbols as for previous figures. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 174

strat. height 6.0 Lu/Hf 0.2

4.0

2.0 0.1

0.0

Nb/Ta Nb/La 1.2 16

12 0.8

8

0.4

4

40 Nb/U [Sm/Yb]n 2.0

30

20 1.0

10

-2 -1 0 1 2 3 -2 -1 0 1 2 3 e(Hf) e(Hf)

Figure 6.5: Relationships between e(Hf) and elemental ratios show the influence of crustal components. Symbols as for previous figures.

These diagrams also show that the Coonter unah and Warrawoona basalts have similar initial e(Nd, Hf) at corresponding trace-element compositions. Hence, there is no evidence to suggest that they were derived from isotopically distinct sources, as perhaps implied by their initial e-values alone. Instead, the Warrawoona basalts represent various degrees of mixing of components that were isotopically similar to those comprising the Coonterunah basalts. This supports a suggestion made earlier that both successions formed by equivalent petrogenetic processes, and thus the Pilgangoora basalts can be generally considered together. The diagrams also show that Coonterunah basalt gc200697 (e(Nd) = -0.73) can be readily explained by a relatively small crustal contribution, and so it is not as anomalous as suggested by its initial e(Nd) alone. Thus, the isotopic composition of gc200697 supports its use as the starting point for the assimilation models presented earlier (Chapter 4.7.5). Moreover, the eccentric basalts CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 175 are not isotopically distinct, but probably reflect addition of a greater amount of the crustal component. The initial Nd-Hf isotopic diversity of the Pilgangoora basalts can thus be accounted for by mixing between at least two isotopically distinct sources. Given the simplest two component system and using Nb/U ratios to measure the crustal contribution, as discussed for the trace-element mixing models (Chapter 4.7.5), then the isotopic composition of the components can be estimated. To start, the mantle source should have the highest Nb/U ratios, and so initial e(Nd) = -1.0 to +1.5 and initial e(Hf) = -2.5 to 0.0. Hence, the magnesian basalts are less evolved but more contaminated than some tholeiitic basalts, as previously noted from their trace element compositions. To estimate the isotopic composition of the crustal component, however, requires using a mixing relationship with the form:

a 144 a a a 144 b b m X [ Nd] e Nd + (1 - X )[ Nd] e Nd e Nd = X a[144Nd]a + (1 - X a )[144Nd]b where subscripts a and b are the components and m is the mixture, X is the proportion of one component and the total abundance of Nd or Hf is used instead of 144Nd or 177Hf (DePaolo & Wasserburg, 1979). The initial e(Nd) of the contaminant can be solved for each basalt assuming the mantle Nd composition, the elemental Nd abundances of the crust and the crustal contribution from previous estimates of their Nb/U ratios (Fig. 4.10). The results are extremely variable depending on these assumed values, such that the model is not very useful for its intended purpose. However, if the mantle source had initial e(Nd) < 1.5 and initial e(Hf) < 0.0, as proposed above, then the magnesian basalts must have assimilated quite distinct crust: very enriched in Nd and Hf, but with initial e(Nd, Hf) > 10. A more reasonable alternative is that the magnesian basalts were derived from a mantle source with initial e(Nd) > 1.5 and initial e(Hf) > 0.0, and so the mantle was more heterogeneous than suggested by the Nb/U ratios alone. Nevertheless, the crustal component must have had, at least in part, initial e(Nd, Hf) greater than the magnesian basalts. In summary, covariations between initial e(Nd, Hf) and certain trace-element ratios indicate that the Pilgangoora basalts formed by mixing between mantle material CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 176 and a crustal component with greater initial 143Nd/144Nd and 176Hf/177Hf than the mantle source. Moreover, the mantle source was relatively heterogeneous (e(Nd) = -1.0 to 2.0, e(Hf) = -2.5 to 1.5), whereas the crustal component is poorly constrained, but it generally had initial e(Nd, Hf) greater than the mantle source.

6.5.4 Origin of the components Given the above estimates of initial e(Nd) and initial e(Hf), the origins of the mantle and crustal components can be constrained. These components are modelled as combinations of four general isotopic domains: primitive and depleted mantle and basaltic and sialic crust. Time-integrated trajectories for these domains are shown in Figure 6.6, where the basalt arrays have the same slopes as the Pilgangoora basalts, and so are more enriched in incompatible elements and thus slightly steeper than for N- MORB. The trondhjemite and tonalite represent typical Archaean sialic crust and the slope of the depleted mantle is consistent with previous estimates (DePaolo, 1981).

tonalite

2 trondhjemite 4

1 2 depleted depleted Pilbara Pilbara mantle mantle ) ) f d 0

H 0 N ( ( e e

-1 basaltic crust -2 basaltic crust

-2 -4

3400 3600 3800 4000 4200 3400 3600 3800 4000 4200 Age (Ma) Age (Ma)

Figure 6.6: Isotopic trajectories of various rock-types and domains. Basaltic crust is Pilgangoora basalts, trondhjemite from Carlindi complex, tonalite from Greenland (~3.7 Ga, Vervoort & Blichert-Toft, 1999) and postulated depleted mantle. Note that the e(Hf) scale is double the e(Nd) scale.

6.5.4.1 Mantle source The Pilgangoora basalts were derived from an isotopically heterogeneous mantle source with initial Nd-Hf values above and below CHUR. Therefore, the mantle source had an averaged time-integrated history with both super- and subchondritic Sm/Nd and Lu/Hf ratios. This diversity cannot be realistically modelled in one step as typical melting CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 177 of primitive mantle would leave a depleted mantle with superchondritic Sm/Nd and Lu/Hf ratios, which would evolve towards positive e(Nd,Hf). Mixing this depleted mantle with primitive mantle would force the combined mantle towards CHUR, but it cannot produce negative e-values. Instead, the mantle source requires a component with time-integrated subchondritic Sm/Nd and Lu/Hf ratios, for example, the recycled products of mantle melting such as basaltic or sialic crust. Thus, the Pilgangoora basalts must have been derived from a mantle that formed by at least two steps: melting of primitive mantle to form a depleted mantle domain with superchondritic Sm/Nd and Lu/Hf ratios that developed positive e(Nd,Hf), and heterogeneous addition of crustal material with subchondritic Sm/Nd and Lu/Hf ratios and negative initial e(Nd,Hf). This crustal addition is distinct from that previously determined for the Pilgangoora basalts with positive initial e(Nd, Hf). A depleted mantle domain with relatively large positive initial e(Nd, Hf) at ~3.5 Ga may have resulted from either typical melting of primitive mantle before 4.0 Ga or a later event with greater incompatible element depletion (Fig. 6.6). The trace element abundances of the least contaminated Pilgangoora basalts indicate that they were derived from a mantle source as depleted in incompatible elements as the present-day MORB source (Chapter 4.7.5), which all but precludes derivation from a mantle domain formed by relatively greater incompatible element depletion. Consequently, the depleted mantle domain sampled by the Pilgangoora basalts probably formed before ~4.0 Ga and was isolated for a prolonged period, enabling positive initial e(Nd, Hf) to develop by ~3.5 Ga. The component with negative initial e(Nd, Hf) is consistent with addition of basaltic or sialic crust to a primitive or depleted mantle domain prior to or during ~3.5 Ga volcanism. For basaltic crust, prolonged isolation is required to produce large negative e(Nd, Hf), whereas the lower Sm/Nd and Lu/Hf ratios of sialic crust necessitate a shorter isolation period (Fig. 6.6). Moreover, at equivalent isotopic compositions, a greater volume of basaltic crust needs to be added to produce the same isotopic shift as sialic crust, because basaltic crust generally has a lower incompatible element content. These two processes counterbalance each other, and as the Pilgangoora basalts also have a significant continental component, discriminating between such subtle additions to the mantle source is not possible. Consequently, the timing of crustal addition is also difficult to constrain. However, significant involvement of primitive mantle can be CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 178 precluded because both basaltic and sialic crust would have been enriched in incompatible elements relative to primitive mantle, and thus crustal addition would have produced a mantle domain with superchondritic elemental abundances, not the depleted compositions sampled by the Pilgangoora basalts. Therefore, the Pilgangoora basalts were derived from a mantle source that was probably depleted before 4.0 Ga, isolated for a period sufficient to develop significantly positive initial e(Nd, Hf), and then mixed with basaltic or sialic crust with negative initial e(Nd, Hf) prior to or during ~3.5 Ga volcanism. Variable mixing produced the observed isotopic heterogeneities.

6.5.4.2 Crustal component The dominant cru stal component in the Pilgangoora basalts had positive initial e(Nd, Hf) and can be modelled with a minimum of three steps. The first step involved melting of primitive mantle to produce depleted mantle with superchondritic Sm/Nd and Lu/Hf ratios, and thus future development of positive e(Nd, Hf). The timing of mantle depletion cannot be constrained, but to remain consistent with the above mantle component it was probably before 4.0 Ga. After the depleted mantle had developed large positive e(Nd, Hf) it was partially melted to form basaltic crust, which was in turn fused to form silicic magmas, as outlined for the petrogenesis of second generation melts (Chapter 5.3.2.1). Since the basaltic crust probably had subchondritic Sm/Nd and Lu/Hf ratios then it must have been fused relatively soon after formation to avoid developing low e(Nd, Hf). The production of silicic magmas may have been contemporary with Pilgangoora volcanism, indicating that the crustal component was added to the mantle from which the basalts were derived, or they may have formed a sialic basement that was assimilated by the basalts. The isotopic data cannot discriminate between these two alternatives. However, the latter requires that the crust was not much older than the Pilgangoora basalts, otherwise the crustal component would have developed significantly lower initial e(Nd, Hf). CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 179

6.5.4.3 Combined Nd-Hf system There is a profound positive correlation between the initial Nd and Hf isotopic composition of most terrestrial rocks, reflecting the general covariant behaviour of Sm- Nd and Lu-Hf in both the mantle and the crust (Patchett & Tatsumoto, 1980; Salters & Hart, 1991; Vervoort et al., 1996, 1999a, 2000; Nowell et al., 1998; Salters & White, 1998; Pearce et al., 1999). Importantly, no evidence has yet been forthcoming to demonstrate a significant secular variation of the terrestrial mantle-crust array, and so the approximate relationship of e(Hf) = 2.0 + 1.4e(Nd) is considered valid for all post- 3.5 Ga terrains (Vervoort et al., 1996, 1999a; Vervoort & Blichert-Toft, 1999). The pronounced displacement of the terrestrial array to greater e(Hf) relative to the chondritic reference (CHUR, BSE) is a feature evident in some of the earliest terrestrial rocks, and an as yet unidentified reservoir has been proposed to complement this significant isotopic shift (Blichert-Toft & Albarède, 1997; Albarède et al., 2000; Vervoort et al., 2000). Alternatively, the present Nd-Hf values for the BSE may be incorrect, and so the terrestrial array may include CHUR, thus avoiding the need for an extra hidden reservoir (Salters & White, 1998). This is discussed further in Chapter 9. The Pilgangoora basalts have smaller initial e(Hf) for any given initial e(Nd) compared with the terrestrial mantle-crust array (Fig. 6.7). Thus, compared with most terrestrial systems, the Pilgangoora basalts record a history with relatively greater growth of radiogenic Nd and indicate decoupling of the generally coherent Nd-Hf system. Such decoupling is most readily explained by garnet fractionation, because Lu is strongly compatible and Sm, Nd and Hf are highly incompatible in garnet, but not for other common rock-forming minerals (Irving & Frey, 1978; Thirlwall et al., 1994; Vervoort & Patchett, 1996; Vervoort et al., 1999a, 2000). Hence, melting in the presence of garnet or crystallisation of garnet from a magma will leave restites with greater Lu/Hf ratios, magmas with smaller Lu/Hf ratios and there will be no corresponding change to the Sm/Nd ratios. As a result, e(Hf) will increase faster in the restite and slower in the melt compared with the typical e(Nd) relationship (Fig. 6.7). CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 180

30 10

20 residue (>Lu/Hf)

5 10 ) ) f f 0 0 H H ( ( e e melt (

-5 -20

-30 -10 -20 -15 -10 -5 0 5 10 15 -5 0 5 10 e(Nd) e(Nd)

Figure 6.7: e(Hf) versus e(Nd) diagrams for Pilgangoora basalts, plotted against a background showing the terrestrial crust-mantle array (see text for references). The magnified view shows the time-integrated evolution of each basalt for 500 myr after formation.

The Pilgangoora basalts show two superimposed trends in initial Nd-Hf space: one approximately orthogonal and the other subparallel to the terrestrial array (Fig. 6.7). The orthogonal trend is related to Nd-Hf decoupling, whereas the parallel trend is consistent with typical mantle-crust behaviour. It is proposed, therefore, that the variations of initial Nd-Hf are the result of mixing where one of the components, at least in part, was derived from magma generated by melting with residual garnet. In general, this model is consistent with straightforward crustal contamination, where the crust was formed by melting of a garnet-bearing source, such as for trondhjemites and tonalites, and there was enough time for divergent e(Hf) to develop. For the Pilgangoora basalts, however, the dominant crustal component had greater initial e(Nd, Hf) than the mantle (Fig. 6.4, 5), and so the observed Nd-Hf decoupling with pronounced negative initial e(Hf) cannot be related to this crustal component. Instead, the decoupling must be related to the subtle crustal component in the mantle source, as previously described (Chapter 6.5.4.1). It follows that this crust formed by melting of a garnet source, although it is not possible to determine whether this crust was basaltic or sialic as both commonly involve garnet fractionation (Salters & Hart, 1989; Johnson et al., 1990; Hirschmann & Stolper, 1996; Arth & Barker, 1976; Winther, 1996; Springer & Seck, 1997). Importantly, this crustal component has the required Nd-Hf composition of the ‘hidden reservoir’ needed to explain the deviation of the terrestrial mantle-crust array from the bulk silicate Earth (Hf-Nd-paradox), a point discussed in Chapter 9. Moreover, the Nd-Hf evolution vectors for the Pilgangoora basalts are subparallel to the terrestrial mantle-crust array (Fig. 6.8), and thus, apart from the initial e(Nd, Hf), the basalts CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 181 themselves show little evidence of garnet fractionation, even though they probably contain two added crustal components, both evidently derived from garnet fractionation. The vectors show the dominance of typical Nd-Hf behaviour and its potential to modulate the garnet effect. In summary, the initial Nd-Hf isotopic compositions of the ~3.5 Ga Pilgangoora basalts are consistent with derivation from a heterogeneous mantle comprising both depleted (positive e(Nd, Hf)) and enriched reservoirs (negative e(Nd, Hf)), and then addition of another isotopically distinct crustal component derived from a depleted source.

6.5.5 Origin of the granitoids The ~3479 Ma Carlindi trondhjemite and ~3468 Ma Carlindi monzogranite have initial e(Nd) = 1.68 and 1.26, respectively, whereas the gneissic xenoliths hosted by the ~2925 Ma monzogranite have initial e(Nd) = 2.34 to 2.43 for their interpreted ages. Hence, these granitoids were derived from sources with greater initial 143Nd/144Nd than CHUR, and so their sources must have evolved with superchondritic Sm/Nd ratios. However, the Carlindi granitoids have extremely fractionated REE, with profoundly subchondritic Sm/Nd ratios, such that growth of 143Nd/144Nd must have been significantly slower than for CHUR (Fig. 6.6). Hence, granitoid formation must have been associated with a marked change of LREE composition from positive to negative e(Nd) evolution (Fig. 6.6). Such a change is consistent with second generation melting (Chapter 5.3.2.1), where basaltic crust is melted with a garnet amphibolite residue. It follows that the Nd composition of the older Carlindi granitoids can be explained with a minimum of three steps. As for the Pilgangoora basalts, the first step involved partial melting of primitive mantle to form a depleted mantle domain with superchondritic Sm/Nd ratios and growth of positive e(Nd) (Fig. 6.6). Growth rates for this depleted mantle must have been faster than indicated in Figure 6.6 to develop e(Nd) > 2.4 by ~3.5 Ga. Moreover, the depleted mantle must have been isolated for a long enough period to allow such isotopic growth. This depleted mantle was melted to form a basaltic crust with positive initial e(Nd) and, in turn and relatively soon afterwards, this crust was fused to form REE fractionated granitoids, also with positive initial e(Nd). CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 182

It was previously shown, however, that some of the features in the Carlindi granitoids are more consistent with formation by fusion of sialic crust derived by the above processes. For this to be plausible, the sialic crust must have been relatively young when melted, otherwise the initial Nd isotopic composition would have been greatly reduced. The timing of such events is discussed in Chapter 6.6. The ~2925 Ma Carlindi granite has initial e(Nd) = -1.04, significantly lower than the other Carlindi granitoids, and thus derived from a source with a time-integrated subchondritic Sm/Nd ratio. Such a model is consistent with that proposed in Chapter 5 where the younger granitoid was derived from fusion of earlier Pilbara crust. However, the granite lies above the time-integrated trajectories of the other Carlindi granitoids at ~2925 Ma, and so cannot be solely derived from these sources (Fig. 6.7). A less evolved source, at least in part, is required. Interestingly, the granite lies within the evolution field defined by the Pilgangoora basalts (Fig. 6.7), which would support a model whereby the granite formed by melting of these basalts. Alternatively, some basaltic crust may have contributed to the fusion of the granitoids, as demonstrated by the inclusion of greenstone xenoliths, or an external mantle source may have added a depleted signature during crustal fusion. This latter idea is further examined in the context of all the published Pilbara isotopic data (see Chapter 6.6). In summary, the older Carlindi granitoids were derived from a source with positive initial e(Nd), and thus time-integrated superchondritic Sm/Nd ratios. The simplest model to account for their composition involves fusion of basaltic crust derived from a previously depleted mantle source, consistent with general models for HREE- depleted granitoids. However, previous considerations suggest that the granitoids formed by fusion of sialic crust, and for this to be true, the crust must have been very young when melted. The ~2925 Ma Carlindi granite has a negative initial e(Nd), and so was derived from a source with time-integrated subchondritic Sm/Nd ratios, perhaps by melting older crust. However, it cannot have been derived solely by fusion of Carlindi granitoids, requiring addition of a more depleted component such as ancient greenstones or even depleted mantle. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 183

6.6 DISCUSSION The following discussion consid ers previous Nd-Hf isotope studies in the Pilbara Craton to elucidate the evolution of regional isotopic domains. This is then used as a base to constrain the isotopic origins of the Pilgangoora basalts and the Carlindi granitoids.

6.6.1 Pilbara evolution Previous Nd isotopic studies in the Pilbara Craton have focussed on four terrains: the ~3.70 Ga Warrawoona Group in the Marble Bar Belt (Hamilton et al., 1981; Jahn et al., 1981; Gruau et al., 1987), granitoids from the Shaw complex (Bickle et al., 1989, 1993) and the west Pilbara (Smith et al., 1998) and basalts and andesites from the <2.76 Ma Fortescue Group (Nelson et al., 1992). The initial e(Nd) of the Pilbara samples are shown in Figure 6.8, where the Pilgangoora basalts are represented as an evolution field and time-integrated trajectories for the Carlindi granitoids and predicted mantle left after extraction of the Pilgangoora basalts are shown. All samples have been plotted for their reported zircon U-Pb ages, except for the Shaw granitoids where whole-rock Pb-Pb methods were used (Bickle et al., 1989, 1993). However, recent zircon U-Pb studies in the Shaw complex have shown a similar age diversity to those obtained by Pb-Pb methods, although they have not sampled the same outcrops (McNaughton et al., 1993; Zegers, 1996; Nelson, 2000). Where required, the initial e(Nd) values have been recalculated with the constants and CHUR values in Appendix 3. Studies in the Marble Bar Belt were primarily undertaken to estimate the age of volcanism and include Sm-Nd analyses from felsic, mafic and ultramafic units (Hamilton et al., 1981; Jahn et al., 1981; Gruau et al., 1987). SHRIMP U-Pb zircon techniques have since been used to determine the precise ages of various felsic units within the succession, all of which are between 3466 and 3477 Ma (McNaughton et al., 1993; Nelson, 1998, 1999, 2000). Thus, given an age of 3470 Ma, the initial e(Nd) compositions are calculated to be between -0.18 and +4.75 (Fig. 6.8). The ultramafic units have the largest initial e(Nd) whereas the felsic volcanics have significantly lower initial e(Nd). This relationship is reflected in the strong positive correlation between initial e(Nd) and the Sm/Nd ratios (Fig. 6.9A), consistent with mixing between mantle sources with large positive initial e(Nd) and chemically evolved sources with low CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 184 positive to negative initial e(Nd). This is a typical crustal contamination trend, where the mantle was depleted and so developed superchondritic Sm/Nd ratios and the crust developed subchondritic Sm/Nd ratios. Unfortunately, elemental analyses of the high field strength elements are not precise or too few (see Chapter 4.8.2), and so cannot be used as for the Pilgangoora basalts. Nevertheless, the mantle source probably had initial e(Nd) = 2.5 to 5.0, whereas the crustal component had initial e(Nd) = -0.5 to 1.5. Interestingly, isotopic divergence between coeval ultramafic and mafic units in other Archaean terrains has been interpreted to reflect derivation from distinct mantle domains (Campbell et al., 1989; Campbell & Griffiths, 1992). The Marble Bar ultramafic units are profoundly depleted in Al 2O3, perhaps also suggesting a discrete mantle source.

However, one Al2O3-depleted sample has an enriched e(Nd) composition, indicating that such a model is not straightforward.

4.0 ?mantle

2.0

Pilgangoora basalts 0.0

-2.0 ) d N (

e -4.0

-6.0 Fortescue Group Carlindi granitoids, lines are evolutionary trends Marble Bar greenstones -8.0 Coonterunah ?contaminant Shaw granitoid

west Pilbara granitoid Marble Bar ?contaminant

2700 2800 2900 3000 3100 3200 3300 3400 3500 3600 Age (Ma)

Figure 6.8: Initial e(Nd) versus age for Pilbara samples. Shaded field shows the evolution of the Pilgangoora basalts. Time-integrated trajectories for Carlindi granitoids and assumed growth of the mantle from which the Pilgangoora basalts were extracted. See text for age and reference details.

A diverse suite of Shaw granitoids were analysed to investigate their isotopic evolution and constrain their petrogenesis (Bickle et al., 1989, 1993). They represent a broad age range and define a general trend where the younger granitoids have lower initial e(Nd) (Fig. 6.8). The oldest Shaw granitoids are broadly contemporary with the Carlindi granitoids (~3470 Ma) and have similar initial e(Nd) (+1.02 to +1.92). The CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 185 younger granitoids have initial e(Nd) that are generally near to the time-integrated trajectories of the ~3470 Ma Shaw and Carlindi granitoids, suggesting that they may have formed by fusion of this older crust. However, three granitoids are below these trajectories indicating a derivation from sources with even lower time-integrated e(Nd).

0.5 0.5 A - Marble Bar Belt B - Fortescue Group 0.4 0.4 d 0.3 d 0.3 N N / / m 0.2 m 0.2 S S Hamilton et al., 1981 0.1 0.1 Jahn et al., 1981 Gruau et al., 1987 0.0 0.0 -1 0 1 2 3 4 5 -5 -4 -3 -2 -1 e(Nd) e(Nd) Figure 6.9: Sm/Nd versus e(Nd) for volcanic rocks from the ~3470 Ma Marble Bar Belt (reference in figure) and the <2.76 Ga Fortescue Group (Nelson et al., 1992).

In the west Pilbara, the Sm-Nd isotopic compositions were determined for eight granitoids adjacent to and within the Sholl Shear Zone (Smith et al., 1998). The study showed that the shear zone represents an important geochemical boundary with distinct initial e(Nd) either side. The west Pilbara granitoids generally plot above the time- integrated trajectories of the older Carlindi granitoids (Fig. 6.8), and hence they were derived from sources with larger initial e(Nd). Therefore, they cannot have been derived solely by fusion of this earlier Pilbara crust, and probably represent, at least in part, juvenile additions to the Pilbara. It is possible that these granitoids represent isotopic mixtures between the Pilgangoora mantle and older granitic crust. The <2.76 Ga Fortescue Group unconformably overlies the Pilbara granite- greenstone terrain and covers a large geographical area. Basalts and andesites from a broad geographic and stratigraphic range in the Fortescue Group have initial e(Nd) between -1.82 to -4.76 (Nelson et al., 1992). Importantly, there is no correlation between initial e(Nd) and the Sm/Nd ratio, and thus there is no evidence that the basalts assimilated crustal material (Fig. 6.9B). This is remarkable considering that the Fortescue Group erupted onto the Pilbara Craton, and thus the magmas passed through a thick continental basement. Nevertheless, the Fortescue Group must have been derived from a mantle source with a time-integrated subchondritic Sm/Nd ratio, which is CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 186 consistent with their reported REE compositions (Nelson et al., 1992). The Fortescue volcanics could not have been derived from the same mantle as the Pilgangoora basalts, but they could have incorporated older Pilbara crust with similar initial e(Nd), particularly the west Pilbara granitoids (Fig. 6.8). Using single-grain TIMS analytical methods, Amelin et al. (2000) obtained the Lu-Hf isotopic compositions and corresponding U-Pb ages of 8 zircons from three separate Pilbara felsic units. These data show no evidence of Hf or Pb mobility and so are considered reliable. However, all the units were described as volcanic, although the Miralga Creek and possibly the West Bamboo porphyries are intrusive, and therefore they may not represent the greenstone successions as intended. Nevertheless, all the zircons have positive initial e(Hf), indicating that they were derived from sources that had time-integrated histories with positive Lu/Hf ratios. Hence, they were derived from depleted sources, which is consistent with initial e(Nd) of equivalent-aged felsic samples, and indeed basalts, in the Pilbara Craton.

6.6.2 Geological history Correlations between initial e(Nd, Hf) and certain trace-element ratios provide evidence for crustal contamination in the ~3.5 Ga Pilgangoora basalts, adding support to an earlier supposition based on their elemental abundances alone (Chapter 4.7). The origins of such contaminants have not been conclusively established, but may include either mantle metasomatism during subduction-like processes or assimilation of continental basement. Importantly, the Warrawoona basalts were erupted onto a basement composed of the Carlindi granitoids, and so these granitoids may have been incorporated by the Warrawoona magmas. The initial e(Nd) of the Carlindi granitoids and gneissic xenoliths are equal to or larger than the largest initial e(Nd) of the Warrawoona basalts (Fig. 6.3, 8), and so addition of these granitoids and gneisses would increase the initial e(Nd) of the Warrawoona basalts, as required. Therefore, the Nd- isotopes are consistent with the Warrawoona basalts assimilating Carlindi granitoids. Given that the Warrawoona basalts assimilated sialic basement, then their geochemical similarities with the Coonterunah basalts suggest that assimilation may have occurred in these basalts, too. The e(Nd) compositions of the required contaminant at 3517 and 3600 Ma are shown in Figure 6.8. Two important points can be made: the contaminant at 3517 Ma coincides with the time-integrated e(Nd) of the Carlindi CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 187 granitoids and gneissic xenoliths and at 3600 Ma the contaminant is within the range of the most depleted part of the Marble Bar mantle source. Therefore, crustal assimilation is plausible for the Coonterunah basalts based on the known Pilbara isotopic constraints, but probably only if the crust had a similar composition to the Carlindi granitoids and separated from a depleted mantle after ~3600 Ma. It follows that this pre-Coonterunah basement must have been young when assimilated. The crustal component in the Marble Bar basalts had negative initial e(Nd), and so was isotopically distinct from that incorporated by the Pilgangoora basalts (Fig. 6.8). If both of these contaminants represent basement, then the early Pilbara crust had marked isotopic variations. Moreover, if the proposed Marble Bar basement was composed of typical early Pilbara granitoids, then its lower initial e(Nd) suggests a profoundly older basement. The formation of pre-Coonterunah sialic crust has important implications for the origins of the Carlindi granitoids. As both have equivalent time-integrated e(Nd), even the oldest Carlindi and Shaw granitoids may have been derived from this earlier crust. If so, this crust must also have been relatively young when fused. The initial e(Nd) of many younger Pilbara granitoids can be modelled as recycled older Pilbara granitoids. Three of the Shaw granitoids had initial e(Nd) that were close to the time-integrated trajectories of the proposed Marble Bar basement, providing additional support for the existence of such a basement. In the west Pilbara, however, the initial e(Nd) of the granitoids suggest that they represent more juvenile additions to the Pilbara crust, although they may still have a incorporated some older sialic crust. The ~2925 Ma Carlindi monzogranite had similar initial e(Nd) as the west Pilbara granitoids, although it also contains gneissic xenoliths and zircon xenocrysts indicating a significant contribution from older crust. The origin of the depleted component in the monzogranite is not constrained, but may represent a mantle signature. The Pilgangoora and Marbl e Bar basalts were derived from a heterogeneous mantle with highly depleted and slightly enriched components. Hence, geochemical domains had been differentiated and isolated for long periods by ~3.5 Ga. Moreover, there is evidence that enriched material had already been recycled back into the mantle. The initial e(Nd) of the Fortescue basalts also supports the recycling of older enriched material into the mantle during the Archaean. CHAPTER 6: RADIOGENIC ISOTOPE GEOCHEMISTRY 188

6.7 SUMMARY Nd-Hf isotopic analyses of the Coonterunah and Warrawoona basalts show that they were derived from sources with diverse time-integrated Sm/Nd and Lu/Hf ratios. Covariations between initial e(Nd, Hf) and certain trace element ratios show that both basalt suites represent mixtures of mantle and crustal material, where the crustal component had greater initial e(Nd, Hf) than the mantle sources. This crustal component was probably juvenile sialic crust which had an initial isotopic composition consistent with the Carlindi granitoids or older equivalents. The mantle sources were heterogeneous, but had significant depleted components that probably formed before 4.0 Ga. The older Carlindi granitoids were derived from sources with time-integrated superchondritic Sm/Nd ratios, either by fusion of basaltic crust derived from previously depleted mantle or by partial melting of juvenile sialic crust. The younger Pilbara granitoids were largely formed by fusion of older crust. CHAPTER 7: DISCUSSION- GEOLOGICAL SYNTHESIS 189

Chapter 7: DISCUSSION – GEOLOGICAL SYNTHESIS

7.1 INTRODUCTION Greenstone belts form some of the oldest extant parts of the Pilbara Craton and thus provide important constraints about the early history of this continental fragment. Of particular interest is whether the greenstones formed in pre-existing continental domains, as this would have profoundly influenced their petrogenesis and provide constraints about certain ancient tectonic processes. Thus, the evidence that the Coonterunah and Warrawoona successions were deposited onto continental basement is discussed here. Of particular interest is the composition and fate of such a basement, if it existed. In Chapter 8, the tectonic setting of the early Pilbara is elucidated, and in Chapter 9, the processes and rates of early continental growth are constrained.

7.2 CONTINENTAL BASEMENT Continental crust has a mean andesitic bulk composition and is predominantly composed of rocks that are more siliceous and enriched in incompatible elements than basalt (Taylor & McLennan, 1985, 1995; Rudnick & Fountain, 1995). In contrast, oceanic crust is entirely composed of basaltic rocks. Since there are many continental regions where basalts form significant fractions of the crust, either actually deposited in continental settings or in oceanic fragments incorporated into the continents, the presence of basalt alone is not indicative of basement type. In some instances, the basal contact of a basalt succession may be preserved, and in such circumstances there is generally little doubt regarding the composition of the basement. In the absence of such a contact, however, geochemical, geological and geophysical features may provide evidence of basement composition. Xenoliths may directly sample the basement, although evolved rocks are not unknown in oceanic domains and primitive rocks can be found in continental settings. The presence of zircon xenocrysts in basalts have been used to imply that the magmas traversed through continental basement (Compston et al., 1986). Crustal contamination is particularly common in continental settings and is typically evident as enrichment of incompatible elements in basalts. However, such features are not exclusive to continental basalts, and there are many modern oceanic domains where basalts have been similarly enriched by addition of crustal material to the mantle (Hofmann, 1997). CHAPTER 7: DISCUSSION- GEOLOGICAL SYNTHESIS 190

Continental crust is typically buoyant and so continental basalts are likely to be deposited in subaerial to shallow subaqueous setting. However, such features are not exclusive to continental settings as clearly demonstrated by modern ocean island volcanism. Moreover, basalts may be deposited onto deeply submerged thin continental fragments (e.g., early Kerguelen Plateau, Storey et al., 1989; Frey et al., 2000). Continental crust is also cool and rigid and so basalt successions may be preserved relatively intact and with minimal metamorphism. In contrast, oceanic crust typically subsides as it cools and so needs to be incorporated into more buoyant continental crust to facilitate its preservation, thus leading to significant tectonism and metamorphism. Moreover, oceanic domains typically have greater geothermal gradients than continental domains, and so metamorphic gradients through volcanic piles can be markedly different. Nevertheless, in the absence of direct evidence and despite all these indicators it is difficult to be certain about the composition of the basement.

7.3 PRESENCE OF EARLY PILBARA CRUST The basal contacts of most early Pilbara greenstone successions are either intrusive or tectonic and so the basements onto which these successions were deposited cannot be directly observed. In the Pilgangoora Belt, however, the upper units of the regional Warrawoona succession unconformably overlie ~3517 Ma Coonterunah greenstones and ~3485-3468 Ma Carlindi granitoids. The unconformity is angular to semi-conformable and sand-filled fissures, ancient topography and a putative palaeosol at the contact indicate that it formed by subaerial erosion of a deformed terrain. Indeed, it is likely that at least 2 km of the Coonterunah-Carlindi basement was removed prior to Warrawoona deposition. In the Pilgangoora Belt, therefore, the Warrawoona succession was deposited onto an emergent basement composed of greenstones and granitoids. Clastic sediments within the SPC-sandstone, the basal Warrawoona unit in the Pilgangoora Belt, also reflect the composition of this early basement. The age distribution of detrital zircons shows that these sediments were largely derived from the immediately underlying Coonterunah-Carlindi terrain. Moreover, the SPC-sandstone is predominantly composed of well-rounded, medium- to coarse-grained quartz that was probably derived from the Carlindi granitoids, because the quartz within the Coonterunah basalts and felsic volcanics is somewhat finer. The abundance and type of inclusions within the quartz and the presence of detrital zircon, monazite and cassiterite CHAPTER 7: DISCUSSION- GEOLOGICAL SYNTHESIS 191 is consistent with a granitic provenance. Thus, the area of Carlindi granitoids exposed to erosion was probably greater than presently represented along the unconformity. In the Pilgangoora Belt, primary features in the Warrawoona succession clearly show that deposition occurred in a shallow-water environment, and so crustal buoyancy must have been maintained during deposition. Moreover, the Warrawoona succession was subjected to only prehnite-pumpellyite to lowermost greenschist-facies metamorphism with minimal deformation, indicating that it escaped major uplift and erosion and also deep burial and high-grade metamorphism. Thus, the Warrawoona succession was deposited onto chemically evolved silicic basement that subsequently remained relatively rigid, cool and buoyant, by all definitions continental crust (Buick et al., 1995). The Warrawoona successions in the Coongan, North Shaw and Kelly Belts and North Pole and McPhee Domes were also deposited in subaerial to shallow subaqueous settings and have also experienced only low-grade metamorphism and minimal deformation (Chapter 2.9.2; Lowe, 1983; Barley et al., 1984; Barley, 1993; Buick & Dunlop, 1990). It is thus likely that these Pilbara terrains were also deposited onto rigid, cool and buoyant basements, although their basal contacts cannot be directly observed. The Warrawoona basalts from th e Pilgangoora, Marble Bar and Kelly Belts and North Pole and McPhee Domes have remarkably similar trace element abundances (Chapter 4.8.2), so they must have been formed by equivalent petrogenetic processes. These basalts have crustal signatures compared with N-MORB, reflecting the addition of a crustal component to a mantle-derived magma either during subduction-related metasomatism or by assimilation of crustal material. Trace element modelling(shows that the Pilgangoora Warrawoona basalts can be readily explained by <10 % assimilation of Carlindi granitoids and subsequent fractional crystallisation (Chapter 4.7). Moreover, Nd-Hf isotope systematics confirm the assimilation of the Carlindi granitoids by the Pilgangoora basalts (Chapter 6.5). Thus, the crustal signature of the Warrawoona basalts in the Pilgangoora Belt and, by extrapolation, to elsewhere in the Pilbara Craton, is most likely a result of assimilation of silicic crust. That such assimilation occurred elsewhere in the Pilbara is supported by the Nd isotope composition of the Marble Bar greenstones (Chapter 6.6.1) and by the common presence of zircon xenocrysts in Warrawoona felsic volcanics in the Coongan, Marble Bar and Kelly Belts and North CHAPTER 7: DISCUSSION- GEOLOGICAL SYNTHESIS 192

Pole and McPhee Domes (Table 3.4; Thorpe et al., 1992a; Barley et al., 1998; Nelson, 2000). Importantly, the ~3517 Ma Coonterunah basalts have equivalent compositions to their Warrawoona counterparts, indicating that they too formed by similar processes. Hence, the Coonterunah basalts evidently also assimilated older silicic crust, and thus were deposited onto continental basement, although this basement cannot now be observed. However, granitic ènclaves preserved within Coonterunah basalts may provide direct evidence of assimilation of this earlier basement, although zircons within these ènclaves have apparently been recrystallised. Therefore, the Coonterunah and Warrawoona successions were probably both deposited onto older continental basement, although direct evidence of this basement has only been noted in the Pilgangoora Belt. Nevertheless, an array of indirect evidence suggests that such basement was widespread and played a crucial role in the petrogenesis of the greenstone successions.

7.4 COMPOSITION OF EARLY PILBARA CRUST The Pilgangoora Belt cont ains the only known exposure of a basal greenstone contact, indicating that the Warrawoona basement was composed of older greenstones (Coonterunah Group) and granitoids (Carlindi). Many of the Warrawoona successions elsewhere in the Pilbara probably had basements with equivalent compositions as they were likewise deposited onto cool, rigid, stable and buoyant basements and have been similarly contaminated. Thus, the Coonterunah-Carlindi terrain probably provides a fair estimate of the general Warrawoona basement. The Coonterunah succession, however, must have been deposited onto even older basement. The Nd-Hf isotopes of the Pilgangoora basalts suggest that they assimilated rather young (<100 myr old) sialic crust. In contrast, Nd isotopes of the Warrawoona succession in the Marble Bar Belt indicate that they incorporated somewhat older crust (Fig. 6.8-9), and thus the pre-Coonterunah crust was probably heterogeneous. Indeed, the great diversity of zircon xenocrysts (Table 3.4) and detrital zircons (Fig. 3.21) from various Pilbara granitoids and greenstone belts supports such basement heterogeneity. Two xenoliths in granitoids provide direct evidence of the composition of the pre-Coonterunah basement (Table 3.4). In the Warrawagine Granitoid Complex, an CHAPTER 7: DISCUSSION- GEOLOGICAL SYNTHESIS 193 intensely deformed £3658 Ma banded gneissic xenolith is hosted by ~3244 Ma granodiorite (Nelson, 1999). The precursor of this xenolith may have been a porphyritic tonalite. Furthermore, the diverse age range of concordant zircons in the xenolith indicates that it experienced a number of pre-3.5 Ga tectonothermal events. A ~3578 Ma gabbroic anorthosite xenolith has been described from Shaw Granitoid Complex (Daltons Pluton), perhaps of plutonic derivation (McNaughton et al., 1988). Thus, the pre-Coonterunah basement was temporally and compositionally heterogeneous and probably experienced many significant tectonothermal events in its early history.

7.5 FATE OF EARLY PILBARA CRUST Pre-Warrawoona basement was evidently widespread in the Pilbara though there are few surviving relics. Consequently, the early Pilbara crust must have been somehow removed or recycled. Large-scale post early Archaean erosion seems unlikely because most of the Warrawoona successions have been relatively stable and buoyant since deposition, which is generally incompatible with the active tectonics required to form significant erosive regimes. Nevertheless, there is strong evidence in the Pilgangoora Belt that >2 km of the Coonterunah-Carlindi basement was removed prior to Warrawoona deposition. The resulting clastic unit is <5 m thick and so most of the sediment must have been removed from the immediate depository. In addition, Warrawoona and post-Warrawoona sediments contain abundant detrital zircons that were derived from Coonterunah-Carlindi and pre-Coonterunah crust, but this can only account for a small part of the early crust. Some of the early basement was incorporated by the Coonterunah and Warrawoona basalts, but the volumes involved were rather insignificant (basalt pile <10 km thick with <10 % contamination implies assimilation of <1 km continental crust). Though it is possible that material was removed from the base of the crust by thermal erosion or crustal delamination, it is unlikely that either could be responsible for such total obliteration of thick ancient continental basement without some resulting evidence. A more likely scenario is that the early crust was remelted to form granitoid complexes. Indeed, the younger Pilbara granitoids were clearly derived from melting of earlier crust as indicated by the presence of granitic xenoliths and zircon xenocrysts (Table 3.4), their trace element compositions (Chapter 5.4.1) and their Nd isotope compositions (Fig. 6.8). These younger granitoids probably formed by small degrees of partial melting at CHAPTER 7: DISCUSSION- GEOLOGICAL SYNTHESIS 194 relatively low temperatures and so have preserved xenoliths and xenocrysts and have been enriched in many incompatible elements (Chapter 5.3.2.2). In contrast, the older granitoids probably formed by large degrees of partial melting at high temperatures and so have not preserved xenoliths and xenocrysts, but have largely retained the composition of their source (Chapter 5.3.2.1). Therefore, most of the early Pilbara crust was probably consumed during intracrustal recycling, although some was clearly removed by erosion. Such recycling may have been important in other Archaean terrains and probably accounts for the general poor preservation of pre-3.5 Ga terrains in Archaean cratons globally.

7.6 SUMMARY The early Pilbara greenstones were deposited onto older continental basement, though basal contacts have generally been removed by later intrusions or tectonism. This early basement was temporally and compositionally heterogeneous and was probably largely recycled to form younger granitoids. CHAPTER 8: DISCUSSION - TECTONIC SETTING 195

Chapter 8: DISCUSSION – TECTONIC SETTING

8.1 INTRODUCTION Deciphering ancient tectonic environments has relied heavily on their geochemical correlation with modern plate tectonic settings. In particular, discriminant diagrams, as pioneered by Pearce & Cann (1971, 1973), provide a simple method for assigning tectonic labels to geochemical data. Such diagrams are widely favoured because they purportedly resolve complicated problems simply, require little understanding of the geochemical processes that produce them and can be supported with a litany of references. Although over-reliance on such diagrams has been rightly challenged (Arculus, 1987; Duncan, 1987; Pearce, 1987), the diagrams are useful monitors of chemical fractionation processes. If fractionation can be uniquely assigned to a specific process, and, in turn, this process can be uniquely assigned to a specific modern tectonic setting, then such diagrams are indeed powerful tectonic indicators. However, such instances have not necessarily been demonstrated for all examples of a tectonic setting for all of time. For instance, although there are profound geochemical differences between island-arc and mid-ocean ridge basalts, such criteria may not discriminate between island-arc and back-arc basalts. It has been suggested, therefore, that a sequence of criteria or ‘screens’ be used to progressively discriminate (Condie, 1989), but again, these fail to separate representatives of all tectonic settings. In a review of commonly used discriminant diagrams, Wang & Glover (1992) showed that continental basalts are most commonly misidentified in these diagrams because they can be derived from variable degrees of partial melting at various depths from a number of possible mantle sources and variably mixed with a diverse range of crustal material. Although uniformitarian principles have provided a fundamental foundation for geological reasoning, there are many reasons to suspect that rates of processes and types of products during the Earth’s early history were considerably different to those prevailing at present. This does not necessarily imply that early processes were not generally analogous to modern tectonic processes (actualistic), though there were probably processes, particularly during the earliest times, for which there are no modern analogues (non-actualistic). Consequently, determining ancient tectonic settings is more difficult. For example, the gradual decrease of radioactivity within the Earth has probably caused secular cooling of the mantle heat flux, and so the temperature and CHAPTER 8: DISCUSSION - TECTONIC SETTING 196 depth of mantle melting have probably decreased with time. Hence, a hotter young Earth is commonly invoked to account for the prevalence of komatiites in the Archaean. In addition, profound chemical changes to the atmosphere, hydrosphere and biosphere have probably fundamentally changed the composition of sediments entering subduction zones and being recycled into the mantle. Such secular changes are not inherently evident in tectonic discrimination diagrams, mainly because their influence on tectonic processes is poorly constrained. Moreover, the sole use of chemical discriminants from one ancient terrain to define the tectonic setting of another terrain produces circular arguments. As a result of these and other more mundane issues, such as alteration intensit y and sampling consistency, geochemical discrimination diagrams are not used here to determine the tectonic setting of the Pilbara terrains. Instead, the tectonic setting of the Pilgangoora greenstones and Carlindi granitoids is deduced by integrating the geochemical features with geological observations. Most of the discriminating characteristics are direct observations from the Pilgangoora greenstones and Carlindi granitoids, but include additional data from other Pilbara terrains, particularly the North Pole Dome. Some derived constraints have been determined relating basalt composition, mantle source temperature, isostatic buoyancy, sea-level, crustal thickness and lithospheric stretching. Specific actualistic and non-actualistic settings are then tested against these characteristics to determine whether any one setting can best account for the tectonic evolution of the early Pilbara. Importantly, the similarity among basalts from all of the ancient Pilbara greenstones suggests a common origin.

8.2 DIRECT OBSERVATIONS 1) Eruption of thick basaltic successions. The Coonterunah succession is ~6.5 km thick, but was possibly >9 km thick before early Archaean erosion assuming that the felsic volcanics immediately beneath the unconformity in the central Pilgangoora Belt can be correlated with those to the east. This assumption is supported by their geochemical and age similarities. The Pilgangoora and North Pole Warrawoona successions are ~3.5 and ~10 km thick, respectively. The observed thicknesses represent minimum estimates because the tops of the successions and the bases of the Coonterunah and North Pole successions have been truncated. CHAPTER 8: DISCUSSION - TECTONIC SETTING 197

2) Bimodal volcanism. The greenstone successions are dominantly basaltic, but the Coonterunah and North Pole successions have several intercalated felsic units up to ~300 m thick. The trace element abundances of the Coonterunah felsic volcanics indicate that they formed by late stage differentiation of comagmatic basalts, possibly requiring extensive, shallow magma chambers. 3) Shallow-water deposition. Persistent shallow subaqueous deposition of the Warrawoona successions indicates, assuming constant sea-level, that buoyancy, subsidence and crustal addition were balanced (see derived constraints below). 4) Minor ultramafic magmatism. Several ultramafic units have been recorded in the Marble Bar Belt (Glikson & Hickman, 1981a, b; Gruau et al., 1987) and North Pole Dome (Green et al., in prep.). Many of these are probably semi-conformable sills, but some may be komatiite flows. They are generally composed of chlorite-talc-tremolite- carbonate and have poorly preserved primary features, thus complicating positive identification. Similar units in the western Pilgangoora Belt are clearly intrusive. The scarcity of widespread, extensive komatiites indicates that the Pilbara basalts were not derived from large degrees of partial melting, and so the mantle sources were generally not substantially hotter than for typical modern basalt sources. 5) Punctuated eruption of lava. The hiatus between the Coonterunah and Warrawoona Groups was between 55 and 80 myr long, with no profound change of basalt source or petrogenesis. Significant hiatuses in the North Pole succession are represented by the North Pole and Strelley Pool Chert, whereas minor hiatuses are indicated by thin chert horizons in all the successions. 6) Limited terrigenous clastic sedimentation. Excluding reworked juvenile volcaniclastics, detrital sediments are restricted to thin quartz-rich sandstones and conglomerates at the base of the Strelley Pool Chert. In the Pilgangoora Belt, the age of detrital zircons, the preservation of abundant euhedral grains and the variety of lithic clasts suggest local provenance. 7) Eruption of basalts onto continental basement. In the Pilgangoora Belt, the Warrawoona basalts were deposited onto the subaerially eroded Coonterunah-Carlindi basement. Primary features in the Warrawoona Group clearly show that deposition was in a shallow-water environment, and the succession was then subjected to only very low- grade metamorphism with minimal deformation. Therefore, the Warrawoona basalts were erupted onto chemically evolved silicic basement that subsequently remained CHAPTER 8: DISCUSSION - TECTONIC SETTING 198 relatively rigid, cool and buoyant, by all definitions continental crust (Buick et al., 1995). Importantly, the geochemical likeness between the Warrawoona and Coonterunah basalts indicates that they formed by similar processes, and thus, it is likely that the Coonterunah basalts were also erupted onto continental basement. Trace element and Nd-Hf isotopic modelling show that assimilation of Carlindi granitoids or something similar can account for the crustal geochemical signature in the Pilgangoora basalts. Similar reasoning can be applied to other Pilbara successions, where further evidence for continental basement has been recorded (see Chapter 7.3). 8) Partial melting of previously depleted mantle. The least contaminated basalts have very low incompatible element concentrations associated with high Nb/U ratios and positive initial e(Nd, Hf). These basalts have trace element compositions that approximate the modern MORB composition. 9) General uniform bulk composition of the mantle source. The elemental abundances of the basalts can be readily explained by varying degrees of partial melting, crustal assimilation and fractional crystallisation with a relatively constant starting composition. Hence, the mantle source probably had a relatively uniform bulk composition. Moreover, the geochemical similarities between the Coonterunah and Warrawoona basalts indicate that they were derived from mantle sources with remarkably similar compositions, although they erupted between 55 and 80 myr apart. The isotopic compositions of the basalts show that the mantle source was isotopically heterogeneous, but these variations are not reflected in their the bulk compositions. 10) Stratigraphic trends of basalt composition. The Coonterunah and North Pole basalts show profoundly different compositional variations with stratigraphic height. In the lower Coonterunah succession, the basalts have greater abundances of incompatible elements with increasing stratigraphic height, suggesting progressive enrichment by assimilation and fractional crystallisation in magma chambers. The North Pole basalts show the reverse relationship with lower abundances of incompatible elements with increasing stratigraphic height, indicating progressive addition of less- fractionated mantle-derived magmas. These opposing trends indicate that temporal variations were caused by late-stage magmatic processes and not systematic changes of mantle compositions. Importantly, there are no profound compositional variations between the Coonterunah and Warrawoona basalts across the Pilgangoora unconformity or between Warrawoona basalts from various Pilbara belts and domes. CHAPTER 8: DISCUSSION - TECTONIC SETTING 199

11) Timing of granitoid emplacement. The Carlindi granitoids intruded the Coonterunah Group and were overlain by the Warrawoona Group. There is no evidence for granitoid intrusion during the Coonterunah event, but there are granitoids in the North Pole Dome and Shaw and Muccan Granitoid Complexes that have ages broadly similar with the Warrawoona event (see Figure 3.18 and Table 3.3). 12) Uplift and erosion. The Coonterunah-Carlindi terrain was uplifted and eroded prior to deposition of the Warrawoona Group. Given the previously outlined correlation between the Coonterunah felsic volcanic units, then most of what now forms the eastern domain may have been removed from on top of the central domain. This implies that at least 2 km of greenstones were eroded prior to Warrawoona deposition. The ~3468 Ma Carlindi microgranite was also exposed during this erosional uplift event, indicating that denudation may have been quite rapid (between 2 and 34 myr depending on Warrawoona age). 13) Low-grade metamorphism. The Warrawoona volcanic piles have experienced sub- to lowermost-greenschist facies metamorphism. Such low metamorphic grades deep in a ~10 km volcanic pile indicate that there was little convective or conductive heat transfer through the pile during volcanism. 14) Minor regional deformation. Most of the Pilgangoora and North Pole greenstones and the Carlindi granitoids are extremely well preserved with only locally developed tectonic fabrics and structures. There is little evidence of syn- or early post- depositional deformation, except for intraformational folds within some Coonterunah chert horizons and the bedding discordance between the eastern Pilgangoora domain and the unconformity. These structures are consistent with doming and related block rotation (Chapter 2.7.1). All other structures are somewhat younger. For instance, the major Lalla Rookh Fault not only truncates the Coonterunah and Warrawoona successions but also displaces sediments from the ~2950 Ma Lalla Rookh Formation (Krapez, 1993). This fault, however, has probably been reactivated as fault-related Pb mineralisation has been dated at ~3.2 Ga (Richards et al., 1981; Thorpe et al, 1992b). The more intense deformation in the western part of the Pilgangoora Belt is confined to that area and has been dated at ~2.9 Ga (Neumayr et al., 1998). 15) Stress regime during deposition. In the Pilgangoora Belt, the SPC does not change thickness across faults and so there is no evidence that faulting was syn- depositional. By contrast, it has been proposed by Nijman et al. (1999) that there were CHAPTER 8: DISCUSSION - TECTONIC SETTING 200 up to ~200 m displacements on normal listric syn-depositional faults in the North Pole Dome indicating a tensional syn-depositional stress regime. However, the evidence is less than convincing. For instance, many listric forms have apparently been inferred from aerial photograph traces of intersecting normal faults and are not truly curved structures, the restored cross-section does not account for strike-slip displacements along faults though such motion was clearly documented, and thickness changes across faults are largely based on mafic units above the North Pole Chert which have been variably inflated by dolerite sills (Green et al., in prep.). Conclusive evidence of thickness changes of sedimentary units across faults has not been documented, and so it is therefore unclear if the syn-depositional stress regime was tensional or not.

8.3 DERIVED CHARACTERISTICS In addition to the direct observations, the composition of the Pilbara basal ts can be used to derive approximations of the temperature and depth of mantle melting, and modelling the balance between crustal addition and sea-level can also provide constraints for tectonic modelling. 1) Temperature. The maximum MgO content of the Pilgangoora and North Pole basalts is ~12 dry wt%, which is slightly greater than the maximum of modern MORB (MgO = ~11 %; Klein & Langmuir, 1987; McKenzie & Bickle, 1988). However, three processes might have influenced the Pilbara basalts such that their measured MgO contents may not be reliably compared with modern MORB: crustal contamination, olivine fractionation and post-magmatic alteration. The first two processes would have lowered MgO contents by adding low-MgO and subtracting high-MgO material, respectively. Hence, ~12 wt% may be a minimum estimate. As Nb/U ratios can be used to estimate the amount of crustal addition (Chapter 4.7.5), then straightforward assimilation of local granitoids by the most magnesian Pilbara basalts would have lowered their MgO contents by <0.5 %. Moreover, there is little evidence for significant involvement of olivine during the late-stage differentiation of the magmas, and so olivine fractionation probably had a minimal affect, although this is poorly constrained. Post- magmatic alteration of the Pilbara basalts may have increased or decreased their MgO contents, although deviations from a magmatic trend are minor (<2 wt%; Fig. 4.4). However, pre-alteration MgO contents significantly greater than 12 wt% are unlikely since the basalts were composed predominantly of plagioclase and pyroxene, not a CHAPTER 8: DISCUSSION - TECTONIC SETTING 201 particularly MgO-rich assemblage. As a result, ~12 wt% MgO is probably a reliable maximum estimate for the Pilbara basalts.

40%

30% ) 0.3 X (

t l e m

f 0.2 o

n o i t c

a 0.1 1 1 20%

r 1 10% 2 3 4 F 8 8 8 0 0 ° 0 ° ° C C C

0 0 50 100 Melting depth (km)

Figure 8.1: Fraction of melt as a function of melting depth with contours of MgO content of a melt derived from adiabatic melting and potential temperature (McKenzie & Bickle, 1988). The x-axis is essentially a solidus. Shaded area indicates the stability of the Pilbara basalts.

If the Pilbara basalts had a maximum MgO content slightly greater than the maximum of modern MORB and greater MgO contents in basalts are the result of larger degrees of mantle melting at higher temperatures (Green & Ringwood, 1967), then the Pilbara basalts were probably derived from a slightly hotter mantle source than modern

MORB. This can be quantified as potential temperature (T P), that is, the temperature the mantle would have if it moved adiabatically to the Earth’s surface without melting (White et al., 1987). Hence, potential temperature allows comparison of magmas generated at various mantle depths. Using an experimentally derived relationship between TP and MgO (Fig. 8.1; McKenzie & Bickle, 1988), the maximum potential temperature for the Pilbara basalts is ~1380° C, which is ~100° C hotter than for modern MORB (TP = ~1280° C; McKenzie & Bickle, 1988). This difference is compatible with the proposed rate of secular cooling of the Earth’s mantle over the last 3.5 billion years (Richter, 1988), and concurs with previous estimates that the ambient Archaean upper mantle was 50-100° C hotter than at present (Jarvis & Campbell, 1983; Campbell & Griffiths, 1992). Therefore, the Pilbara basalts were probably not derived from anomalously hot mantle, a proposition supported by the paucity of ultramafic volcanics. CHAPTER 8: DISCUSSION - TECTONIC SETTING 202

2) Depth and fraction of melting. The relationship between MgO and potential temperature is also dependent on the fraction and depth of melting (Fig. 8.1; McKenzie & Bickle, 1988). Given a maximum potential temperature of 1380° C, then the Pilbara basalts were either derived from small degrees of partial melting at ~75 km or larger degrees of melting (up to 20 %) at shallower depths (Fig. 8.1). The thick successions of tholeiitic and magnesian basalt in the Pilbara probably preclude their derivation by small degrees of partial melting. Experiments have also shown that garnet is not stable in the modern mantle along the lherzolite solidus at depths shallower than ~78 km (Hirschmann & Stopler, 1996), although this would have been deeper in a hotter Archaean mantle. Nevertheless, since garnet is an aluminous mineral that preferentially incorporates the HREE, melts derived from a garnet-bearing source should have high

MREE/HREE ratios corresponding with low Al 2O3 contents (Sun & Nesbitt, 1978). Therefore, if the depth estimate of <75 km is correct, the Pilbara basalts should not have these garnet features. Indeed, the Pilgangoora basalts show a positive correlation between Al2O3 and Gd/Yb (MREE/HREE), suggesting no garnet fractionation (Fig. 8.2). Decoupling of the typical covariant behaviour between initial e(Nd) and e(Hf) has also been interpreted to reflect garnet fractionation. The Pilgangoora basalts do indeed show strong decoupling (Fig. 6.7), but this is most likely an inherited feature (Chapter 6.5.4.3). Moreover, the time-integrated evolution trends of the basalts are subparallel to the terrestrial array, indicating that their present Sm/Nd and Lu/Hf ratios were not the result of garnet fractionation (Fig. 6.7). Therefore, the Pilgangoora or North Pole basalts were probably not derived from a garnet-bearing source, supporting the proposition that they formed at <75 km depth. 3) Buoyancy versus loading. The North Pole succession is ~10 km thick and was continuously deposited in a shallow subaqueous to subaerial environment (Buick & Dunlop, 1990; Buick et al., 1995). In the Pilgangoora Belt, the Warrawoona Group similarly records persistent shallow-water deposition but the succession is much thinner, whereas the Coonterunah succession is relatively thick but was deposited at greater water depth. Nevertheless, the persistent deposition of a thick volcanic succession near sea-level, as shown in the North Pole Dome, must be accounted for in any tectonic model. The lithosphere is generally considered to be resting on a fluid asthenosphere of uniform density, and hence, the elevation of the Earth’s surface is a measure of the CHAPTER 8: DISCUSSION - TECTONIC SETTING 203 lithospheric buoyancy. Since the buoyancy depends only on the density and thickness of the lithosphere, then a change to either will result in a corresponding change to the surface elevation. Therefore, all parts of the Earth that have the same elevation must have the same lithospheric buoyancy, and vice versa. As a result, addition of magma to the crust should produce thicker, more buoyant lithosphere and a greater surface elevation. In the North Pole Dome, large volumes of basalts were added to the crust and deposition was persistently at sea-level. This implies that the pre-existing lithosphere must have been modified, probably by thinning, to accommodate such basaltic additions while maintaining a constant surface elevation. These changes can be approximately quantified using a 3-layered model (crust, lithospheric mantle, asthenosphere) and assuming that: i) the ~1350° C geotherm defined the boundary between the lithosphere and asthenosphere in the Archaean (as for the modern Earth), ii) the asthenospheric density was 3.30 g cm-3, iii) the lithospheric mantle density can be calculated by its mean temperature, and so is a function of the 1350° C geotherm and the temperature at the base of the crust (Parsons & Sclater, 1977); here assumed to be equivalent to modern stable crust (base of crust = 460° C for all thicknesses), iv) the original crustal density was 2.82 g cm -3 (average crustal density of the modern Pilbara is ~2.81 to 2.84 g cm -3 (Drummond, 1988); this is reasonable because the pre-Warrawoona basement was composed of both Coonterunah greenstones and Carlindi granitoids), v) basaltic additions had a density of 3.02 g cm -3 (as calculated for a 10 km thick basaltic pile derived from a mantle with T P = ~1380° C; White & McKenzie, 1989), vi) the hydrogeoid (sea-level relative to the asthenosphere geoid) was equal to today (-3.5 km; Turcotte et al., 1977). This equates to a ‘dry’ asthenosphere geoid of -2.4 km and implies that the volume of water in the Archaean was equivalent to the present, and vii) the Pilbara basalts were derived from the asthenospheric mantle (their MORB-like characteristics support this claim). The relationship between surface elevation and lithospheric density and thickness are shown in Figure 8.2, where the lithosphere has been divided into its crustal and mantle components (Lachenbruch & Morgan, 1990). Lithospheric mantle is CHAPTER 8: DISCUSSION - TECTONIC SETTING 204 compositionally similar to the asthenosphere, but is cooler and denser (3.35 g cm -3 is used here) and thus has a negative buoyancy. For example, the present Pilbara crust has been estimated from seismic profiling to be ~30 km thick (Drummond, 1988), and so should be ~1850 m above sea-level (Fig. 8.2A). However, the average surface elevation is ~250 m (estimate from topographic maps of northern part of craton) and so the lithospheric mantle must be pulling the crust down by ~1600 m. This requires that the lithospheric mantle is ~100 km thick (Fig. 8.2B).

6 1 2.8 2.9 3.0 5 A 2.75 2.82 B density = 2.65 )

3.02 ) m 4 m k k ( (

l l 0 density = 3.30 e e v

3 v e e l l - - 3.32 = 1000° C a 2 a e e s

s 3.33 = 815° C

e e 3.34 = 640° C v 1 v -1 o o 3.35 = 460° C b b a a

0 sea-level 3.36 = 280° C n n

o 3.37 = 100° C o i i t -1 t a a v v e e

l -2 l e -2 e depth of unloaded asthenospheric mantle (-2.4 km) -3 0 10 20 30 40 50 60 70 80 90 100 0 10 20 30 40 50 60 70 80 90 100 crustal thickness (km) lithospheric mantle thickness (km)

Figure 8.2: Contributions to elevation of continental surface from A) crust, where 2.82 g/cm3 is average Pilbara crust and 3.02 g/cm 3 is density of North Pole basalts, and B) lithospheric mantle, where 3.35 g/cm 3 is the density calculated from a temperature at the base of the crust of 460° C. Other densities refer to different temperatures at the base of the crust, that is, the top of the lithospheric mantle (adapted from Lachenbruch & Morgan, 1990).

These relationships can be extended to the time when the North Pole succession was deposited. If the pre-Warrawoona crust had a density of 2.82 g cm -3 and there was no lithospheric mantle, then crust thinner than ~17 km would have been submerged. This is the minimum initial crustal thickness needed to maintain sea-level, as lithospheric mantle is negatively buoyant and so requires more crust to counterbalance its presence. As the Warrawoona basalts added 10 km of crust at 3.02 g cm -3 and the rest of the crust was still at 2.82 g cm -3, then the minimum crustal thickness to maintain sea-level was ~22 km (10 km of basalts plus 12 km of initial crust; combined density of ~2.91 g cm-3). Consequently, if there was no lithospheric mantle, then ~5 km of the earlier crust must have been removed during Warrawoona volcanism (~29 % of the original crust) to maintain the same buoyancy. Recalculating with a 20 km thick lithospheric mantle before and after eruption requires a pre-Warrawoona crust that was ~19 km thick and a CHAPTER 8: DISCUSSION - TECTONIC SETTING 205 post-eruption crust that was ~23 km thick (10 km of basalts plus 13 km of initial crust), and hence, ~6 km (~32 %) of the earlier crust must have been lost to maintain sea-level. With a thicker lithospheric mantle, therefore, marginally more of the earlier crust needs to be removed to maintain sea-level after addition of 10 km of basalts. A similar relationship is found irrespective of the density of the lithospheric mantle. Furthermore, the heat associated with Warrawoona volcanism probably decreased the depth of the 1350° C geotherm, and so it is unlikely that there was a thick lithospheric mantle during this time. If the 20 km thick lithospheric mantle was totally removed during eruption, then the pre-Warrawoona crust was ~19 km thick, the post-eruption crust was 22 km thick (10 km of basalts plus 12 km of initial crust) and ~7 km (~32 %) of the earlier crust must have been lost. In summary, it seems that at least 5 -7 km or ~30 % of the pre-Warrawoona crust must have been removed to allow 10 km of basalt to be deposited at constant surface elevation. The removal or thinning of pre- existing basement is clearly an important characteristic for the tectonic model. The Warrawoona basalts have compositions that reflect some crustal assimilation, but this can only account for <1 km of basement removal (<10 % maximum contamination for ~10 km of basalts). In the Pilgangoora Belt, there is evidence that at least 2 km of the Coonterunah succession was eroded to form the unconformity. Such erosion may account for a thinner Warrawoona succession above the unconformity, but cannot explain the loss of basement in the North Pole Dome. Indeed, there is little evidence of terrigenous clastic deposition during Warrawoona volcanism. The most likely cause of crustal thinning was lithospheric stretching, and, using the above values, a stretching factor of ~1.4 to 1.5 is estimated. The relationship between lithospheric stretching and crustal addition required to maintain a constant surface elevation is shown graphically for a setting where the lithospheric mantle was 40 km thick and stretching was uniformly distributed between the crust and mantle (Fig. 8.3A). The case where there was no lithospheric mantle is also shown for sea-level (dashed line), and indicates that changes to the mantle contribution do not significantly increase the required stretching. CHAPTER 8: DISCUSSION - TECTONIC SETTING 206

20 3 km A 2 km ) B

m 1 km k (

) 1480° C 30 15 d m e k (

d 0 km

d s a s

e s 20 10 n s k e c i n k h 1380° C t c

i t 10 l h 5 t e

t l e = -2.5 km m a s

a 1280° C

b 0 0 1 2 3 4 1 2 3 4 5 stretching factor (b) stretching factor (b)

Figure 8.3: A) Relationship between thickness of basalt added and the stretching factor ( b) required to maintain original surface elevation ( e) with uniform isothermal stretching and a 40 km thick lithospheric mantle. Dashed line is 0 km (sea-level) for a crust with no lithospheric mantle. Values used are discussed in text. Adapted from Lachenbruch & Morgan (1990). B) Relationship between stretching factor and thickness of basaltic melt derived by adiabatic decompression associated with lithospheric stretching. Three potential temperature ranges are shown where the upper and lower curves for each potential temperature represent the mechanical boundary layer at 100 and 70 km, respectively (McKenzie & Bickle, 1988).

8.4 POSSIBLE TECTONIC SETTINGS Most tectonic models proposed for the formation of Archaean greenstone belts have been actualistic, using modern plate tectonic settings but with slight variations due to differing Archaean conditions such as higher heat flow. These models include accreted oceanic crust (ophiolites) and plateaux, volcanic island arcs, marginal basins (back-arc and pull-apart basins), continental rifts and crustal overthickening by flood basalt volcanism. However, non-actualistic processes and environments, those with no known modern analogues, may also have been responsible for forming Archaean greenstone belts. For example, meteorite impacts may have formed basalt-filled basins, similar to lunar mare, which were then accreted to continental crust to form greenstone belts. However, the geological and geochemical details of non-actualistic models are poorly defined, and so devising tests to discriminate them from other greenstone models is difficult. Below, a broad range of possible actualistic models and one of the better- CHAPTER 8: DISCUSSION - TECTONIC SETTING 207 constrained non-actualistic models are tested against the geological, geochemical and stratigraphic characteristics of the Pilbara greenstones listed above.

8.4.1 Meteorite impact basins The collision of large meteorites with Earth may have formed basalt-filled basins, similar to lunar mare, which were then preserved as greenstone belts (Green, 1972; Glikson, 1996, 1999). Lunar cratering rates and sizes indicate that the flux of large meteorites was greater during the Moon’s early history, and so there should have been a corresponding higher flux of very large terrestrial impacts (Ryder et al., 1991; Culler et al., 2000). Therefore, the formation of impact basins should have been much more common in the Archaean than it is now. Under such conditions, the impact of very large meteorites causes mantle melting due to rebound-induced adiabatic melting and shock heating (Glikson, 1996). The resulting melts would erupt into the impact basin. This model, therefore, would require that the Coonterunah and Warrawoona successions were the result of two very large impacts. It follows that there should be evidence of the cataclysmic Warrawoona event in the earlier Coonterunah-Carlindi rocks, but despite searching, no such evidence has been discovered. Possible impact spherules have been reported from the North Pole Chert (Lowe & Byerly, 1987), but their origin has been disputed (Buick, 1987; French, 1987), and even if they are spherules, they must be related to a later event because they are >1 km above the base of the succession. Although it is unclear what style of evolving thermal anomaly and associated basalt stratigraphy would be induced by a very large impact, it can be assumed, as long as tectonic rifting did not commence, that the system would cool and attain isostatic stability. As the thermal anomaly dissipates, there should be a corresponding decrease of basalt flow thickness and also a systematic variation of basalt composition due to progressively smaller degrees of partial melting. The stratigraphic trend of basalt composition in the lower Coonterunah succession is similar to that proposed, but this trend is not present for the entire Coonterunah succession and is reversed for the North Pole basalts. Thus, there is little evidence that the Pilbara greenstone belts represent ancient basalt-filled impact basins. CHAPTER 8: DISCUSSION - TECTONIC SETTING 208

8.4.2 Ocean crust Accretion of oceanic crust (de Wit et al., 1987; Kusky & Kidd, 1992) has been proposed as an origin for greenstone belts, although Bickle et al. (1994) present compelling evidence that ophiolites, and thus irrefutable evidence of oceanic crust, have not yet been described in any Archaean greenstone belt. Certainly the early Pilbara greenstones cannot represent ancient oceanic crust because they were deposited onto continental basement, have appreciable bimodal volcanism and were intruded by coeval granitic plutons. In addition, some of the defining features of oceanic crust, such as abundant cross-cutting dykes related to lateral accretion and sheeted dyke complexes, are absent in the Pilbara successions, even though they are up to 10 km thick. Moreover, the lack of convergence-related structures in the Pilgangoora and North Pole domains provides no evidence for a preservation mechanism, such as accretion or obduction.

8.4.3 Oceanic plateaux Oceanic plateaux are thick subaqueous basalt-dominated volcanic piles which may be accreted to continental crust and preserved in the rock record (Abouchami et al., 1990; Storey et al., 1991; Desrochers et al., 1993; Puchtel et al., 1998). The Pilbara greenstones have many features in common with modern oceanic plateaux, particularly their great thickness of subaqueous pillowed basalts without sheeted dykes (Kerr et al., 2000). However, oceanic plateaux are typically floored by oceanic crust, not enriched in incompatible elements and rarely contain felsic volcaniclastic deposits (Kerr et al., 2000). As for oceanic crust, the lack of convergence-related structures in the Pilbara provides no evidence for a preservation mechanism. A special case, however, may be where plateaux are developed on attenuated continental crust, as proposed for parts of the Kerguelen Plateau (Storey et al., 1989; Frey et al., 2000), and such a setting is considered with continental flood basalts (Chapter 8.4.7).

8.4.4 Volcanic arcs Volcanic arcs built on both oceanic or continental substrates are commonly proposed as settings for greenstone belts because they have bimodal volcanism, a broad range of submarine to subaerial environments and voluminous coeval plutonism (e.g., Slave Province, King & Helmstaedt, 1997; Karelian Province, Sorjonen-Ward et al., 1997; Pilbara Craton, Barley et al., 1984; Barley, 1993). Moreover, the plutonic rocks CHAPTER 8: DISCUSSION - TECTONIC SETTING 209 associated with granite-greenstone terrains commonly have a tonalite-trondhjemite- granodiorite composition, which has been explained by partial melting of subducted altered oceanic crust (see Chapter 5). However, volcanic arc basalts are dominantly enriched in crustal material which has been derived from subduction-related mantle metasomatism, but, in contrast, the Pilbara basalts are best interpreted as having assimilated crustal material (Chapter 7). There are also well-established temporal and spatial variations of basalt composition in volcanic arcs (Baker, 1982), which are inconsistent with the contrasting stratigraphic trends in the Coonterunah and North Pole successions. Moreover, temporal trends are not reflected between the Coonterunah and Warrawoona basalts, although they erupted ~ 55-80 myr apart. Volcanic arcs may also develop paired metamorphic belts, with a high-pressure, low-temperature blueschist- to eclogite-facies belt adjacent to a low-pressure, high-temperature metamorphic belt (Miyashiro, 1973). However, in the North Pole Dome and Pilgangoora Belt, metamorphism was at low temperatures and pressures, and no blueschist-facies rocks have been discovered in the Pilbara. Finally, since volcanic arcs form at convergent margins they are often intensely deformed soon after deposition. In contrast, the North Pole and Pilgangoora domains have been only weakly deformed. Therefore, although a volcanic-arc setting has been previously proposed for the Warrawoona Group (Barley et al., 1984; Barley, 1993), there is little supporting evidence for this interpretation, particularly in the Pilgangoora Belt and North Pole Dome.

8.4.5 Back-arc basins Back-arc basins are a type of marginal basin that have been proposed as modern analogues for greenstone belts because they contain mostly basaltic rocks with a subordinate felsic component, the basalts have a crustal signature (subduction-related), and their association with convergent margins provides an obvious mechanism, obduction, for addition to and preservation on pre-existing continental basement (Saunders & Tarney, 1991). Back-arc basins typically have oceanic basements and deep seafloors (Saunders & Tarney, 1991), which would preclude them as Pilbara analogues. However, some back-arc basins, such as in the Bransfield Strait (Keller et al., 1991; Lawver et al., 1995) and the Sea of Japan (Tamaki, 1995), have basements of attenuated continental crust. Bathymetric surveys of these basins, however, show that shallow-water domains are rare and typically small, and the basin topography is CHAPTER 8: DISCUSSION - TECTONIC SETTING 210 extremely rugged (Klepeis & Lawver, 1994; Tamaki, 1995). Furthermore, in basins with continental basement, true MORB is erupted only when the basin becomes wide and deep (Saunders & Tarney, 1991). Thus, continued widespread eruption of MORB-like basalts in shallow subaqueous to subaerial environments, as seen in the Pilbara, is not evident in modern back-arc basins. Moreover, the crustal component of back-arc basin basalts, including those with continental basements, is dominantly derived from subduction-related mantle enrichment, and there is little evidence of basement assimilation (e.g., Bransfield Strait; Keller et al., 1991). In contrast, the crustal composition of the Pilbara basalts is apparently due solely to crustal assimilation, with no geochemical evidence of subduction. Since back-arc basins are extensional basins, they commonly form sheeted dykes (Kerr et al., 2000), a feature not present in the Pilbara greenstones. As in volcanic arc settings, the mantle source in modern back-arc environments evolves chemically such that there are temporal trends of basalt composition (Ewart et al., 1998). Such trends are not evident as chemical variations between the Coonterunah and Warrawoona basalts, although they erupted ~55-80 myr apart, and stratigraphic trends within individual Pilbara successions contrast markedly. Therefore, the evidence for a back-arc basin setting, even one with a continental basement, is not strong for the Pilbara greenstones.

8.4.6 Rifts To maintain constant sea-level during deposition of the North Pole basalts, the initial crust must have been thinned, most likely by stretching, and thus the basalts were probably deposited in an extensional tectonic setting. Besides those already discussed, there are other modern extensional settings that form basalt-filled basins, such as pull- apart basins (e.g. Gulf of California, Saunders & Tarney, 1991), continental rifts (e.g. Oslo rift, Neumann et al., 1995) and highly extended terranes (e.g. Basin and Range Province, Axen et al., 1993). Since these settings are not necessarily associated with subduction zones, then the crustal signatures of basalts erupted in them are likely to result from assimilation. Moreover, bimodal volcanism and coeval granitoid emplacement are not uncommon in these settings. For instance, the extrusive and intrusive volcanic units in the Oslo rift have a wide range of compositions from granitic to basaltic (Neumann et al., 1995). CHAPTER 8: DISCUSSION - TECTONIC SETTING 211

Typical rifts, however, have characteristics that differ gr eatly from the Pilbara greenstones. Volcanic rocks in continental rifts are generally alkaline with a decrease in alkalinity with time (Bailey, 1983), a trend perhaps evident in the North Pole succession, but not in the Coonterunah succession. Modern rifts also subside more than expected for a given crustal thickness, and thus deviate from Airy isostatic equilibrium (Bott, 1981). This is due to the rifted fault block being wedge-shaped, and thus differing from the assumed parallel-sided column used for Airy isostatic calculations (Heiskanen & Vening Meinesz, 1958), and also because of mechanical differences between the brittle upper crust and the ductile lower crust during rifting (Bott, 1981). Subsidence is then compounded by increasing water depth and clastic sediment loading. Therefore, maintaining sea-level, even with mantle additions, is not typical of rift systems. Moreover, clastic sediments derived from the basin flanks are common in continental rifts (Olsen, 1995; and papers therein), but are generally lacking in the Coonterunah and Warrawoona successions. Therefore, typical modern rift settings are unlikely analogues to the Pilbara greenstones. A special case, however, has been noted in the North Atlantic where early Tertiary rifting was coincident with the deposition of thick basalt successions (White et al., 1987; White & McKenzie, 1989). This relationship was caused by rifting above anomalously hot mantle (1400 to 1450° C), which increased the volume of mantle melting, and hence increased the crustal thickness and buoyancy (White et al., 1987). If the potential temperature of the Archaean upper mantle was ~100° C hotter than typical modern upper mantle (MORB source), then perhaps the Pilbara successions simply reflect passive continental rifting above hotter mantle. However, at a potential temperature of ~1380° C, stretching factors of >5 are required to produce ~10 km of basalts (MgO = 12 wt%) by adiabatic decompression of the mantle (Fig. 8.3B; McKenzie & Bickle 1988), which is greater than the estimated stretching required to have maintained North Pole deposition at constant sea-level. Moreover, large stretching factors should lead to the development of an oceanic regime, as demonstrated in the North Atlantic. Furthermore, for a rift setting to b e plausible, it must provide a mechanism for superposition of two distinct successions 55-80 myr apart at different base levels (Coonterunah in deep water followed by Warrawoona at sea-level). Clearly, there must have been a discrete and significant (>1 km) uplift event prior to deposition of the CHAPTER 8: DISCUSSION - TECTONIC SETTING 212

Warrawoona Group to explain the emergence of the Coonterunah-Carlindi terrain. Pronounced pre-Warrawoona deformation fabrics were not developed in either the Coonterunah succession or the Carlindi granitoids, and so the intervening event was not obviously convergent. Hence, uplift was probably caused by some form of vertical tectonics, and indeed, several previous studies in the Pilbara have concluded that emplacement of granitoid batholiths was by solid-state diapirism, a form of vertical tectonics (Collins, 1989, 1993; Collins & Van Kranendonk, 1999). However, since the granitoids have been interpreted as partial melts of pre-existing crust (Chapter 5, 7), then such diapirism would have only redistributed material within the crust. In contrast, isostatic compensation of the entire crust must be accommodated in the asthenosphere, and so such a significant uplift event was probably induced from the mantle. One possible solution is that the Warrawoona event occurred above actively upwelling mantle, for instance, above a mantle plume, while the preceding Coonterunah event was just caused by passive rifting above hot mantle. The proximity of the Iceland plume has been proposed to explain the anomalous heat source required for the North Atlantic basalt successions (White et al., 1987), so the superposition of a rift and plume has a possible modern exemplar, albeit in reverse order. However, the compositions of the Coonterunah and Warrawoona basalts are so similar that they indicate a similar petrogenesis; that is, the elemental and mineralogical composition of the mantle source, depth of melting, potential mantle temperature and fractional crystallisation processes were essentially the same. It is thus difficult to conclude that the Coonterunah basalts formed in a passive rift and the Warrawoona basalts formed in a rift above actively upwelling mantle. It is more likely that they both formed above actively upwelling mantle, perhaps similar to modern flood basalt provinces, a possibility considered in the following section.

8.4.7 Flood basalts Flood basalt provinces contain vast volumes (>2 x 10 6 km3, up to 5 km thick,) of mafic volcanics that were erupted subaerially onto continental crust, and so could be analogous to the early Pilbara greenstone successions (Arndt, 1999). They are interpreted to form as a result of anomalous thermal conditions in the mantle due to the upwelling of hot material, possibly plumes, and have formed periodically since, at least, the late Archaean (Nelson et al., 1992; Blake, 1993). Modern examples are dominated CHAPTER 8: DISCUSSION - TECTONIC SETTING 213 by basalt, although bimodal associations are not uncommon (Arndt, 1999). Flood basalts typically have a crustal signature in their composition which has been attributed to either melting of lithospheric mantle (Gallagher & Hawkesworth, 1992) or crustal assimilation during eruption (Larsen & Watt, 1985; Nelson et al., 1992; Arndt et al., 1993). Modern events are associated with up to 1000 m of uplift and varying degrees of lithospheric extension. Continental breakup has postdated the development of some continental flood basalt provinces (e.g., North Atlantic, Paraná-Etendeka, Afar, Karoo), whereas other major events, such as the Siberian traps, have not lead to the formation of oceanic basins (Hill et al., 1992). Anatectic melting of continental basement, and thus felsic magmatism, may be delayed by up to tens of millions of years due to a thermal lag in heating the crust to melting temperatures (Campbell & Griffiths, 1992). So, the Coonterunah and Warrawoona basalts could reasonably be interpreted as two stacked flood basalt successions. In such a model, the first flood basalts were erupted onto submerged continental basement to form the ~3517 Ma Coonterunah Group and delayed anatectic melting of the basement formed the ~3484-3468 Ma Carlindi granitoids. The Warrawoona event uplifted and weakly deformed the Coonterunah-Carlindi terrain to produce a topographic high, which was, in turn, eroded to form the Pilgangoora angular unconformity. Crustal extension was probably induced by the uplift and continued such that the pre-North Pole crust was stretched by a factor of ~1.5. The crust was thinned by stretching and subsequent basalt deposition was maintained at sea-level. While the Pilgangoora domain was emergent, basalts were being deposited in the North Pole domain, which must have been topographically lower. The two domains achieved the same topographic level by the time the Strelley Pool Chert was deposited. The broad regional difference in base-level between the two eruption events may be due to deposition of the Coonterunah basalts on thinner continental basement or distal from the mantle upwelling. The former would imply that in the Pilgangoora domain, the Coonterunah event added to the total crustal thickness, in contrast to the Warrawoona event which involved thinning of the original crust. Since the upper part of the Coonterunah succession has been truncated by an erosional unconformity, determining better isostatic constraints is not now possible. For instance, the missing upper part of the Coonterunah succession may have been emergent. A significant aspect of the continental flood basalt model is that there is no requirement for CHAPTER 8: DISCUSSION - TECTONIC SETTING 214 horizontal tectonic activity, such as obduction, for epicratonic preservation, which is consistent with the limited deformation of the Pilgangoora and North Pole successions. Hence, the eruption of stacked flood basalt successions can accommodate the observed and derived characteristics of the Pilbara successions, except that Phanerozoic flood basalts are typically <5 km thick and exclusively subaerial, whereas the Pilbara successions are up to ~10 km thick and were deposited in subaqueous environments. If the Archaean mantle was hotter than its Phanerozoic counterpart, as is generally believed, then the Pilbara successions should be correspondingly thicker than Phanerozoic flood basalt successions. However, a ~10 km thick basalt succession with up to ~12 % MgO is not consistent with simple adiabatic melting of mantle with a potential temperature of ~1380° C (Fig. 8.3B; McKenzie & Bickle, 1988). Indeed, a similar problem applies to Phanerozoic flood basalt successions, but at lower potential temperatures. To resolve this, it has been proposed that upwelling physically moves hot mantle to shallower depths where large degrees of partial melting are induced (Campbell et al., 1989; White & McKenzie, 1989; Campbell & Griffith, 1992). Moreover, with melting limited to the top 78 km of the mantle (absence of garnet signature in Pilbara basalts) and beneath a ~20 km thick crust, then to form a ~10 km thick basalt pile requires large melt fractions (~20 %) or steady replenishment of the source region during eruption. Such conditions are consistent with models explaining eruption of mantle plume heads (Campbell et al., 1989). Stratigraphic trends of basalt composition in the North Pole basalts supports steady replenishment, whereas the inverse trend of the Coonterunah basalts suggests isolation in shallow magma chambers after magma separation from the mantle. Many models have been proposed to explain the frequent occurrence of submerged basalt volcanism on continental basement in the Archaean, including greater crustal densities, larger water volumes and shallower ocean basins (see Arndt, 1999 for review). Generally, these models imply that there was little emergent continental crust in the Archaean because relative sea-level was much higher. However, the Coonterunah- Carlindi terrain was clearly emergent and provides evidence that the ocean did not cover all of the continents at ~3.5 Ga. Hence, continental surface elevation was more likely a local effect, with apparent widespread Archaean submergence perhaps caused by crustal stretching and thinning concomitant with deposition. Indeed, such processes most readily explain the eruption of flood basalt successions onto submerged continental crust CHAPTER 8: DISCUSSION - TECTONIC SETTING 215 in the Pilbara. Moreover, these submerged terrains are likely to be preserved because, unlike oceanic crust, they are not readily subducted and, unlike subaerial deposits, they are not readily eroded. The Coonterunah and Warrawoona Groups, therefore, could plausibly represent two stacked flood basalt successions that were erupted onto continental basement. Uplift predating the Warrawoona eruptions in the Pilgangoora domain suggests that mantle upwelling may have been an important process inducing mantle melting. Moreover, the impingement of hot mantle at the base of the crust may have induced melting of the continental basement and thus formed the Pilbara granitoids.

8.5 DISCUSSION The discovery that the Pilbara greenstone successions are most reasonably explained as submerged continental flood basalts has important implications for Archaean mantle dynamics and crustal growth. The former is discussed here and the latter is discussed in Chapter 9.

8.5.1 Mantle upwelling The flood basalt model is the most viable tectonic model fo r early Pilbara greenstone evolution and implies active mantle upwelling. Moreover, the tectonic similarity of the Pilbara greenstones with continental flood basalt provinces may imply that they were formed by similar mantle perturbations, that is, a plume originating from the core-mantle boundary (Campbell et al., 1989; Hill et al., 1992). However, this is not necessarily the case. Although upwellings above subduction zones have been largely discounted by the Pilbara geochemical data, upwellings that originated elsewhere in the mantle are not precluded. Indeed, some models of Earth’s evolution have a strongly layered mantle in the Archaean (McKenzie & Richter, 1978; Allègre, 1997) and so upwellings may have originated from an intramantle boundary layer. Since upwellings derived from greater depths will have corresponding larger ‘plume heads’, the most compelling evidence for the depth of origin of an upwelling is the aerial extent of its magmatism (Hill et al., 1992). Unfortunately, the North Pole and Pilgangoora greenstones represent only part of the original volcanism, the rest having been covered or eroded, so the full extent of the Pilbara basalts cannot be determined. Geochemical characteristics also provide little constraint upon the mode of upwelling. For example, a geochemical compilation of possible plume-related picrites and CHAPTER 8: DISCUSSION - TECTONIC SETTING 216 komatiites shows only that modern eruptions are remarkably heterogeneous and that possible Archaean plume-related basalts were different (Campbell & Griffiths, 1992). However, conclusive evidence that these Archaean basalts were indeed plume-related has not been established. It has also been argued that komatiites are an indication of a hot, deep plume source (Campbell & Griffiths, 1992; Storey et al., 1991), but contradictory geochemical and mineralogical modelling suggests that some komatiites may have been derived from a volatile-rich mantle and so were subduction-related (Parman et al., 1997). However, the evidence that komatiites were derived by wet melting is hotly disputed (Arndt et al., 1999), and so komatiites remain uncertain indicators of upwelling modes. Regardless, komatiites are rare in the Pilbara successions. Hence, there are no criteria that discriminate between different modes and sources of the mantle upwelling responsible for the eruption of the Pilbara basalts.

8.6 SUMMARY The geochemical, geological and geophysical characteristics of the Pilbara greenstone successions preclude their formation in many modern and ancient basalt- dominated tectonic setting. The Pilbara characteristics can be best explained as flood basalt successions deposited onto thin, submerged continental basement. Hence, the Pilgangoora Belt evidently comprises two flood basalt successions separated by an erosional unconformity. Heating associated with the basalts apparently melted pre- existing crust and formed the Carlindi granitoids. CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 217

Chapter 9: DISCUSSION – MANTLE AND CRUSTAL EVOLUTION

9.1 INTRODUCTION The silicate portion of the modern Earth is not homogeneous, and so it does not represent the primitive mantle. One distinct component is continental crust, a relatively thin layer that is highly enriched in many incompatible elements relative to the primitive mantle. In this chapter, the formation and evolution of continental crust are discussed, with particular focus on its abundance and origin in the early Archaean. The discussion starts with an overview of crustal genesis and proposed growth models. The relevant Pilgangoora data are then considered, in particular Nd-Hf isotopes and Nb-U systematics. From this, there seems to be sound evidence that the modern mass of continental crust had formed by 3.5 Ga, and perhaps during the first few hundred millions years of Earth history. If so, the present age distribution of the continental crust is an artifact of preservation and recycling, and is not representative of crustal growth.

9.2 CONTINENTAL CRUST Continental crust covers ~40 % of the Earth’s surface, has an average thickness of ~40 km and comprises ~0.6 % of the mantle’s mass (Cogley, 1984; Christensen & Mooney, 1995). It is composed of a diverse range of rock types but has a mean andesitic composition (Taylor & McLennan, 1985, 1995; Rudnick & Fountain, 1995). The oldest known continental rocks are ~4.03 byr old (Bowring & Williams, 1999), although much older zircons with continental affinities have been found (Froude et al., 1983; Compston & Pidgeon, 1986; Maas et al., 1992; Mojzsis et al., 2001; Wilde et al., 2001). In contrast, oceanic crust is dominantly basaltic, thinner (<7 km thick) and younger (<200 myr old). As a result, continental crust is more buoyant with a mean elevation of ~5 km above the sea- floor or ~125 m above sea-level (Cogley, 1984; Taylor & McLennan, 1985, 1995). Although r emarkably small, continental crust forms a major global reservoir for many incompatible elements, and so there should be a complementary depleted reservoir within the mantle. Indeed, the commonest mantle-derived rocks, mid-oceanic ridge basalts (MORB), have compositions that indicate melting of such a depleted source. This relationship is shown in Figure 9.1. Compared to primitive mantle, continental crust is more CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 218 enriched in the more incompatible elements with negative Nb-Ti and positive U-Pb anomalies. MORB are also more enriched than the primitive mantle, but are less enriched in the most incompatible elements than in the moderately incompatible elements and, notably, have positive Nb and negative U-Pb anomalies, an inverse trend to continental crust (except for Ti). Hence, MORB must have been derived from a source previously depleted in incompatible elements and possessing these Nb-U-Pb anomalies. This is also consistent with the depleted compositions of peridotite massifs and xenoliths, which are direct samples of the mantle (McDonough & Sun, 1995). An example of a peridotite xenolith is shown in Figure 9.1 (fertile lherzolite 2905; Eggins et al., 1998), though this sample is relatively enriched in Pb and has been replotted using Ce/Pb = 25, the typical mantle value (Hofmann et al., 1986; Sims & DePaolo, 1997). Other mantle-derived rocks, such as ocean island basalts, have features which indicate that they are mixtures of depleted mantle and enriched components, thus complicating any simple mass balance relationship between continental crust and depleted mantle (Zindler & Hart, 1986; Rudnick, 1995; Hofmann, 1997). Complications also arise due to mass balance deficiencies of Nb, Ta and Ti between the crust and mantle (McDonough, 1991; Barth et al., 2000; Rudnick et al., 2000), although recent experiments suggest that some Nb may have been sequestered into the core (Wade & Wood, 2001). Nevertheless, continental crust and depleted mantle can be considered to be generally complementary (Hofmann, 1988, 1997).

9.2.1 Crustal genesis Even with such a simple complementary relationship, there are many complications constraining crustal growth processes, particularly the need to explain its bulk composition with respect to the depleted mantle and its compositional and temporal diversity. For instance, mantle melts are overwhelmingly basaltic, and thus it is unlikely that direct mantle melting could have formed a crust with an andesitic bulk composition. This would have been even less likely during the Archaean when the mantle was hotter and melt fractions were greater (Chapter 8.2.2; McKenzie & Bickle, 1988). In part, however, the continents are composed of basaltic rocks, including flood basalts, ophiolites, oceanic plateaux and greenstone belts. Moreover, it has been estimated that basaltic terrains associated with mantle plumes may have contributed 5-20 % of the continental La/Nb budget (Rudnick, CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 219

1995; Barth et al., 2000), which is probably also a fair estimate of the total basalt component in the continents. Thus, direct mantle melting has probably been a minor, albeit important, contributor to the total continental mass but could not have produced an andesitic bulk crust. As a result, profoundly different processes must have produced the distinctive continental composition.

Continental crust 100 n o i t a r t n

e 10 c n

o MORB c

d

e 1 Primitive mantle s i l a

m Depleted mantle r o 0.1 n Rb Th U Ce Nd P Hf Ti Dy Er Al Fe Mg Ba Nb La Pb Sr Sm Eu Gd Y Lu Ca Si

Figure 9.1: Primitive mantle normalised concentrations of average continental crust and MORB (from Hofmann, 1997). Fertile lherzolite xenolith is proxy for depleted mantle (Eggins et al., 1998; Pb recalculated to Ce/Pb = 25; dashed line is measured value). Element compatibility increases to the right.

At first glance, the prevalence of andesites in some modern subduction settings suggests a relationship with continental growth (Taylor, 1967, 1977). However, andesites are most common above subduction zones associated with pre-existing continental crust and are relatively rare in juvenile volcanic arc systems (Rudnick, 1995). Furthermore, melting in modern subduction zones predominantly occurs in the mantle wedge and thus the net flux from the mantle to the crust is essentially basaltic with more evolved rocks typically forming by intracrustal fractionation (Anderson, 1982; Grove & Kinzler, 1986; Brenan et al., 1994; Pearce et al., 1995). An andesite-accretion growth model also struggles to account for significant chemical differences between modern andesites and bulk crustal estimates of Cr and Ni contents and Th/U (Taylor & McLennan, 1985, 1995) and Rb/Sr ratios (Ellam & Hawkesworth, 1988). With this in mind, Taylor & McLennan (1985, 1995) proposed that crustal growth processes may have profoundly changed through time. CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 220

Other models explaining the andesitic bulk crust include partial melting of subducted oceanic crust to form tonalite-trondhjemite-granodiorite (TTG) magmas with a garnet amphibolite residue (see Chapter 5.3.2.1; Arth & Barker, 1976; Martin, 1987, 1993; Drummond & Defant, 1990; Rapp et al., 1991; Winther, 1996; Springer & Seck, 1997), intracrustal differentiation of basaltic crust followed by delamination of dense ultramafic to eclogitic lower crust leaving an enriched complementary residue (Anderson, 1982; Arndt & Goldstein, 1989; Kay & Kay, 1991; Davies, 1992) and preferential recycling of Mg back into the mantle (Anderson, 1982; Albarède & Michard, 1986; Albarède, 1998). Of these, the TTG model is the most satisfactory because it can explain the previously mentioned geochemical features. Moreover, TTG form a significant proportion of the continents (Condie, 1993; Rudnick, 1995), perhaps up to 75 % (Taylor & McLennan, 1985, 1995). Nevertheless, any of the proposed models could have contributed to the growth of continental crust.

9.2.2 Crustal growth Given the present diversity of continental ro cks and the likelihood that the crust formed by various processes, then it is no wonder that there has been little agreement about when continental crust attained its present mass. There are two general models; i) gradual growth, perhaps punctuated by periods of more rapid growth (Fig. 9.2A(c-e); Hurley & Rand, 1969; Reymer & Schubert, 1984; Taylor & McLennan, 1985, 1995; McCulloch & Bennett, 1993, 1994; Albarède, 1998; Collerson & Kamber, 1999), and ii) steady-state growth, where the present crustal mass was attained very early and has been maintained through growth concomitant with recycling (Fig. 9.2A(d-e); Armstrong, 1968, 1981, 1991; Fyfe, 1978; Bowring & Housh, 1995; Sylvester et al., 1997). Importantly, crustal growth refers to the complimentary extraction of sialic crust from the mantle, and hence basaltic melts are not readily accounted for in most growth models, though they are important components of the crustal mass. At first glance, it would seem that the modern Earth could be used to distinguis h between the models because gradual growth predicts that continental crust is presently growing, whereas a steady-state predicts no net flux. Although there is widespread agreement that continental material is consumed in subduction zones (McCulloch, 1989; CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 221

White, 1989), estimates of the flux in individual subduction zones (Plank & Langmuir, 1993; Pearce et al., 1995) cannot be easily extrapolated to a global flux. Other poorly constrained processes, such as removal of dissolved continental material during alteration of oceanic crust and tectonic erosion of the lower continental crust, also complicate mass balance estimates. Hence, it is unclear whether the continental mass is currently increasing, remaining constant or even decreasing, and so modern subduction fluxes have not been able to distinguish between the models.

1.2 1.2 B high T / low P modern island A e subduction regime ???? arc regime Intense d 1.0 1.0 Meteorite Bombardment s h s t r a a

0.8 m 0.8

c E l

a e t h s t

u b f a b r o c

0.6 0.6 n n r o i e t e d r o major episodes of c c m

crustal growth 0.4 f 0.4 a A o

n o i t c

a 0.2 0.2 r f

supercontinents in place 0 0 0 1 2 3 4 4 3 2 1 0 Age (Ga) Age (Ga)

Figure 9.2: A) Crustal growth models; a) Hurley & Rand (1968), b) Reymer & Schubert (1984), c) McLennan & Taylor (1985), d) Armstrong (1981), e) Fyfe (1978). Adapted from Bowring & Housh (1995). B) Schematic model of growth and evolution of continental crust from Taylor & McLennan (1995); a) crust dominated by greenstone belt (low-grade) terranes with lesser regions of cratonic (high-grade) terranes, b) massive crustal melting to produce granite and granodiorite.

Consequently, it has been necessary to use the exposed rock record to constrain growth models, although this may include a sampling bias. Fortunately, there is little evidence to suggest any significant age differences between the upper and lower crust (Rudnick, 1992), and so the age of the exposed crust is probably representative of the entire crust. Rocks older than 2.5 Ga account for less than 15 % of the present exposed continental area (Windley, 1995) with the oldest-known terrestrial rocks exposed over less than ~40 km 2 (~4.03 Ga Acasta gneisses; Bowring & Williams, 1999). Zircon fragments predate the Acasta gneisses with a diversity of ages up to ~4.40 Ga and indicate the existence of even older crust (Froude et al., 1983; Compston & Pidgeon, 1986; Mojzsis et CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 222 al., 2001; Wilde et al., 2001). In the gradual growth models, the present age distribution of continental crust broadly corresponds with the crustal volume at a particular time; that is, there is little Archaean crust because not much was formed. In contrast, the steady-state model envisages that the exposed ancient rocks represent the remaining relics of vigorous recycling. There is also clear evidence that large parts of the continental crust were formed very quickly (DePaolo, 1981; Patchett & Bridgwater, 1984; Nelson & DePaolo, 1985; Abouchami et al., 1990), suggesting that continental growth reflects some sort of periodicity within the mantle. Again, these events may represent positive crustal growth or just preferential preservation. A number of these features are evident in the gradual growth model shown in Figure 9.2B (Taylor & McLennan, 1985, 1995), where the small volumes of crust that formed before ~3.8 Ga were vigorously recycled, perhaps by intense meteorite bombardment, followed by rapid growth by addition of TTG magmas during the late Archaean and then slower growth as the modern island-arc regime started to dominate the tectonic cycle. In a steady-state model, these apparent changes of tectonic processes would correspond with changes to the recycling rate; that is, recycling was far more vigorous in the Archaean and so there is little preserved crust from this time. Again, the supporting evidence is equivocal and can be modelled accordingly. So, given the paucity of the Archaean rock record, perhaps due to the reasons cited in favour of either model, constraining the timing of crustal growth directly from the age distribution of the rock record is not feasible. An alternative approach, and one that has found favour recently, has been to use the complementary chemical relationship between continental crust and depleted mantle (Fig. 9.1). In this approach, crustal growth should correspond with the timing of mantle depletion, and hence, gradual growth should have gradually depleted the mantle, whereas in a steady-state system the mantle should always have been as depleted as the modern mantle. Since the greatest discrepancy between the models is in the early Archaean, the most ancient rocks are probably the most informative. Monitoring mantle depletion has been achieved with Nd isotopes, and more recently Hf isotopes, both of which seem to indicate that the oldest terrestrial rocks, including ancient zircons, were derived from a depleted source (see section 9.4 below). However, these results have also been conversely modelled, CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 223 or dismissed as unreliable, in favour of a gradual-growth model. With all these issues in mind, the Pilbara data are marshalled to further constrain crustal growth processes.

9.3 IMPORTANCE OF THE PILBARA To understand mantle differentiation and crustal evolution, it ha s been necessary to study very old rocks and thus there have been many detailed studies of early Archaean terrains. In general, however, pre-3.5 Ga terrains are intensely deformed and highly metamorphosed (Windley, 1995), and so deciphering primary geological and geochemical features has been problematic. For instance, there has been much debate about the validity of extremely depleted initial e(Nd) compositions from the ~3.8 Ga Isua Greenstone Belt (see 9.4.1.1). Nevertheless, these results cannot simply be discarded a priori because it is unclear whether early Archaean processes were profoundly different from those that followed. As a result, it is important to compare these oldest terrains with those that are slightly younger but better preserved and less equivocal. The Coonterunah and Warrawoona successions are among the oldest known well-preserved supracrustal rocks on Earth and are roughly contemporary with well-preserved basalt-komatiite successions in the Kaapvaal Craton (~3.48 Ga Komati Formation; Armstrong et al., 1990). Consequently, the Pilbara and Kaapvaal Cratons provide important references against which to compare the results from the older terrains. Furthermore, their improved preservation allows additional constraints to be established.

9.4 CRUSTAL GROWTH 9.4.1 Nd-Hf isotopic constraints The primitive mantle had chondritic Sm/Nd and Lu/Hf ratios because these elements are refractory and lithophile, and so any measured deviations from these chondritic ratios must be due to fractionation within the silicate portion of the Earth. During typical magmatic processes, Sm and Lu are more compatible than Nd and Hf, respectively, and thus melts have lower Sm/Nd and Lu/Hf ratios than their residues. This relationship can be seen in Figure 9.1, where MORB has near-chondritic to subchondritic (melt) and depleted CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 224 mantle (residual) has superchondritic Sm/Nd and Lu/Hf ratios. Continental crust also has subchondritic Sm/Nd and Lu/Hf ratios, and hence, crustal differentiation has fractionated Sm-Nd and Lu-Hf relative to the primitive mantle. Fortunately, Sm and Lu have long-lived radiogenic isotopes that decay to Nd and Hf, respectively, and so Sm/Nd and Lu/Hf fractionations are traceable through time. The details of Nd and Hf isotopes have been presented in Chapter 6. Briefly, the initial Nd-Hf isotopic composition of a rock can be measured relative to the primitive mantle reference (CHUR), and thus it can be determined whether a rock was derived from a source with time-integrated superchondritic (depleted) or subchondritic (enriched) Sm/Nd and Lu/Hf ratios. Depleted and enriched components develop positive and negative initial e(Nd, Hf), respectively, where e is the difference from the time-integrated evolution of the primitive mantle (CHUR). Moreover, the magnitude of initial e(Nd) and e(Hf) are functions of the Sm/Nd and Lu/Hf ratios and the period of isotopic isolation. For instance, modern MORB have very large positive initial e-values (e(Nd) > 8, e(Hf) > 15; Fig. 9.3), and so have been derived from either a moderately depleted source with prolonged isolation or an extremely depleted source with shorter isolation. This concept can be extended to crustal growth because continental crust and depleted mantle have complementary Sm/Nd and Lu/Hf ratios. Hence, by determining the isotopic composition of mantle-derived rocks (e.g., basalts, komatiites) through time, a proxy for mantle depletion can be obtained.

9.4.1.1 Mantle isotope evolution Mantle evolution diagrams are shown in Figure 9.3 where the initial e(Nd, Hf) of juvenile volcanic rocks have been plotted at their time of formation. The data are from the literature (Table 9.1), include only a representative selection of post-3.5 Ga terrains and have been recalculated using the same constants (Appendix 3). Data from komatiites and basalts have been used to compile the diagrams, although it has been necessary to include felsic rocks for the early Archaean. The diagrams are similar to previous compilations (Shirey & Hanson, 1986; Bennett et al., 1993; McCulloch & Bennett, 1993; 1994; Bowring CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 225

& Housh, 1995; Nägler & Kramers, 1998; Vervoort & Blichert-Toft, 1999), except for the addition of new data, and show that the mantle has become steadily more depleted with time (increasing e(Nd, Hf)). This, however, is not supportive of either growth model because isotopic changes are time-dependent. Moreover, there is a profound positive e(Nd) spike in the early Archaean that does not have an obvious corresponding e(Hf) spike, and so the validity of these Nd data has been questioned. Since these ancient samples are critical to defining mantle evolution, they are reviewed here and revised isotopic evolution diagrams emphasising data quality are presented (Fig. 9.4).

15 30 B MORB A MORB 10

20

5 ) ) d f 0 10 N H ( S&W ( e e

-5 S&W 0

-10

-15 -10 0 1000 2000 3000 4000 0 1000 2000 3000 4000 Age (Ga) Age (Ga)

Figure 9.3: A) e(Nd) and B) e(Hf) evolution diagram of ancient and modern rocks (see text for details). Solid line represents approximate time- integrated mantle evolution required for modern MORB source. Dashed line (S&W) represents proposed BSE of Salters & White (1998).

There are two main issues that must be addressed before interpreting the isotopic composition of ancient rocks: are their ages satisfactorily constrained and do their measured isotopic abundances represent original compositions? These issues are particularly relevant for the beginning of the rock record because most early Archaean rocks have been considerably metamorphosed and deformed. All the data assembled here are plotted using their published ages unless a more robust age is available. For example, initial e-values for the Isua supracrustal rocks are calculated for the same depositional age, although estimates of this have varied. Since there are many assumptions that are difficult to validate when CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 226 determining Sm-Nd isochrons, rocks with independent age evidence (e.g., zircon U-Pb) are preferred. Age errors produce smaller initial isotopic variations for basaltic than felsic rocks because their Sm/Nd and Lu/Hf ratios are typically closer to chondritic, although basalts with the wrong age still plot incorrectly in mantle evolution diagrams. Ultramafic rocks have low Sm, Nd, Lu and Hf concentrations and so large amounts of sample have to be dissolved for precise analyses, which may lead to greater analytical errors than for more enriched felsic rocks. Moreover, even slight chemical changes to rocks with low elemental abundances can profoundly change their isotopic compositions. The revised isotope evolution diagrams highlight samples where both the age and initial e(Nd, Hf) can be confidently established (Fig. 9.4; solid circle ), discriminating them from samples where there is less confidence in one or other parameter (open circle). This does not imply errors in the downgraded results, just doubt about the certainty of Nd-Hf isotope ratios or their accredited age.

Acasta gneisses, Slave Province, Canada Zircons in the Acasta gneisses have ancient cores (<4.03 Ga) which provide strong evidence that the gneisses had, at least in part, ancient precursors (Bowring et al., 1999). Establishment of initial e(Nd) has been attributed to these ancient times and defines a maximum of +3.5 at ~4.0 Ga (Bowring et al., 1989; Bowring & Housh, 1995). However, the assumption that zircon core ages correlate with the timing of whole-rock Sm-Nd equilibrium has been questioned by Moorbath et al. (1997). Indeed, the evidence that the gneisses re-equilibrated at ~3.4 Ga is strong, supported by the ages of zircon overgrowths (sample SAB91-37; Bowring & Williams, 1999) and whole grains (Amelin et al., 2000). Thus, there is little evidence to suggest that whole-rock Sm-Nd equilibrium corresponds with the ~4.0 Ga zircon ages, although there can be no doubt that the Acasta gneisses had ancient precursors. The whole-rock Nd and zircon Hf isotopes are consistent with an ancient precursor, but whether this precursor was derived from an extremely depleted source is unclear.

North Atlantic Craton CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 227

The North Atlantic Craton contains extensive outcrops of early Archaean rocks and isotope studies have focussed on the Isua Supracrustal Belt and Am îtsoq gneisses (west Greenland) and the Saglek-Hebron segment (northern Labrador). The Isua Supracrustal Belt i s the largest ancient greenstone belt and has an age between 3.7 and 3.8 Ga, although it is probably closer to 3.77-3.80 Ga (Nutman et al., 1996; Moorbath et al., 1997; Blichert-Toft et al., 1999a). The belt has been metamorphosed to at least mid-amphibolite facies and intensely deformed such that it is difficult to determine the origins of many rock units. For instance, some felsic gneisses are purported to have volcanosedimentary protoliths, despite limited supporting evidence for this interpretation (Jacobsen & Dymek, 1988; Moorbath et al., 1997). Widespread Sm-Nd alteration of mafic rocks, particularly within the extensive garbenschiefer unit, has been documented, possibly downgrading all these analyses (Gruau et al.; 1996; Rose et al., 1996). However, an integrated Nd-Hf study by Blichert-Toft et al. (1999a) used parent- daughter correlations ( 147Sm/144Nd versus 176Lu/177Hf) to show that five mafic samples have similar compositions to modern basalts. Although this in itself does not prove that these five samples have retained their magmatic compositions, it does confirm the modification of the other mafic samples. Importantly, the five non-suspect samples define a straight line in the parent-daughter diagram, consistent with a magmatic trend ( Fig. 2 in Blichert-Toft et al., 1999a; Fig. 1 in Albarède et al., 2000) and supporting their preservation of original Nd-Hf compositions. Furthermore, since they have near-chondritic Sm/Nd and Lu/Hf ratios, they should be little affected by slight age discrepancies. Thus, they are here considered to be the most likely Isua samples to have preserved their initial e(Nd, Hf) with geochronologic integrity (solid circles). The Akilia gabbroic bodies and their host Am îtsoq gneisses have initial e(Nd) = +4.5 to -4.6 and e(Hf) = +3.3 to +0.6 for interpreted ages between 3729 to 3858 Ma (Bennett et al., 1993; Vervoort & Blichert-Toft, 1999). It is difficult to assess these samples as there is little supporting data, although extreme care has been taken to analyse samples that are moderately homogeneous and contain uncomplicated zircons (Bennett et al., 1993). However, these samples may suffer from the same problems as the Acasta gneisses (Moorbath et al., 1997; Vervoort & Blichert-Toft, 1999; Whitehouse et al., 1999), though the evidence for re-equilibration of all the gneisses at the same time is not as convincing as CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 228 for the Acasta gneisses. Moreover, it has been suggested that since the gneisses do not plot within the general terrestrial Nd-Hf mantle-crust array, then they must have been altered (Vervoort & Blichert-Toft, 1999). However, this assumes that the early Earth was controlled by the same processes as the modern Earth, which is not necessarily the case (see discussion of Hf paradox below). But given the array of uncertainties, these samples are here considered to be slightly suspect (open circles). Ultramafic rocks from the Saglek-Hebron segment have initial e(Nd) between +3.1 and -4.0 at ~3.8 Ga (Collerson et al., 1991). These rocks have moderately coherent trace element patterns which are generally consistent with magmatic compositions and provide little evidence of REE alteration at the sampling scale. However, replicate analyses have very large differences (e(Nd) = (+2.3, –0.1), (+1.1, –0.2)), suggesting substantial isotopic heterogeneity within the samples, which might be a reflection of their low Sm-Nd abundances. Regardless, even allowing for the youngest age defined by their isochron (~3.7 Ga), they would still have e(Nd) = +2.2 to +3.2. These samples are also considered to be slightly suspect (open circles). CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 229

Table 9.1: Data sources for isotopic compilations (whole-rock). Terrain Age (Ma) Nd References Hf References Slave 3700 Bowring et al. (1989) Bowring & Housh (1995) Moorbath et al. (1997) 2712 Cousens (2000) North 3800 Collerson et al. (1991) Atlantic 3770 Hamilton et al. (1983) Vervoort & Blichert-Toft (1999) Craton Gruau et al. (1996) Vervoort et al. (1996) Moorbath et al. (1997) Blichert-Toft et al. (1999a) 3754-3872 Bennett et al. (1993) Pilbara 3517-3434 herein herein 3470 Hamilton et al. (1981) Jahn et al. (1981) Gruau et al. (1987) 2684-2764 Nelson et al. (1992) Kaapvaal 3472 Hamilton et al. (1979) Gruau et al. (1990a) Jahn et al. (1982) Blichert-Toft & Arndt (1999) Gruau et al. (1990b) Lécuyer et al. (1994) 3275-3472 Lahaye et al. (1995) 3170-3300 Wilson & Carlson (1989) Hebei 3500 Xuan et al. (1986) Superior 2990 Tomlinson et al. (1998) 2700-2720 Vervoort et al. (1994) Vervoort & Blichert-Toft (1999) 2710 Lahaye et al. (1995) Blichert-Toft & Arndt (1999) Baltic 2795 Puchtel et al. (1998) west Africa 2100 Abouchami et al. (1990) Blichert-Toft et al. (1999a) Cape Smith 1920-1998 Hegner & Bevier (1991) Blichert-Toft & Arndt (1999) Vervoort & Blichert-Toft (1999) Ketilidian 1800 Patchett & Bridgwater (1984) east Africa 743-782 Claesson et al. (1984) ophiolites 100-387 Sharma & Wasserburg (1996) ocean basalts 0 Salters & White (1998) Salters & White (1998) Blichert-Toft et al. (1999b) Blichert-Toft et al. (1999b) CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 230

~3500 Ma terrains There are several exceptionally well-preserved ~3500 Ma greenstone belts where unequivocal initial e(Nd, Hf) can be obtained, and so these can provide a reference for understanding the somewhat ambiguous data from the older terrains. Precise ages are well established in the Pilbara Craton and the Barberton Greenstone Belt of South Africa from zircon dating of felsic volcanics. Moreover, trace element abundances can be confidently determined from magmatic trends of cogenetic samples within thick mafic successions, and hence, strong supporting evidence is available to determine whether a sample has retained its initial e(Nd, Hf). Therefore, only the best supported analyses are included at ~3500 Ma (Fig. 9.4). Samples not plotted include: i) those from stratigraphically beneath the Komati Formation at Barberton as these have been strongly deformed and their ages are poorly constrained (Hamilton et al., 1979; Jahn et al., 1982; Gruau et al., 1990b), ii) komatiites from the Mendon Formation at Barberton as there is significant variation between samples from the same flow and between pyroxene separates (Lahaye et al., 1995), iii) sample B15 of Blichert-Toft & Arndt (1999) as it plots off the trend defined by the other data, and iv) data from Schapenburg, South Africa (L écuyer et al., 1994) and Hebei, China (Xuan et al., 1986) as their ages are defined only by Sm-Nd isochrons.

8 6 A B 6

4 4

2 ) 2 d ) f N ( 0 H e ( e 0 -2

-4 -2

-6 3400 3600 3800 4000 3400 3600 3800 4000 Age (Ga) Age (Ga)

Figure 9.4: A) e(Nd) and B) e(Hf) evolution diagram of Archaean rocks (see text for details). Solid and dashed lines as for Figure 9.3. CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 231

Implications Even after removing all the suspect samples, there are still many early Archaean units with large positive and negative initial e(Nd) (Fig. 9.4). Moreover, the Akilia and Amîtsoq units define a possible positive e(Hf) spike that approaches the time-integrated trajectory to modern MORB, but was not noted in previous studies of larger data sets, excepting Vervoort & Blichert-Toft (1999). Importantly, the five best-constrained Isua samples have diverse initial e(Nd, Hf) at ~3770 Ma indicating a heterogeneous source with time-integrated superchondritic and subchondritic Sm/Nd and Lu/Hf ratios. Although these five samples do not have the largest initial e(Nd), they plot above the time-integrated evolution of modern MORB, as do samples from the Pilbara Craton (~3470 Ma; Gruau et al., 1987), Abitibi Belt (Fig. 9.3A, ~2710 Ma; Lahaye et al., 1995) and Voykar Massif (Fig. 9.3A, ~387 Ma; Sharma & Wasserburg, 1996). Interestingly, all of these younger samples come from ultramafic units, perhaps suggesting a common petrogenetic origin. Further evidence for significant isotopic diversity and extreme depletion in the early Archaean has been reported in Hf isotopes from ancient zircons (Fig. 9.5; Vervoort et al., 1996; Amelin et al., 1999, 2000). Zircons are a particularly powerful data source because precise U-Pb ages can be determined and they preferentially incorporate Hf over Lu, thus minimising the effects of 176Hf ingrowth. The main problems of the method, such as small sample size and inheritance within zircon grains, have been partly addressed, although further technical improvements are likely. Despite this, the zircon data indicate greater Hf depletion of their sources than the whole-rock data. There is strong isotope evidence that many Archaean rocks, including the Pilbara samples, were derived from isotopically enriched and depleted sources. Moreover, the relative diversity seems to have been extreme during the Earth’s earliest history, and perhaps transient (Bennett et al., 1993; Albarède et al., 2000). Thus, the mantle must have been differentiated during its earliest history, either by segregation of sialic crust (Bennett et al., 1993; Bowring & Housh, 1995), mafic crust (Chase & Patchett, 1988; Galer & Goldstein, 1991) or a magma ocean (Collerson et al., 1991; Albarède et al., 2000). Each of these processes can be mathematically modelled to satisfactorily account for the observed isotopic diversity, although the most extreme e(Nd) values at ~3800 Ma are difficult to CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 232 reconcile with a purely mafic crust (Bennett et al., 1993; Bowring & Housh, 1995). However, the identification of an early depleted mantle, even if extremely depleted, does not by itself demonstrate the existence or extent of ancient continental crust, although it is consistent with such a proposition. Additional chemical evidence is required to discriminate between the various alternatives.

9

6

3 ) f 0 H ( e

-3

-6 3400 3600 3800 4000 4200 Age (Ma)

Figure 9.5: e(Hf) evolution diagrams for early Archaean zircons. Data from Vervoort et al., 1996; Amelin et al., 1999, 2000. Line is the same from Figure 9.3.

9.4.2 Nb-U systematics It is widely held that Nb and U are refractory lithophile elements and so the primitive mantle would have had a chondritic Nb/U ratio (~32.4; Sun & McDonough, 1995). However, Nb may be moderately siderophile at high pressures such that up to 23 % of the Earth’s total Nb budget may have been sequestered into the core (Wade & Wood, 2001). If so, the primitive mantle would have had a subchondritic Nb/U ratio (~25.0). Modern mid-ocean ridge and ocean island basalts have remarkably uniform superchondritic Nb/U ratios (47 ± 10) that do not vary as a function of elemental concentration, indicating that Nb and U are not fractionated during typical mantle melting (Hofmann et al., 1986; Sims & DePaolo, 1997). Consequently, the modern mantle has the same Nb/U ratio as modern oceanic basalts (47 ± 10; Hofmann et al., 1986). In contrast, continental crust has very low subchondritic Nb/U ratios (~5.6; Nb from Barth et al., 2000; U from Rudnick & CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 233

Fountain, 1995), and thus balances the depleted composition of the modern mantle with respect to the primitive mantle. This complementary relationship is also shown in Figure 9.1, where continental crust has a negative Nb and positive U anomaly, and the depleted mantle and MORB have the converse. Thus, the increase in the Nb/U ratio of the mantle from its primitive value (25.0 to 32.4) to the modern depleted value (47 ± 10) almost certainly reflects the extraction of continental crust. Significantly, the incorporation of ~23 % of the Earth’s Nb into the core (Wade & Wood, 2001) lowers the Nb/U ratio of the primitive mantle (32.4 to 25.0), but does not negate the necessary Nb/U balance between the crust and depleted mantle because of the rapid and very early segregation of the core. The various global reservoirs are shown in Figure 9.6A, including two estimates of the primitive mantle composition depending on whether Nb was sequestered into the core.

100 70 A OIB B MORB 60

DM 50 C1 pyrolite 40 U / U 23%core / a b

b 30 N $ N upper 10 b$ continental 20 crust (1995) UCC (2000) continental Pilbara 10 pelagic crust Barberton sediments 1 1 0 0.1 1 10 100 1000 0 1000 2000 3000 4000 [Nb] Age (Ma)

Figure 9.6: A) Nb/U versus [Nb] with data from Pilbara basalts (Pilgangoora herein & North Pole from Green et al., in prep.) and Barberton komatiites and basalts (Lahaye et al., 1995; Dann, unpubl. data), fields for modern ocean island basalts (OIB) and mid-ocean ridge basalts (MORB) (Sims & DePaolo, 1997), shaded field is range of modern oceanic basalts (47 ± 10; Hofmann et al., 1986), DM = depleted mantle, C1 chondrite and pyrolite model (McDonough & Sun, 1995), 23 % core (Nb sequestered into core; Wade & Wood, 2001), upper continental crust (Rudnick & Fountain, 1995; new Nb value Barth et al., 2000), pelagic sediments (Taylor & McLennan, 1985) adapted from Sims & DePaolo, 1997. B) Nb/U evolution of basalts and range of modern values (shaded area), a $ = chondritic, b$ = 23 % core (refs. as for A) see text for data sources; adapted from Collerson & Kamber, 1999.

9.4.2.1 Mantle Nb/U evolution CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 234

Given that Nb/U ratios of depleted mantle inversely correlate with those of continental crust, then it follows that a greater continental mass should correspond with increasing Nb/U ratios in the mantle, if the volume of depleted mantle has remained constant. Therefore, by determining the Nb/U ratio of the mantle through time, the extent of mantle depletion, and hence the continental mass, can be constrained. Fortunately, mantle Nb/U ratios can be estimated from ancient basalts because they are derived from the mantle and Nb and U are not fractionated during basalt petrogenesis. To adequately assess ancient Nb/U ratios, the effects of post-magmatic alteration and crustal contamination must be removed. Alteration can be assessed by showing that Nb and U define magmatic trends for an entire volcanic succession (Chapter 4.5), after excluding anomalous samples from the database. Crustal contamination may lead to lower Nb/U ratios and so a simple average from a contaminated succession will underestimate the mantle Nb/U ratio. Instead, basalt suites can be plotted in entirety (e.g., Fig. 9.6A) and the largest Nb/U ratios taken to represent the mantle composition. Crustal contamination is evident in the ~3.5 Ga Pilbara and Barberton greenstones, where the least contaminated samples plot at high Nb/U ratios, similar to modern ocean basalts, and define a contamination trend towards continental crust (Fig. 9.6A, also see Fig. 4.10). A mantle Nb/U evolution diagram (Fig. 9.6B) has been defined using the same data sources as Collerson & Kamber (1999), except the diagram has been extended to older terrains and now includes the ~3517 to 3434 Ma Pilbara (Pilgangoora (herein), North Pole (Green et al., in prep.)), ~3470 Ma Barberton (Lahaye et al., 1995; Jesse Dann, unpubl. data) and ~2840 Ma Kostomuksha greenstones (Puchtel et al., 1998). Moreover, the two oldest values used by Collerson & Kamber (1999), and indeed the most important values defining their inflected trends, have been reassessed. Their Nb/U ratio for the ~2970 Ma Red Lake greenstones (Tomlinson et al., 1998) has been recalculated to ~40 as the previous value of 33.4 is clearly an average of samples that have been contaminated, and their ~3300 Ma average value includes samples from the ~3470 Ma Barberton greenstones (Jochum et al., 1991). In addition, all average values for specific times, as plotted by Collerson & Kamber (1999), have been excluded as they are probably biased by altered or contaminated samples, thus not measuring the mantle Nb/U ratio. Three samples have also been precisely measured from the ~3470 Ma Marble Bar Belt, Pilbara (Jochum et al., 1991), but are not CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 235 included in the compilation presented here as there are too few samples to assess contamination and alteration. The evolution diagram shows that all of the ~3500 Ma basalt suites were derived from a mantle source with Nb/U ratios indistinguishable from modern depleted mantle (47 ± 10, shaded area Fig. 9.6B). Moreover, uniformly high Nb/U ratios have been maintained since ~3500 Ma, in contrast to the inflected trend proposed by Collerson & Kamber (1999). Indeed, there is little evidence for a profound Nb/U change in the late Archaean associated with the oxidation of the atmosphere and hydrosphere (Collerson & Kamber, 1999). Instead, the ancient mantle was as depleted in Nb/U as the modern mantle, suggesting that a similar mass of continental crust had been extracted. Significantly, the ~3517 Ma Coonterunah succession may be the oldest from which Nb/U can be satisfactorily determined since all older greenstone belts have been intensely metamorphosed. The Nb/U data do not uniquely constrain the total volume of depleted mantle since ancient basalts may have preferentially sampled small depleted reservoirs or the size of the depleted mantle may have grown as continental crust was progressively extracted. In either case, the reservoirs must have been isolated because mixing with primitive mantle (Nb/U = 25 to 32) would have lowered depleted mantle Nb/U ratios. This requires a mechanism for preventing small depleted reservoirs from homogenising in a hotter Archaean mantle and a process whereby continental crust can be extracted from the remaining primitive mantle, but the primitive mantle is never sampled by ancient basalts. This seems extremely unlikely given the diversity of tectonic settings for ancient greenstone successions.

9.4.3 Synthesis Nd-Hf isotopes and Nb/U ratios provide independent evidence for the existence of an ancient depleted mantle. However, Nd-Hf isotopes cannot readily distinguish whether depletion was associated with differentiation of continental or mafic crust or the formation of a magma ocean. However, Nb and U are not fractionated during basalt petrogenesis, and so ancient superchondritic Nb/U ratios cannot be solely due to the formation of mafic crust. Furthermore, formation of a magma ocean probably also did not significantly fractionate Nb and U as they have similar partition coefficients for garnet and perovskite, the phases most likely to have crystallised from a magma ocean (Kato et al., 1988a, b). Thus, the Nd-Hf CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 236 isotopes and Nb/U ratios of Archaean rocks are most explicable in a model whereby a mass of continental crust roughly equivalent to its modern volume was extracted from the mantle in the Earth’s earliest history. Thus, a steady-state model regulated by post-Archaean continental addition coupled to compensatory recycling is most tenable.

9.5 CRUSTAL GENESIS 9.5.1 Nd-Hf isotopic correlation The initial Nd-Hf isotopic compositions of mantle and crustal rocks from both modern and ancient terrestrial terrains show a strong positive correlation illustrating the typical covariant behaviour between these two isotopic systems (Chapter 6.5.4.3; Fig. 6.7, 9.7). However, the terrestrial array does not correspond with the composition of the bulk silicate Earth (BSE), a feature known as the Hf-paradox (Salters & Hart, 1989). Three explanations of this paradox have been proposed: a complementary reservoir exists within the Earth, estimates of the BSE are incorrect, or the Earth does not have a chondritic composition (Salters & Hart, 1989; 1991; Blichert-Toft & Albarède, 1997; Salters & White, 1998; Blichert-Toft & Arndt, 1999; Vervoort et al., 1999). The issue of a non-chondritic Earth, though favoured by Blichert-Toft & Arndt (1999), is not discussed here because such a finding conflicts with a vast array of geochemical and cosmochemical data and would profoundly alter our present understanding of planetary formation. Interestingly, the Pilbara basalts (Chapter 6.5.4.3) and some other early Archaean rocks (Vervoort & Blichert-Toft, 1999; Blichert-Toft et al., 1999a; Albarède et al., 2000) have initial Nd-Hf compositions that are adjacent to the terrestrial array (Fig. 9.6B), suggesting that these ancient rocks have either been modified by post-magmatic processes (Vervoort et al., 1996; Vervoort & Blichert-Toft, 1999) or that they sampled reservoirs with memories of distinct fractionation processes (Albarède et al., 2000). This anomalous behaviour may be related to the Hf-paradox and thus may have important implications for crustal genesis. The Pilbara basalts are crucial to this debate because they clearly have not been altered and so their anomalous Nd-Hf behaviour is almost certainly an original feature. CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 237

10 low-Ti lunar low-Ti lunar (a) 40 A B low-K high-Ti lunar

martian 5 terrestrial array (b) 20 chondrites (e) low-K high-Ti lunar (c)

) $ ) early Archaean (d) f

f ?complement 0 0 H H b e a ( ( c e residue e S&W(1998) 5 (>Lu/Hf) $ -20 d Isua (mafic) -5 Isua (felsic) Amîtsoq -5 -40 melt Barberton -5 5 (

Figure 9.7: A) eNd-eHf diagram shows the general trends of various planetary reservoirs; terrestrial array (references in text), early Archaean gneisses (Vervoort & Blichert-Toft, 1999), chondrites (Salters & White, 1998), lunar basalts (Beard et al., 1998) and SNC Martian meteorites (Blichert- Toft et al., 1999c). Lines are extrapolations and do not represent the range of data arrays. B) Early Archaean samples; Isua (Blichert-Toft et al., 1999a), Amîtsoq gneisses (Vervoort & Blichert-Toft, 1999), Barberton (Blichert-Toft & Arndt, 1999) and Pilbara (herein). Star represents the BSE estimate of Salters & White (1998).

9.5.1.1 Hf-paradox - complementary reservoir If the present isotopic estimates of the BSE are correct (Jacobsen & Wasserburg, 1980; Goldstein et al., 1984; Blichert-Toft & Albarède, 1997), then there must exist a reservoir with positive e(Nd) and negative e(Hf) to complement the terrestrial array (Fig. 9.6A). Some early Archaean rocks have such compositions (Fig. 9.6B) and so they may have sampled this complementary reservoir. If so, the reservoir must have become hidden because there is no evidence for its later existence despite investigation of a diverse range of oceanic settings (Salters & White, 1998), ancient terrains (Vervoort & Blichert-Toft, 1999; Blichert-Toft & Arndt, 1999), granites (Vervoort & Patchett, 1996), sediments (Patchett et al., 1984; Vervoort et al., 1999a) and lower crustal xenoliths (Vervoort et al., 2000). The origin and composition of this reservoir can be inferred from its complementary Nd-Hf composition. Geochemical differentiation into terrestrial and complementary reservoirs indicates that there were significant events during early Earth history where the typical covariance between Nd and Hf was decoupled. Such decoupling suggests the presence of garnet during CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 238 melting because Lu is strongly compatible and Sm, Nd and Hf are highly incompatible in garnet (Irving & Frey, 1978; Thirlwall et al., 1995; Vervoort & Patchett, 1996; Vervoort et al., 1999a, 2000). Involvement of other mantle minerals such as perovskite may also decouple Nd-Hf, but the partitioning is analogous to garnet, and so garnet can be considered to be representative of these minerals (Kato et al., 1988a, b; Blichert-Toft & Albarède, 1997; Salters & White, 1998). Hence, melting in the presence of garnet or crystallisation of garnet from a magma will leave restites with greater Lu/Hf ratios, magmas with smaller Lu/Hf ratios and there will be no corresponding change to the Sm/Nd ratios. As a result, e(Hf) will increase faster in the restite and slower in the melt compared with typical Nd-Hf behaviour (Fig. 9.6B). In this instance, the terrestrial array correlates with the restites and the hidden reservoir with the magmas. Thus, a plausible melt-segregation model would require the formation of a garnet-depleted layer that was subsequently buried and hidden in the mantle, with most currently accessible terrestrial rocks derived from the garnet-rich residue. The garnet-depleted layer would be less dense than a garnet-rich residue and so would need to cool, solidify and become more dense in order to sink and then remain in the deep mantle. In other words, the garnet-depleted layer would have formed in the upper part of the early terrestrial mantle, perhaps as a crust. Such a model, however, is problematic. First, a garnet-depleted layer would be hard to bury. For instance, cool basaltic crust is readily returned to the mantle, but typically converts to garnet amphibolite while sinking into the mantle, and thus the buried layer becomes garnet-rich with the opposite Lu/Hf ratio to the complementary reservoir. In addition, sialic crust is very buoyant and thus very difficult to bury en masse into the mantle. Nevertheless, if a garnet-depleted layer with the necessary subchondritic Lu/Hf ratio was sequestered into the deep mantle it would be difficult to preserve because the layer would incorporate relatively greater amounts of heat- producing elements (K, Th, U) than the surrounding rock, and thus would gradually heat up and become more buoyant. Incorporation of K, Th and U into a garnet-depleted layer would be particularly pronounced during differentiation of primitive mantle, prior to enrichment of these elements into continental crust. Moreover, such preservation would be especially difficult in a young, hot, vigorously convecting mantle. The greatest difficulty for the model, however, is conceiving a process whereby the garnet-depleted layer is hidden CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 239 from subsequent magmatic sampling given that modern ocean island basalts sample a variety of distinct ancient domains sequestered in the mantle (Zindler & Hart, 1986; Salters & White, 1998), including oceanic crust (Hofmann & Jochum, 1996; Lassiter & Hauri, 1998; Sobolev et al., 2000), oceanic plateaux (Gasperini et al., 2000) and pelagic sediments (Hauri, 1996; Blichert-Toft et al., 1999b). Therefore, it seems unlikely that a reservoir with subchondritic Lu/Hf ratios could have formed, passed into the deep mantle intact and then remained isolated and untapped for ~3.5 billion years.

9.5.1.2 Hf-paradox - reappraisal of BSE A more tenable solution to the discrepancy between the BSE and terrestrial ar ray is that the BSE estimate is incorrect. This issue has been explored by Salters & White (1998), who argue that if terrestrial planets are broadly chondritic, then the Earth should have a chondritic Nd-Hf composition. However, there have been no systematic Nd-Hf studies of chondritic meteorites. For instance, only one meteorite from the Nd study (Jacobsen & Wasserburg, 1980, 1984) was also analysed for Hf (Blichert-Toft & Albarède, 1997). To overcome this shortcoming, Salters & White (1998) grouped the chondrites into their compositional classes to determine an approximate chondrite Nd-Hf trend, and then proposed that the BSE can be estimated from the intersection between the terrestrial array and the chondritic trend (Fig. 9.6A; 143Nd/144Nd = 0.51259, 176Hf/177Hf = 0.2828, 147Sm/144Nd = 0.1952, 176Lu/177Hf = 0.0335). If this finding is supported by a truly systematic study of chondritic meteorites, then the BSE composition will change for both Sm-Nd and Lu-Hf, removing the need to invoke a complementary reservoir. Importantly, the revised BSE compositions do not significantly affect the initial e(Nd, Hf) values of ancient rocks, and so the conclusions derived from the mantle evolution diagrams are unchanged (dashed lines in Fig. 9.3). A revised BSE composition is supported by Nd-Hf data from lunar basalts and early Archaean rocks. All of the inner rocky planets can be considered to be differentiated chondrites (silicate mantle–metallic core), and so, like the Earth, Nd-Hf fractionations were the result of differentiation within their silicate portions. Hence, the lunar and Martian trends should intersect the chondrite trend at their respective primitive Nd-Hf compositions (Fig. 9.7A). Assuming Sm/Nd and Lu/Hf were not fractionated prior to or during the early CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 240 stages of planetary formation, then the intersection points should coincide with the BSE composition, unless the planets accreted from substantially different chondrite mixtures. A common initial composition would be especially likely if the Moon and Earth were derived from the same chondritic mass, as proposed in the giant impact theory (Hartmann, 1986; Lee et al., 1997). Two distinct lunar trends have been defined which indicates the failure of the lunar mantle to homogenise after differentiation of a magma ocean (Fig. 9.7A; Patchett & Tatsumoto, 1981; Unruh et al., 1984; Beard et al., 1998). The lunar trends (a, c) intersect the chondrite trend (e) either side of the terrestrial array (b), and both intersect at higher initial e(Hf) and lower initial e(Nd) than the present BSE composition (Fig. 9.7A). The Martian shergottites define a trend, albeit with only five points, and one point forces the trend to the positive e(Nd) side of BSE (Fig. 9.7A; Blichert-Toft et al., 1999c). Nevertheless, the Moon provides evidence, admittedly on limited data, that it was derived from material with Nd-Hf compositions quite dissimilar to the present BSE estimate, but relatively close to the estimate of Salters and White (1998). Early Archaean rocks should most closely resemble the BSE composition because they have had less opportunity to fractionate and less time to develop radiogenic compositions. Interestingly, many early Archaean rocks lie off the terrestrial array, with others plotting near the proposed BSE of Salters & White (1998). An early Archaean trend (d) can be defined that intersects the chondrite trend (e) almost coincidental with the mantle-crust array (b; Fig. 9.7A). This trend is similar to that outlined by Vervoort & Blichert-Toft (1999). Importantly, the negative e(Nd), positive e(Hf) part of the trend is anchored by a least-suspect mafic sample from Isua (see 9.4.1.2). The present BSE plots just within the field defined by the early Archaean samples and so is not excluded by the data. However, the Nd-Hf compositions of the early Archaean samples suggest that they were derived from a primitive mantle with a composition more similar to the revised BSE proposed by Salters & White (1998). In summary, it is probably no coincidence that all of these trends inter sect the chondrite trend at higher initial e(Hf) and lower initial e(Nd) than the present BSE estimate. Therefore, the Hf-paradox is probably a result of an incorrect estimate of the BSE, and thus there is no need for a complementary reservoir. CHAPTER 9 DISCUSSION – MANTLE AND CRUSTAL EVOLUTION 241

9.5.1.3 Early Archaean Nd-Hf decoupling As the early Archaean trend can be defined by the Pilbara and some Isua basalts and these rocks have probably preserved their magmatic compositions, then the trend is probably real. As a result, discordance with the mantle-crust array cannot be used as a criterion to claim that an ancient sample has been altered, and so the Am îtsoq gneisses may well have preserved their early Archaean initial isotopic compositions (Vervoort et al., 1996; Vervoort & Blichert-Toft, 1999). Nevertheless, the Pilbara and Isua basalts indicate that they were derived from a source that had a previous history of Nd-Hf decoupling. Since decoupling can be related to various petrogenetic processes (garnet fractionation), then maybe early Archaean petrogenesis can be further constrained. Modern examples of Nd-Hf decoupling have been noted for MORB where residual garnet is present during partial melting at deep ridges (e.g., Australian-Antarctic Discordance; Salters, 1996). Given that the young Earth was hotter and so there was less likelihood of small degrees of partial melting, then mafic magmas with residual garnet are unlikely to have been significant in the early Archaean. Anomalous Nd-Hf behaviour has also been documented in island arcs where subduction inputs are mixed (Pearce et al., 1999), but this requires the isolation of distinct domains for prolonged periods to develop marked isotopic heterogeneities. In short, modern geological processes provide little guidance for explaining the early Archaean Nd-Hf trend. Instead, the early Archaean samples have probably preserved the isotopic signature of a significant differentiation event quite dissimilar to typical modern processes. Such an event may have been associated with the segregation of continental crust or the formation of a mantle magma ocean. Notably, preserved early Archaean continental crust contains a disproportionately large amount of rocks with TTG affinities (Condie, 1993), which probably formed by partial melting of garnet amphibolite (see Chapter 5.3.2.1). So, if large-scale continental crust formation produced voluminous TTG magmas, then this process may have caused Nd-Hf decoupling in the early Archaean rock record.

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