Tectonic Evolution of the Early Archaean Doolena Gap Greenstone Belt, East Pilbara Terrane, Western Australia

TECTONIC EVOLUTION OF THE EARLY ARCHAEAN DOOLENA GAP GREENSTONE BELT, EAST PILBARA TERRANE, WESTERN AUSTRALIA

Daniel Wiemer (Dipl. Geol./M.Sc.)

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Submitted in total fulfillment of the requirements for the degree of

Doctor of Philosophy (Ph.D.)

Brisbane, 2017

School of Earth, Environmental and Biological Sciences Science and Engineering Faculty Queensland University of Technology

Supervisor: Dr. David T. Murphy Co-supervisor: Dr. Christoph E. Schrank

QUT Verified Signature

iii Abstract

Earth’s oldest preserved crust is characterised by a distinctive crustal architecture of large ovoid granitic domes surrounded by curvilinear supracrustal greenstone belts. Such dome-and-keel morphology is not generated in the present-day Earth as part of plate tectonics. This indicates that the crustal fragments preserved from the early Earth, which had higher heat generation and therefore higher temperatures, formed in a distinct non-plate tectonic style tectono-magmatic regime. Here I investigate the lithological associations, the magmatic contributions and the structural kinematics that gave rise to the distinctive Archean dome and keel morphology. In doing so I test the model of partial convective overturn for the formation of the dome-and-keel crustal architecture in which dense mafic crust overlying lower density felsic crust leads to buoyancy instabilities and crustal-scale reorganization. The 3530- 3225 Ma East Pilbara Terrane, Western Australia, contains the best preserved lithological and structural inventory associated with dome-and-keel formation, making it the archetypical early Archean dome-and-keel terrane and the perfect natural laboratory for investigating early Archaean tectonic processes. Here I present a detailed lithostratigraphic, structural, petrological, geochemical, and geochronological study of a well exposed geological transect through the dome-and-keel terrane of the transition between the western Doolena Gap greenstone belt and the Muccan Granitic Complex. I aim to address the following fundamental research problems and questions: i) What is the lithostratigraphic and structural anatomy of a well exposed transect between little altered lowermost greenschist-facies lithologies in the Doolena Gap greenstone belt and granitiod gneisses of the adjacent Muccan Granitic Complex and how do they relate to dome-and-keel formation? ii) Is there evidence for the development of a buoyantly unstable crustal configuration, such as significant felsic crustal components that predate mafic volcanism and lateral thickness and compositional variations of greenstone formations? iii) When did the major phase(s) of partial convective overturn that led to the dome-and-keel morphology occur? I have identified a deformation event that includes tight fold development in both the Doolena Gap greenstone belt and marginal gneisses of the Muccan Granitic Complex that predates the formation of a 45o unconformity upon which the 3427-3350 Ma Strelley Pool Formation was deposited. The nature and kinematics of this deformation event are consistent with dome-up and keel-down relative motion and indicates a partial convective overturn event between 3460 Ma, the lower age of the Duffer Formation of the greenstone belt and 3427 Ma, the minimum age of the Strelley Pool Formation, that precedes the regional 3310 Ma partial convective overturn event in the East Pilbara Terrane. I have identified a new suite of co-magmatic granitic gneisses, the Doolena Suite (ca. 3500-3590 Ma) along the margin of the Muccan granitic Complex that represents the oldest coherent lithological formation in the East Pilbara Terrane and predates the oldest greenstone belts. This demonstrates that felsic crust was present in the precursor crust upon which the Pilbara Supergroup was deposited.

Keywords: East Pilbara Terrane, Archaean tectonics, dome-and-keel, crustal evolution, Doolena Gap greenstone belt;

iv Publications

Chapter 2: Wiemer, D., Schrank, C. E., Murphy, D. T., and Hickman, A. H., 2016: Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt, East Pilbara Terrane, Western Australia. Precambrian Research, 282, 121-138

Conference abstracts

Wiemer, D., Allen, C. M., Murphy, D. T., and Kinaev, I., 2017: Effects of thermal annealing and chemical abrasion on the U-Pb systematics, and the microstructure of complex, metamict ~3.5 billion year old zircon: insights for U-Pb LA-ICP-MS dating. AMAS 14th Biennial Symposium, Brisbane

Allen, C. M., Wiemer, D., and Murphy, D. T., 2016: Improving LA-ICP-MS dating techniques: experiments on zircon from a 3.51 Ga dioritic gneiss, East Pilbara Terrane, Western Australia. GSA Annual Meeting, Denver

Wiemer, D., Schrank, C. E., Murphy, D. T., and Hickman, A. H., 2015: Structural development of the early Archaean Doolena Gap greenstone belt, East Pilbara Terrane (Western Australia). SGTSG, Caloundra

Burke-Shyne, D., Wiemer, D., Schrank, C. E., and Murphy, D. T., 2015: Carbonate alteration is the dominant weakening mechanism in the Doolena Gap Greenstone Belt preserving the keel rock in a dome-and-keel terrane. SGTSG, Caloundra

Murphy, D. T., Trofimovs, J., Hepple, R. A., Wiemer, D., Kemp, A. I. S., and Hickman, A. H., 2015: Pillow basalts from the Mount Ada Basalt, Warrawoona Group, Pilbara Craton: implications for the initiation of granite-greenstone terrains. Goldschmidt, Prague

Wiemer, D., Schrank, C. E., and Murphy, D. T., 2014: Lithostratigraphy and structure of the early Archaean Doolena Gap Greenstone Belt, East Pilbara Terrane (EPT), Western Australia. AGU Fall Meeting, San Francisco

Acknowledgements

I highly acknowledge my supervisory team David Murphy and Christoph Schrank for their trust, patience, guidance and mentoring. Thank you! You are both remarkable personally and as scientists. I thank Charlotte Allen for her assistance and mentoring in zircon geochronology and for giving me shelter for a while. Scott Bryan is acknowledged for critically reviewing various stages of my PhD research. David Gust is thanked for letting me take part in teaching of undergraduate students. Arthur Hickman and the Geological Survey of Western Australia are highly acknowledged for fieldwork logistics. Critical review by Arthur Hickman has greatly improved this thesis, particularly the second Chapter. I thank the administrative backbone of the QUT system: Sarie Gould, Courtney Innes, Noelene Davis, Tiziana La Mendola and others. I highly acknowledge following QUT staff members and colleagues for training and assistance in various analytical methods: Aarshi Bhargav, Henny Cathey, Karine Moromizato, Llew Rintoul, Gus Luthje, Donald McAuley, Irina Kinaev, Sanjleena Singh, Peter Hines, Mitch DeBruyn, Chris East, Linda Nothdurft, Shane Russell, Natalia Danilova, Alex Hepple, David Steele. I acknowledge Charles Verdel, Gang Xia and Dario Hogg from the University of Queensland for guidance and assistance in lab work I highly acknowledge John De Kruijff, Thomas Spring, Shosh O’Connor, Ali Sternes and Duncan Burke-Shyne for helping hands during my fieldwork. Thanks to the Ultimate Frisbee Team and the EEBS Seminars I thank all my housemates for tolerating me over the years: Joania, Yazu, Ella, Julia, Carol, Joakim, Bronwen, Charlotte & David, and James. Finally, my family, friends and mates – I love You!, You know who you are!

“The Earth is not forever, But just to remain for a short while…”

(Scott R. Weinrich, Dreamwheel, 1999)

vii

viii TABLE OF CONTENTS Page

Declaration of Original Authorship iii Abstract iv Publications v Acknowledgements vi List of Figures xi List of Tables xii ______

CHAPTER 1 1 INTRODUCTION 1.1 Research Background 1 1.1.1 Earth’s thermal evolution, and the theoretical consequences for Archaean geodynamics 5 1.1.2 Archaean dome-and-keel terranes 6 1.1.2.1 Model for dome-and-keel formation in the East Pilbara Terrane 7 1.2 Research Problems and Questions 7 1.3 Aim of the Research 10 1.4 Research Philosophy and Methodology 11 1.5 Outline of Chapters 11 ______

CHAPTER 2 13 LITHOSTRATIGRAPHY AND STRUCTURE OF THE EARLY ARCHAEAN DOOLENA GAP GREENSTONE BELT, EAST PILBARA TERRANE, WESTERN AUSTRALIA 2.1 Abstract 13 2.2 Introduction 15 2.3 Regional Geology and Tectonic Framework of the East Pilbara Terrane 16 2.4 Geological Setting of the Study Area 18 2.5 Lithostratigraphy and Structure 21 2.5.1 Muccan Granitic Complex 21 2.5.1.1 Lithological components and structure within the Muccan Granitic Complex 21 2.5.1.2 Deformational micro-textures and metamorphism within the Muccan Granitic Complex 23 2.5.2 South Muccan Shear Zone 25 2.5.2.1 Lithological components and structures within the South Muccan Shear Zone 25 2.5.2.2 Deformational micro-textures and metamorphism within the South Muccan Shear Zone 25 2.5.3 Central Fold Belt 27 2.5.3.1 Lithological components within the Central Fold Belt 27 2.5.3.2 Structures and deformational events within the Central Fold Belt 27 2.5.3.3 Micro-textures, deformational intensity and metasomatism within the Central Fold Belt 30 2.5.4 Southern Low-Strain Belt 32 2.5.4.1 Lithological components and structures within the southern Low-Strain Belt 32 2.5.4.2 Micro-textures, deformation and metamorphism within the Low-Strain Belt 35 2.5.4.3 Lithostratigraphy of the Low-Strain Belt 35 a) Mount Ada Basalt 35 b) Felsic volcanic – sedimentary succession (Duffer Formation) 37 c) Strelley Pool Formation 39 2.6 Discussion 39 2.6.1 Deformation history 40 2.6.2 Structural evolution and tectonically controlled deposystems 43 2.6.2.1 Depositional environment and early extension 43 2.6.2.2 Main phase of granitic doming and associated greenstone sagduction within the western 44 Doolena Gap greenstone belt 2.6.2.3 Progressive dome-and-keel formation: partial convective overturn versus horizontal 47 tectonics 2.6.3 Implications for early Archaean tectonics 48 2.7 Conclusions 50 ______

CHAPTER 3 53 FORMATION, DIFFERENTIATION AND REWORKING OF EARLY (ca. 3500-3590 Ma) EAST PILBARA TERRANE CONTINENTAL MATERIAL AND IMPLICATIONS FOR DOME-AND-KEEL INITIATION – EVIDENCE FROM THE WESTERN DOOLENA GAP GREENSTONE BELT 3.1 Abstract 53 3.2 Introduction 55 3.3 Brief Geology and Tectonic Model of the East Pilbara Terrane 57 3.4 Study Area and Sample Selection 60 3.5 Analytical Methods 61 3.6 Petrography 64 3.6.1 Dioritic to granodioritic gneisses 65 3.6.2 Tonalitic-trondhjemitic and granitic gneisses 65 3.6.3 In situ leucosomes, late-stage pegmatites and leucocratic domains of migmatites 66 3.7 Geochemistry 68 3.7.1 Major elements 68 3.7.2 Trace elements 71 3.7.2.1 Rare earth elements (REE) 71 3.7.2.2 High field strength (HFSE), large ion lithophile (LILE) and other trace elements 73 3.8 U-Pb Zircon LA-ICP-MS Geochronology 74 3.8.1 Muccan Granitic Complex gneiss components 77 3.8.1.1 Dioritic gneiss 370-1 77 3.8.1.2 Granodioritic gneiss 309-2 78 3.8.1.3 Tonalitic-trondhjemitic gneiss 310-4 78 3.8.1.4 Granitic gneiss 308-1 78 3.8.2 Detrital zircon from the quartz-sandstone 79 3.8.2.1 Quartz-jasper base conglomerate 332-1 and quartz-sandstone 335-1 79 3.9 Discussion 81 3.9.1 Petrogenetic constraints on the formation of the 3502-3583 Ma Doolena Suite gneisses 81 3.9.1.1 Formation of parental continental magma through partial melting of hydrous metabasalt 83 a) Constraints from thermodynamic modeling (Theriak-Domino) 83 b) Constraints from geochemical trace element modeling 84 3.9.1.2 Derivation of compositional types: fractional crystallization versus depth and degree of 87 partial melting 3.9.2 Reworking and un-roofing of the Muccan Granitic Complex – implications for early (>3427 89 Ma) dome-and-keel initiation 3.9.2.1 Time constraints and detrital provenance 89 3.9.2.2 Constraints on reworking mechanism and dome initiation 91 3.9.3 Implications for early continental crust formation in the East Pilbara Terrane 94 3.10 Conclusions 96 ______

CHAPTER 4 97 GEOCHEMISTRY OF THE ca. 3470-3460 Ma MOUNT ADA BASALT AND DUFFER FORMATION IN THE WESTERN DOOLENA GAP GREENSTONE BELT (EAST PILBARA TERRANE, WESTERN AUSTRALIA): IMPLICATIONS FOR EARLY ARCHAEAN UPPER CRUSTAL GROWTH 4.1 Introduction 97 4.2 Brief Geology and Previous Constraints on the Warrawoona Group Greenstone Volcanic 100 Rocks 4.3 The Doolena Gap Greenstone Belt Transect: Field Relationships and Sample Selection 102 4.3.1 Low-Strain Belt 103 4.3.2 Central Fold Belt and South Muccan Shear Zone 104 4.3.3 Sample selection 105 4.4 Analytical Methods 107 4.5 Petrography and geochemistry of Rocks of the Duffer Formation and Mount Ada Basalt 107 from the Low-Strain Belt

x 4.5.1 Petrography 107 4.5.1.1 Mount Ada Basalt 107 4.5.1.2 Duffer Formation 108 4.5.2 Major and trace element geochemistry 109 4.5.2.1 Mount Ada Basalt 110 a) High-MgO group 111 b) Low-MgO group 112 4.5.2.2 Duffer Formation 113 4.6 Petrography and Geochemistry of Mafic Rocks from the Central Fold Belt and South 115 Muccan Shear Zone 4.6.1 Petrography 115 4.6.1.1 Central Fold Belt 115 4.6.1.2 South Muccan Shear Zone 117 4.6.2 Major and trace element geochemistry 117 4.6.2.1 Mafic rocks from the Central Fold Belt and South Muccan Shear Zone 117 4.7 Discussion 121 4.7.1 Effects of alteration and metamorphism/metasomatism 123 4.7.1.1 Assessment of least altered geochemical signatures 124 4.7.2 Petrogenetic relationship between the Duffer Formation and the Mount Ada Basalt in the 127 Low-Strain Belt 4.7.3 Petrogenetic processes and correlation between mafic rocks from the Low-Strain Belt and 128 mafic rocks from the Central Fold Belt and South Muccan Shear Zone 4.7.3.1 Model I: distinct mantle sources 130 4.7.3.2 Model II: significance of intra-crustal processes – differentiation and assimilation 132 4.8 Conclusions 134 ______

CHAPTER 5 137 SYNTHESIS AND DISCUSSION 5.1 Summary of Key Findings of the Thesis 137 5.2 Discussion 139 5.2.1 Tectonic evolution of the western Doolena Gap greenstone belt 139 5.3 Remaining Questions and Future Research Directions 144 ______

References 147 Appendices 159 Appendix 1: AGU Fall Meeting 2014 abstract Appendix 2: Goldschmidt 2015 abstract Appendix 3: SGTSG 2015 abstract Appendix 4: SGTSG 2015 abstract Appendix 5: GSA Annual Meeting 2016 abstract Appendix 6: AMAS Biennial Symposium 2017 abstract Appendix 7: Table 3.A-I: XRF detection limits Appendix 8: Table 3.A-II: ICP-MS detection limits Appendix 9: Table 3.A-III: Zircon chemical data (LA-ICP-MS)

List of Figures Chapter 1 Fig. 1.1: Thermal evolution curve Fig. 1.2: mobile and stagnant lid regime Chapter 2 Fig. 2.1: Simplified geological map of the East Pilbara Terrane Fig. 2.2: Generalized stratigraphy of the Pilbara Supergroup Fig. 2.3: Geological map of the Doolena Gap – Marble Bar greenstone synclinorium Fig. 2.4: Geological-structural map of the studied ‘1 Mile Creek’ study area Fig. 2.5: Geological cross sections through the studied granite-greenstone traverse Fig. 2.6: Field photograph and stereo MGC

xi Fig. 2.7: Petrographic image MGC Fig. 2.8: Detailed structural map (1) Fig. 2.9: Petrographic images SMSZ Fig. 2.10: Field photographs CFB Fig. 2.11: Detailed structural map (2) Fig. 2.12: Petrographic images CFB Fig. 2.13: Detailed structural map (3) Fig. 2.14: Field photographs LSB Fig. 2.15: Petrographic images LSB Fig. 2.16: Stratigraphic columns Fig. 2.17: Synoptic timeline and interpretative synthesis of major structural events Fig. 2.18: Cartoon illustration of major structural events Chapter 3 Fig. 3.1: EPT regional map Fig. 3.2: Study area and sample locations Fig. 3.3: Field photographs MGC Fig. 3.4: Petrographic images MGC Fig. 3.5: Harker variations Fig. 3.6: Discrimination plots Fig. 3.7: Chondrite normalized REE Fig. 3.8: Chondrite normalized multi-element spidergram Fig. 3.9: U-Pb Concordia plots Fig. 3.10: Detrital zircon Concordia plots Fig. 3.11: Pseudoscection Fig. 3.12: Geochemical model source Fig. 3.13: TTG field plots Fig. 3.14: Geochemical model fractionation Fig. 3.15: Geochemical model anatexis Fig. 3.16: Histogram illustration Chapter 4 Fig. 4.1: EPT map and Pilbara Supergroup stratigraphy Fig. 4.2: Study area and sample locations Fig. 4.3: Field photographs Fig. 4.4: Petrograpic images LSB Fig. 4.5: Discrimination diagrams Fig. 4.6: Harker variations Fig. 4.7: Chondrite normalized REE Fig. 4.8: Primitive mantle normalized trace elements Fig. 4.9: Petrograpic images CFB and SMSZ Fig. 4.10: Discrimination diagrams Fig. 4.11: Harker variations Fig. 4.12: Chondrite normalized REE Fig. 4.13: Primitive mantle normalized trace elements Fig. 4.14: Zr (ppm) versus element Fig. 4.15: Discrimination plot Fig. 4.16: Volcanic cycles Fig. 4.17: Source, fractionation, contamination plots

List of Tables Chapter 3 Table 3.1: Sample list with mineral assemblages, 59 Table 3.2: XRF major element data, 65 Table 3.3: ICP-MS trace element data, 69 Table 3.4: U-Pb zircon data, 74-75 Chapter 4 Table 4.1: Sample list and mineralogy Table 4.2: XRF data LSB Table 4.3: LA-ICP-MS data LSB

xii Table 4.4: XRF data CFB and SMSZ Table 4.5: LA-ICP-MS data CFB and SMSZ

xiii “Whatever you do will be insignificant, But it is very important that you do it.” (Mahatma Gandhi, 1869-1948)

Chapter 1 ______

INTRODUCTION

1.1 Research Background

The internal heat of terrestrial planetary bodies is dissipated into the surrounding space by volcanic activity and radiation and thermal diffusion from surface rocks (e.g. Lay et al., 2008). To a large extent, the mantle geodynamics are responsible for the transport of heat from the planet’s core to the surface by thermal convection (e.g. Breuer and Moore, 2007). As a result of the convective motion of the mantle, the planet’s surface is being deformed. Magmatic activity and its spatial distribution, as well as deformation of the surface, are manifestations of the internal heat flow of terrestrial planets (e.g. Breuer and Moore, 2007). Present-day Earth, from a thermomechanical point of view, is a three-layer system formed by the core, mantle and lithosphere. Large contributions of the Earth’s total heat flow (~46 TW) are generated by conductive heating of the lowermost mantle by the hot core, mantle cooling, and heat from the decay of radioactive elements (e.g. Bercovici et al., 2000; Lay et al., 2008). From the core-mantle boundary, situated at approx. 2900 km depth, hot material rises buoyantly in cylindrical mantle plume upwellings creating large convective cells (Bercovici et al., 1989; Tackley, 2000; Lay et al., 2008). The rigid lithosphere responds to the mantle convection and is integrated into its motion (e.g. Tackley, 2000). Along sheet-like planar downwellings of the mantle convective system the lithosphere is being recycled into the mantle forming subduction- zone systems between converging lithospheric plates. The elastoplastic rheology of subducting lithosphere allows the internal build-up of large tectonic stresses that tear the essentially rigid plates apart far away from the subduction zones (e.g. Bercovici et al., 1989). The resulting break- up of the lithosphere (e.g. East African rift system, mid-ocean ridges) causes a passive upwelling of the creeping fluid-like asthenosphere that produces new oceanic crust by decompression melting of the upper mantle (e.g. Key et al., 2013). Above subduction zones new crust is formed when down-going oceanic crust de-hydrates to release fluids that infiltrate the overlying mantle wedge resulting in partial melting (McCulloch and Gamble, 1991; Zellmer et al., 2015). In oceanic-oceanic and oceanic-continental plate converging subduction zones the newly generated magma forms typical island- (e.g. Japan) or continental- volcanic arcs (e.g. Andes), respectively.

Chapter 1 Introduction ______

Where two continental plates collide the crust is being over-thickened by thrusting and experiences intensive deformation that characterizes present-day mountain ranges (e.g. Himalayas). In the latter case, new crust is produced following the deep burial of the lower crust resulting in partial melting through heating and de-hydration. This unique mobile-lid plate tectonic geodynamic system stands in contrast to the common mode of heat transfer in terrestrial planets, characterized by mantle convection beneath a stagnant- lid upper thermal boundary layer that develops because of large temperature (and temperature- dependent viscosity) differences between mantle and surface (Solomatov and Moresi, 1997; Tackley, 2000; Bercovici et al., 2000; Breuer and Moore, 2007; Korenaga, 2009; O’Rourke and Korenaga, 2012). At the surface, radiation and thermal diffusion dissipate heat into space. In addition, large quantities of heat are released through mantle-derived volcanism that is directly responsible for the generation of new crust (e.g. Breuer and Moore, 2007). The formation of new crust in a stagnant-lid regime can occur either locally in the form of hot mantle upwellings, such as plumes or heat pipes, or globally through whole-crustal re-surfacing if the upper mantle temperatures are higher than the solidus (e.g. Kamber et al., 2005; Breuer and Moore, 2007; Moore and Webb, 2013).

How was it possible for Earth to develop its present complex geodynamic mechanism of heat dissipation? How did the early crust form that eventually broke-up into rigid plates subject to recycling in subduction zones?

It is generally accepted that the oldest physical evidence for the operation of plate tectonic (i.e. mobile-lid) processes, preserved in the rock record, reaches back into the Mesoarchaean (~3200 Ma; e.g. Cawood et al., 2006; Van Kranendonk et al., 2010; Hickman and Van Kranendonk, 2012). However, relatively voluminous continental crust had already formed, at least since the early Archaean (3850 Ma), and rare evidence for continental material (i.e. up to ~4374 Ma detrital zircon) dates back into the Hadean (>3850 Ma; e.g. Harrison, 2009; Valley et al., 2014; Kamber, 2015). The early Archaean 3850-3200 Ma crustal remnants on Earth comprise a distinct crustal architecture, known as granite dome – greenstone keel terranes, which is hard to reconcile with formation through modern-style horizontal plate tectonic processes (e.g. Van Kranendonk et al., 2007). Instead, it is proposed that these dome-and-keel terranes represent the manifestations of a gravity-driven geodynamic mode in a hot stagnant-lid early Earth (e.g. Debaille et al., 2013), in which crust was formed through “heat-pipes” (Moore and Webb, 2013), or massive mantle plume- derived volcanic resurfacing events (e.g. Kamber et al., 2005). The preserved deformational style and composition of early Archaean rocks (3850-3200 Ma) is unique and has no equivalent in post-Archaean rock associations. Furthermore, the early Archaean mantle must have exhibited its peak temperatures during the thermal evolution of the Earth (e.g. Korenaga, 2006). The prominent dome-and-keel terranes therefore present the exceptional opportunity to study a distinct early Earth tectonic regime, representing an early or intermediate

2 Chapter 1 Introduction ______

Fig. 1.1: Predicted thermal history of the Earth; potential mantle temperatures (Tp) estimates are based on a constant Urey ratio model using the present value of Ur (0) = 0.34 (from Korenaga, 2013); solid circles represent petrological estimates by Herzberg et al. (2010); Note that peak temperatures were reached around 3500 Ma during the early Archaean.

stage of planetary evolution, before secular cooling provided thermal and mechanical conditions of the mantle-lithosphere system that allowed the subsequent development of a plate-tectonic regime. Most of the dome-and-keel terranes experienced intensive late and post-Archaean deformation and metamorphic overprinting and reworking, obscuring their primary geological features. This modification of primary structures and magmatic protoliths hampers the investigation of the petrogenesis and deformation of the lithological inventory directly associated with dome-and-keel formation. In this thesis I focus on part of the 3530-3225 Ma East Pilbara Terrane of Western Australia, which forms one of the oldest and best-preserved dome-and-keel terranes. I critically examine the local tectonic evolution of the western Doolena Gap greenstone belt in the light of previously established models of regional dome-and-keel formation in the East Pilbara Terrane. I address some critical remaining controversies and knowledge gaps in the tectono-magmatic evolution of the East Pilbara Terrane, outlined below, to provide an improved understanding of dome-and-keel tectonics in the study area. In the following sections, a brief review of the overall background of Earth’s thermal history and Archaean geodynamics is provided, followed by an overview of the structural and lithological make-up of dome-and-keel terranes, and currently proposed models of their formation.

3 Chapter 1 Introduction ______

Fig. 1.2: see next page for caption.

4 Chapter 1 Introduction ______

Fig. 1.2: Results from numerical simulations of mixed internal and basal heating for complex 2D (a) and 3D spherical (b) domains (from Weller and Lenardic, 2012); the Heating Ratio (H) is defined as the ratio of the Rayleigh number of purely internally heated convection to the Rayleigh number for purely basally heated convection; grey shells, high viscosity “plates”; yellow bands, regions of upwelling material; thermal profiles from the core-mantle boundary are colour-coded with decreasing temperature from red to blue; Note that the steady-state model parameters are the same, but different tectonic regimes are simulated depending on the evolution history in terms of increasing versus decreasing lid yield strength.

1.1.1 Earth’s thermal evolution, and the theoretical consequences for Archaean geodynamics

Earth’s present-day global heat flux is estimated to be approximately 40 TW (Korenaga, 2008), while internal heat production by radioactive elements, predominantly long-lived isotopes that are characterized by high decay energies and/or are present in sufficient abundance within the Earth’s mantle (238U, 235U, 232Th, 40K), is estimated to be ~20 TW (Korenaga, 2008). The difference in surface heat flux and heat production (quantified by the Urey ratio, the quotient of internal heat production and surface heat flux) is explained by secular cooling of the Earth (Korenaga, 2008). Secular cooling is in agreement with theoretical predictions of a hotter early Earth. Initial heating of the embryonic Earth by planetary accretion friction, core formation, and the greater abundance in heat producing elements is postulated to have reached maximum internal mantle temperatures in the Archaean (Korenaga, 2006; 2013; Herzberg et al., 2010; Fig. 1.1), with an estimated heat flux up to three times higher than today (e.g. Sleep, 2010). It is argued that the thermal conditions of the early Earth resulted in thermo-physical properties of the lithosphere (e.g. thick, weak and buoyant) that did not allow the initiation and/or sustainability of modern-style plate tectonics (Hamilton, 1998; van Thienen et al., 2004a; van Hunen and van den Berg, 2008). The higher mantle temperatures would have promoted higher degrees of partial melting within deeper parts of the mantle resulting in the formation of hot and thick oceanic lithosphere that would have been too weak to break into plates, and too buoyant to subduct (e.g. van Thienen et al., 2004a). Although numerical models have demonstrated the possibility of subduction initiation, it is proposed that break-off of the weak lithospheric slab would have prevented subduction to be long-lived (van Hunen and van den Berg, 2008; Moyen and van Hunen, 2012). Some authors therefore argue for short-lived episodic subduction events (Moyen and van Hunen, 2012). Furthermore, horizontal plate tectonic processes as a mode of mantle convection might not have been able to efficiently dissipate Earth’s internal heat to cool the hotter Archaean planet (Vlaar et al., 1994; van Thienen et al., 2004b). Gravity-driven tectonic models in a stagnant-lid geodynamic regime are proposed for the early Earth (Debaille et al., 2013). Episodic outpourings of massive mantle plumes could have led to crustal re-surfacing causing large amounts of heat dissipation into space (Kamber et al., 2005; Moore and Webb, 2013). A stagnant-lid geodynamic model has also been demonstrated for the early thermal history of Earth-like systems by thermodynamic modeling (Weller and Lenardic, 2012; Griffin et al., 2014). Although high basal heat in a mantle convective system produces great buoyancy favoring mobilization of its upper thermal boundary layer (mobile-lid), the process is muted by the internal heat production within the mantle (increased radioactive decay), causing

5 Chapter 1 Introduction ______stagnation of the ‘lid’. Because of the ineffective conductive heat loss through an immobile lid, the internal heat builds up until it is released in episodes of massive mantle overturn and crust recycling (Griffin et al., 2014). Importantly, based on numerical simulations it has been shown that the tectonic regime of a planet strongly depends on its evolutionary history (e.g. Weller and Lenardic, 2012; O’Neill et al., 2016). For example, surface expressions of mantle convection are controlled by the directionality of the lid yield stress evolution (Weller and Lenardic, 2012). Figure 1.2 demonstrates that under the same steady-state planetary model parameter values, multiple solutions (i.e. either mobile- or stagnant-lid solutions) are possible. The observed state depends on prior (initial) conditions, and whether the yield strength increases or decreases (Fig. 1.2). Therefore, predictions of convective regimes are difficult in cases where no geological information is available (Weller and Lenardic, 2012). In other words, to conclusively determine the geodynamic state of a planet at a particular time, geological constraints are key.

1.1.2 Archaean dome-and-keel terranes

Dome-and-keel terranes often represent the most ancient nuclei of Archaean Cratons. The crustal structures comprise high-wavelength (~60 km) mid- to upper-crustal granitic domes, surrounded by narrow (~5-10 km wide) synclinal supracrustal greenstone belts. The domal antiforms are characterized by ovoid outlines in map view, but complex internal structures comprising multiple granitic components (e.g. Van Kranendonk et al., 2007). The domes often feature radial faults and ring faults at and around their margins (Van Kranendonk et al., 2007). As evident for example in the East Pilbara Terrane and the Western Dharwar Craton, older granitic components are found in the marginal areas of domes, while younger, re-worked components intrude the centers (e.g. Van Kranendonk et al., 2007; Rao et al., 1991). The older granitic components mainly comprise tonalitic-trondhjemitic to granodioritic gneisses (TTG in the following), while the volume of potassic granites increases with decreasing age (e.g. Smithies et al., 2007). The volcano-sedimentary greenstone keels form tight synclines wrapping around the domal complexes. The contact between the granitic domes and the greenstone keels is either intrusive, or tectonic (e.g. Van Kranendonk et al., 2007). Where preservation is good, consistent outward younging stratigraphy, away from the domes, is observed within the keels (Van Kranendonk et al., 2007). The synclinal greenstone limbs are often steeply tilted or overturned, parallel to the inferred margins of adjacent domes. Within the highly deformed central parts of the keel synclines, L- tectonites are developed (e.g. Collins et al., 1998; Chardon et al., 1998; Parmenter et al., 2006). The greenstone successions are dominated by tholeiitic pillow basalts, but varying amounts of felsic volcanic rocks, sediments, and ultra-mafic komatiites are often present. In the case of the East Pilbara Terrane, the greenstone successions, known as the 3530-3225 Ma Pilbara Supergroup, comprise numerous ultra-mafic to mafic to felsic volcanic cycles, separated by unconformities (e.g. Hickman and Van Kranendonk, 2012). The Pilbara Supergroup is interpreted as an autochthonous succession with a thickness of 20 km. However, a maximum 12 km is preserved in

6 Chapter 1 Introduction ______single greenstone belts (e.g. Hickman, 2012). Emplacement ages of East Pilbara Terrane TTG components within the domes coincide with major felsic volcanic formations at the top of the volcanic cycles. Many of the dome-and-keel terranes, such as the East Pilbara Terrane, the Barberton area (South Africa) and the Western Dharwar Craton (South India) comprise evidence for precursor continental basement, inferred from i) detrital zircon within younger sediments (e.g. Kemp et al., 2015), ii) geochemical and isotopic signatures within both dome and keel rocks indicating crustal reworking and/or assimilation with pre-existing crust (e.g. Green et al., 2000; Kröner et al., 2013; Van Kranendonk et al., 2015), and iii) rare findings of older xenoliths (e.g. McNaughton et al., 1988). The nature of this precursor basement remains largely unclear.

1.1.2.1 Model of dome-and keel formation in the East Pilbara Terrane

Dome-and-keel formation is explained by partial convective overturn of a gravitationally unstable mid- to upper crust (e.g. Collins et al., 1998). It is proposed that the thick autochthonous layering of mantle plume-derived dense mafic volcanic material on top of a partly continental, soft and buoyant felsic substrate caused a crustal-scale re-organization, driven by Rayleigh-Taylor flow (e.g. McGregor, 1951; West and Mareschal, 1979; Mareschal and West, 1980; Anhaeusser, 1984; Bouhallier et al., 1995; Chardon et al., 1996; 1998; Collins et al., 1998; Smithies et al., 2007; 2009; Van Kranendonk et al., 2007; 2015). Partial melting of the felsic mid-crust caused the buoyant rise of felsic magma domes, and the sinking (i.e. sagduction) of dense supracrustal rocks into the mechanically weakened mid-crust (e.g. Collins et al., Sandiford et al., 2004). The dominant mechanism of partial melting of the felsic mid-crust remains unclear, and is interpreted as the result of either i) heat from radioactive decay in the U-, Th-, K-enriched granitic rocks (e.g. Sandiford et al., 2004), ii) plume heat from below (e.g. Van Kranendonk et al., 2015), iii) burial of colder granitic crust into the warm mid-crust (Collins et al., 1998), or iv) a combination of these processes (e.g. Rey et al., 2003; Van Kranendonk et al., 2015). Regional partial convective overturn (i.e. dome-and-keel formation) in the East Pilbara Terrane occurred around 3310 Ma (e.g. Sandiford et al., 2004).

1.2 Research Problems and Questions

In a comprehensive review of the available literature on granite dome – greenstone keel formation, I identified three major controversies:

I) The geometry, kinematics, and tectonic significance of an intense deformation phase predating the regional dome-and-keel event in the East Pilbara Terrane (EPT)

The relative timing, geometry, and kinematics of deformation events in dome-and-keel terranes constitute critical evidence for their tectonic evolution and thus the ongoing debate of gravity- driven versus horizontal tectonics. Obtaining these data is challenging because even the best-

7 Chapter 1 Introduction ______

preserved dome-and-keel terranes, such as the East Pilbara Terrane, were affected by post- dome-and-keel development plate tectonic-related deformational modification (e.g. Van Kranendonk et al., 2002; Hickman and Van Kranendonk, 2012). In addition, significant horizontal stresses and resulting cylindrical, linear upper-crustal deformation features have been shown to accompany gravity-driven tectonic processes, such as crustal-scale partial convective overturn or Rayleigh-Taylor instabilities of the mantle lithosphere (e.g. Mareschal and West, 1980; Lin, 2005; Parmenter et al., 2006; Van Kranendonk et al., 2007; Erickson, 2010, Pysklywec and Cruden 2004). This has led to much debate in the past. The proposed gravity- driven structural development of dome-and-keel terranes has been constantly challenged by uniformitarian, horizontal plate tectonic models (e.g. van Haaften and White, 1998; Kloppenburg et al., 2001; Blewett, 2002). It is critical to clearly identify the relative timing, geometry, kinematics and mechanisms of successive deformation events during Archaean crustal evolution, in order to conclusively demonstrate the nature of predominant tectonism (e.g. Burg et al., 2004). Specifically, early tight folds below a regional 3427 Ma unconformity have been observed adjacent to the southeastern margin of the Carlindi Granitic Complex in the East Pilbara Terrane (Buick et al., 1995). The development of these high-strain folds remains unexplained. Importantly, their relative timing indicates major deformation prior to the generally accepted dome-and-keel event at 3310 Ma in the East Pilbara Terrane (e.g. Sandiford et al., 2004). Therefore, interesting questions arise: Can this early high-strain event be seen elsewhere in the East Pilbara Terrane? If so, what are the geometric, kinematic and metamorphic properties of this event? How does it relate to the formation of the dome-and-keel terrane, if at all? Is there perhaps a first main doming event in the research area that is substantially older than the regional age? Here, I examine the Doolena Gap greenstone belt, of the East Pilbara Terrane, which comprises some of the oldest and most highly deformed greenstone rocks in the East Pilbara Terrane (e.g. van Kranendonk, 2010). I show that a pre-regional-doming deformation event is clearly present and widespread, analyse it carefully, and interpret it in the context of dome-and-keel formation.

II) The timing, nature and formation of pre-dome-and-keel granitic crust

A major knowledge gap in the evolution of dome-and-keel terranes is the formation of the most ancient granitic mid-crust that is thought to have represented a basement onto which greenstones were deposited (e.g. Smithies et al., 2009; Hickman and van Kranendonk, 2012; Kröner et al., 2013). Both field-based and numerical models of partial convective overturn are successful in that they demonstrate that a gravitationally instable crustal configuration arising from dense greenstones overlying buoyant felsic substrate results in mid- to upper crustal Rayleigh-Taylor flow (e.g. Chardon et al., 1996; Collins et al., 1998; Sandiford et al., 2004; Bodorkos and Sandiford, 2006; Thébaud and Rey, 2013). However, the models do not account for the formation of the mid-crustal layer (Sizova et al., 2015). In the East Pilbara Terrane, a

8 Chapter 1 Introduction ______

pre-existing continental basement is inferred from the detrital zircon record and isotope data. However, its nature and formation remains unclear. How did the early granitic melts form? The lack of spatiotemporal and chemical constraints on the formation of the ancient granitic basement hampers our understanding of the tectono-magmatic processes and timeframes of early Archaean crustal growth that ultimately led to partial convective overturn. In this thesis I provide new geochronological, petrological and geochemical data from ancient rocks of the Muccan Granitic Complex to discuss the generation of early granitic melts in the East Pilbara Terrane.

III) Lateral variability in the architecture, thickness and composition of the greenstones of the East Pilbara Terrane

There are two problems regarding the formation of supracrustal greenstone successions that are critical to reconstruct the tectonic evolution of dome-and-keel terranes: i) Most forward models of partial convective overturn make use of the reported maximum thickness of supracrustal greenstone successions (12-20 km) and assume that dense ultramafic to mafic rocks constitute their bulk composition. This parameter combination creates the expected gravitationally unstable layer of volcanic rock (e.g. Sandiford et al., 2004; Moore and Webb, 2013). However, greenstones generally exhibit significant spatiotemporal variations of thickness and compositions laterally (Hickman, 1984, 2008, 2011; Van Kranendonk, 2010). This fact is often disregarded in partial convective overturn simulations (e.g. Sandiford et al., 2004; Thébaud and Rey, 2013; O’Neill and Debaille, 2014). Greenstone stratigraphy is often interrupted by numerous, sometimes angular, unconformities. Single lithostratigraphic formations may vary in thickness from a few meters to hundreds of meters between greenstone belts of the same craton, or may even be absent in other belts (e.g. Hickman, 2008; Van Kranendonk, 2010). Although mafic rocks often dominate, felsic volcanic and sedimentary formations are abundant in some greenstone belts, while truly ultra-mafic rocks are subordinate (e.g. Furnes et al., 2013; Condie et al., 2016). Besides the conclusively demonstrated autochthonous nature of the greenstone stratigraphy in the East Pilbara Terrane, there are drastic variations in thickness and composition of single greenstone formations (e.g. Hickman, 2008). This lateral variability requires explanation and might reveal local differences in the tectonic evolution. In the studied Doolena Gap greenstone belt, for example, a sandstone ascribed to the regional Strelley Pool Formation reaches exceptional thickness of 1 km, while in most other greenstone belts the Strelley Pool Formation reaches thickness of only ~25 m and formed by stromatolite, evaporite, chert and carbonaceous sediments (e.g. Lowe, 1983; Hickman, 2008; Van Kranendonk, 2010). Here, I attempt to reconcile and interpret this dramatic local variation of depositional environment within a greenstone belt with the regional tectonic model. This may provide new detail in the understanding of the dynamics of early Archaean

9 Chapter 1 Introduction ______

upper crustal growth in a dome-and-keel terrane and help to explain apparent inconsistencies of the local rock record with regional models. ii) The petrogenesis of felsic and mafic greenstone volcanic rocks is debated. Classically, mafic-felsic volcanic suites in Archaean greenstone belts have been interpreted as bimodal suites, in which felsic volcanics are derived from crustal melting (tonalite-trondhjemite- granodiorite/TTG derivatives), and mafic volcanics are derived from mantle melts (e.g. Kamber, 2015). However, in the East Pilbara Terrane most of the felsic volcanic rocks form the tops of volcanic cycles, and are interpreted as products of tholeiitic differentiation (e.g. Smithies et al., 2007). Furthermore, although the greenstone volcanic rocks are mostly interpreted as formed by hot mantle plumes, ultra-mafic komatiites are relatively rare or absent, and the greenstones are dominated by low-MgO tholeiitic basalts (e.g. Smithies et al., 2007). The generation of compositional types within the basalts has been interpreted as the result of either the contribution from different mantle sources (e.g. Smithies et al., 2007), varying amount of crustal contamination (e.g. Green et al., 2000), or fractional crystallization processes in crustal magma chambers (e.g. Van Kranendonk et al., 2015). To address this problem, I am testing these contrasting petrogenetic models of mafic-felsic bimodal volcanism in the Doolena Gap greenstone belt. By doing so, I provide insights into the upper crustal growth and overall magmatic structure of the early Archaean crust.

1.3 Aim of the Research

In this thesis, I reconstruct the primary depositional, magmatic, and structural history of a granite dome – greenstone keel transect in the early Archaean Doolena Gap greenstone belt. This allows me to:

[1] Investigate the lithostratigraphy of the greenstone belt with the purpose to explain local variations in the depositional environment. Moreover, I provide relative and absolute age constraints to test the proposed stratigraphic correlation of greenstone formations and therefore better constrain the relative timing of magmatic and structural events.

[2] Examine the formation of and the petrogenetic relationships between greenstone volcanic rocks and dome components, and between the mafic-felsic bimodal volcanic suite. In this context, I present new geochronological constraints for the timing of emplacement of granitic dome components to investigate the temporal relationship between greenstone belt and dome, and potentially identify pre-East Pilbara Terrane basement components.

[3] Unravel the relative timing, geometry, and kinematics of deformation structures related to dome-and-keel development. This is necessary to reveal the history of pre- and syn-deformational sedimentary and magmatic processes and to tie in the local evolution of the greenstone belt with regional tectonic models.

10 Chapter 1 Introduction ______

1.4 Research Philosophy and Methodology

A combination of methodological approaches is used to address each of the main research problems. It is fundamental to elucidate the structural history based on field relationships and kinematics of structural features. This requires a spatial understanding of the distribution of rock types that I achieved through field-based lithostratigraphic investigations. Observed structural features were correlated with the previously established craton-wide features based on the literature and satellite images. The greenstone belt and map-scale features were further evaluated in petrographic thin sections. Detailed mapping and structural analyses resulted in a comprehensive understanding of the structural-lithostratigraphic anatomy of the studied dome-keel transect. Based on the established structural-lithostratigrapic framework, most suitable and freshest samples were collected for petrological and geochemical investigations. The selection of the most suitable (fresh) samples for laboratory analyses was based on petrographic characterization. Petrographic observations provided a petrological background for bulk-rock geochemical interpretations to investigate petrogenetic processes and relationships between the various rock types. The identification of pre-East Pilbara Terrane basement material has been addressed by U- Pb (LA-ICP-MS) zircon geochronology of granitic gneisses from the dome and detrital zircon provenance studies of a sandstone formation from within the greenstone belt.

1.5 Outline of Chapters

The first chapter provides the background of the thesis, and outlines research problems, aims of the present contribution, and the research philosophy. Chapter 2 constitutes the basis of my enquiry: a mostly field-based study of the western Doolena Gap greenstone belt, presenting the structural and lithostratigraphic anatomy of the studied dome-and-keel transect. A model for the lithostratigraphic-structural evolution is presented and discussed in context with the established regional model for East Pilbara Terane dome-and-keel formation. In Chapter 3, U-Pb zircon data from gneissic components of the dome and provenance ages from the supposed Strelley Pool Formation sandstone are presented, as well as petrological and geochemical data of host gneisses and associated leucosomes from the dome. The chapter provides evidence for pre-East Pilbara Terrane basement and discusses its formation. The conclusive stratigraphic correlation of the sandstone with the regional Strelley Pool Formation, provide critical time constraints on the previously established structural history. In Chapter 4, geochemical and petrological data from a suite of ultramafic-mafic to more evolved greenstone rocks are presented. The data confirms the previously proposed petrogenetic relationship of mafic and more evolved rocks within the greenstone suite. The petrogenesis of the rocks is discussed in the light of the contrasting models of crustal magmatic processes versus different mantle sources. Chapter 5 represents a brief synthesis and discussion and highlights the contributions of the thesis. The currently proposed model of the Doolena Gap greenstone belt tectonic and magmatic evolution is presented.

11 Chapter 1 Introduction ______

“You don’t blame us for being here, do you? After all, we have no place to go. No home… Incidentally, what an excellent day for an exorcism...” (William P. Blatty, The Exorcist, 1971)

Chapter 2 ______

LITHOSTRATIGRAPHY AND STRUCTURE OF THE EARLY ARCHAEAN DOOLENA GAP GREENSTONE BELT, EAST PILBARA TERRANE, WESTERN AUSTRALIA

Wiemer, D., Schrank, C. E., Murphy, D. T., and Hickman, A. H., 2016: Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt, East Pilbara Terrane, Western Australia. Precambrian Research, 282, 121-138 Copyright © 2016 Elsevier

2.1 Abstract

The early Archaean East Pilbara Terrane of the Pilbara Craton represents the archetypical granite dome – greenstone keel terrain and preserves much of the primary geological features of dome- and-keel formation. Here I present the first detailed lithostratigraphic and structural transect from marginal gneisses of the Muccan Granitic Complex through 6 km of steeply dipping strata of the western Doolena Gap greenstone belt.

Based on overprinting relationships, I identified five deformational events (D1-D5) that can be correlated between four structural domains: i) marginal orthogneisses of the polyphase Muccan Granitic Complex, ii) a dominantly mylonitic shear zone (South Muccan Shear Zone) representing a granite dome – greenstone keel detachment zone, iii) a Central Fold Belt of dominantly mafic greenschists, and iv) a lower greenschist- to sub-greenschist facies southern Low-Strain Belt.

The D1 event in the Low-Strain Belt is recorded in syn-depositional normal faulting (f1) during eruption of deep-water low vesicular pillow basalts of the ca. 3470 Ma Mount Ada Basalt, followed by a conformable transition towards shallow water stromatolites in the Duffer Formation.

I interpret the D1 event as the result of upper-crustal extension and subsequent uplift above the Muccan Granitic Complex.

The D2 event shows the most intense deformation with syn-anatectic tight folds within the

Muccan Granitic Complex and tight to isoclinal transposed folds (F2) and fluid-induced normal faults (f2) within the Central Fold Belt. The D2 event resulted in a regional EW composite S1/2 foliation, and in the development of the mylonitic South Muccan Shear Zone along the dome-keel

13 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______interface and within the Central Fold Belt, mineral lineations (L2) in conjunction with increasing southward D2 L-tectonite. This D2 event is not observed in a ~1 km-thick sandstone package (ascribed to the <3427 Ma Strelley Pool Formation) in the Low-Strain Belt that was deposited over an angular unconformity indicating 45º SSW tilting of the underlying Mount Ada Basalt and

Duffer Formation. This constrains the D2 event to between 3470-3427 Ma.

I propose that the D1 and D2 events record locally significant initiation of Muccan Granitic Complex doming and associated sagduction of greenstone keel rocks, suggesting that partial convective overturn initiated early in the evolution of the East Pilbara Terrane.

Keywords: East Pilbara Terrane, dome-and-keel, Archaean tectonics, Doolena Gap greenstone belt, Muccan Granitic Complex;

14 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

2.2 Introduction

The tectonic evolution of Archaean granite-greenstone terranes remains a focus of frontier research (e.g. Bédard et al., 2013; Piper, 2013; Anhaeusser, 2014; O’Neill and Debaille, 2014; Furnes et al., 2014). Vertical geodynamic models are proposed to explain the development of Archaean dome-and-keel cratonic nuclei, such as the East Pilbara Terrane (Western Australia), the Barberton area (Kaapvaal Craton, South Africa), the Western Dharwar Craton (South India), and part of the Superior Craton (Canada) (e.g. Anhaeusser, 1984; Hickman, 1984; Chardon et al., 1996, 1998; Collins et al., 1998; Bédard et al., 2003; Hickman and Van Kranendonk, 2004; Lin, 2005; Van Kranendonk, 2011; Van Kranendonk et al., 2014). It is suggested that the nascent lithosphere was too thick, weak and buoyant to allow the initiation and sustainability of modern- style plate tectonics (e.g. van Thienen et al., 2004; van Hunen and van den Berg, 2008). The increased heat-flow would have promoted higher degrees of partial melt extraction from the mantle to form thick mafic proto-crust (e.g. Smithies et al., 2005). Through intra-crustal partial anatexis of the overthickened mafic crust, differentiated crustal layers evolved (Smithies et al., 2009; Johnson et al., 2014). Ongoing voluminous mafic plume-volcanism onto the differentiated basement resulted in gravitational (i.e. Rayleigh-Taylor) instabilities, which led to partial convective overturn and development of dome-and-keel crustal geometry (e.g. Collins et al., 1998). As O’Neill and Debaille (2014) remarked, a crucial problem with recent numerical models in favor of gravity-driven Archaean lithosphere tectonics is embroiled in the initial conditions and associated parameterizations of the ancient upper mantle and crust (see also Weller and Lenardic, 2012). Most of the models make use of 15-20 km greenstone thickness dominated by mafic material – as reported from some well-preserved Archaean crustal cross-sections (e.g. Barberton, East Pilbara Terrane; e.g. Van Kranendonk et al., 2006; Stiegler et al., 2011). However, the presence of significant spatial and temporal variations in thickness and greenstone composition across the lateral extents of ancient greenstone depositional regimes (e.g. Hickman, 1984; 2008; 2011; Van Kranendonk, 2010) often remains disregarded. Local deformational overprinting relationships and associated unconformities within the greenstone pile attest for more complex Archaean lithosphere dynamics (e.g. Moore and Webb, 2013) that may be characterized by differential vertical tectonic movement during ongoing bulk crustal construction over a protracted time period (e.g. Hickman, 2011). In the present study I examine part of the East Pilbara Terrane, which represents the type area for dome-and-keel geometry and has been one of the key areas for the development of gravity- driven tectonic models. Previous lithostratigraphic and structural studies of the East Pilbara Terrane have led to some dispute regarding the causal and temporal relationship between local depositional processes, stratigraphic continuity, and proposed regional tectonic models (e.g. Collins et al., 1989; Nijman et al., 1999; Van Kranendonk et al., 2002, 2004; de Vries et al., 2006). Correlations between adjacent greenstone belts demonstrate the existence of significant compositional and thickness variations across the East Pilbara Terrane that may suggest

15 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.1: a) Simplified geological map of the East Pilbara Terrane (modified after Van Kranendonk et al., 2006); NP, North Pole; WA, Warrawagine; YG, Yilgalong; rectangle indicates outline of Figure 3; b) Location of the East Pilbara Terrane within Australia.

spatiotemporal differences in the local tectonic evolutions of different greenstone belts (Hickman, 2011). To assess the importance of such lateral variation to East Pilbara Terrane evolution it is vital that the lithostratigraphic and structural anatomy of greenstone belts is investigated and related to models for regional dome-and-keel development. Here, I re-examine the lithostratigraphy and structure of the thickest, most cylindrical part of the early Archaean Doolena Gap greenstone belt to [1] decipher the initial configuration of the upper crust in this area, [2] reconstruct the subsequent structural evolution with a focus on the identification of dome-and-keel formation related deformation and deformation mechanisms, and [3] discuss how the local observations can be accommodated within the regional East Pilbara Terrane dome-and-keel evolution.

2.3 Regional Geology and Tectonic Framework of the East Pilbara Terrane

The East Pilbara Terrane forms the ancient nucleus of the Pilbara Craton and comprises numerous antiformal granitic domes (~60 km diameter) that display complex internal structures, surrounded by synclinal keels of curvilinear greenstone belts (Van Kranendonk et al., 2004; 2006; Fig. 2.1).

16 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.2: Generalized stratigraphy of the Pilbara Supergroup (Pilbara Supergroup; modified from Hickman, 2012; timing of granitic supersuites from Van Kranendonk et al., 2006).

Based on field and geochronological data, four granitic supersuites have been identified within the domes (e.g. Van Kranendonk et al., 2006): the 3490-3460 Ma Callina, the 3450-3420 Ma Tambina, the 3324-3290 Ma Emu Pool, and the 3275-3225 Ma Cleland Supersuites (Fig. 2.1). At the margins of the granitic domes, either sheared- or intrusive contacts with adjacent steeply dipping to overturned greenstone belts are observed (Van Kranendonk et al., 2006). The greenstone belts comprise volcano-sedimentary successions of the early Archaean Pilbara Supergroup. Based on the recognition of two regional erosional unconformities at ca. 3427 Ma and ca. 3270 Ma (beneath the Strelley Pool and the Leilira Formations, respectively), the Pilbara Supergroup has been divided into the lowermost Warrawoona Group, the Kelly Group, and the uppermost Sulphur Springs Group (Hickman, 2012; Fig. 2.2). The Warrawoona Group has been

17 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.3: Geological map of the Doolena Gap – Marble Bar greenstone synclinorium (based on and modified after geological maps of the Geological Survey of Western Australia: Williams, 1999; Van Kranendonk, 2010; Hickman, 2010). Geochronological data (red dots; ages in Ma) by (1) Thorpe (1991), written comm. in Williams (1999), conventional U-Pb zircon; (2) McNaughton et al. (1993), SHRIMP U- Pb zircon; (3) Nelson (1999), SHRIMP U-Pb zircon; (4) Beintema (2003), Ar-Ar hornblende; (5) deVries et al. (2006), SHRIMP U-Pb zircon.

subdivided further into the Coonterunah, the Talga Talga, the Coongan, and the Salgash Subgroups (Fig. 2.2). The entire Pilbara Supergroup exhibits a cumulative thickness of over 20 km. However, a maximum of 12 km thickness is exposed within single greenstone belts (Van Kranendonk et al., 2006). The greenstone successions comprise dominantly tholeiitic basalts, subordinate ultramafic volcanic and intrusive rocks, but also felsic volcanic formations, and sediments (e.g. Hickman, 2012). At least eight ultramafic-mafic-felsic volcanic cycles, separated by unconformities or interlayered sedimentary rocks, have been observed in the Pilbara Supergroup (Hickman, 2012; Fig. 2.2) that are interpreted as the result of episodic mantle plume activity (Smithies et al., 2007). Emplacement ages of the four granitic supersuites coincide with deposition of thick felsic volcanic formations at the top of major volcanic cycles (Van Kranendonk et al., 2007; Fig. 2.2). The East Pilbara Terrane is partly overlain by the 3020-3010 Ma Gorge Creek Group (De Grey Superbasin), and the 2780-2630 Ma Fortescue Group (Fortescue Basin), across regional unconformities (Fig. 2.1; e.g. Hickman et al., 2010).

2.4 Geological Setting of the Study Area

The Doolena Gap greenstone belt forms the northern limb of an EW-trending synclinorium, which together with the northern arm of the Marble Bar greenstone belt, as its fault-bound southern counterpart, is sandwiched between the domal antiforms of the Muccan Granitic Complex to the north and the Mount Edgar Granitic Complex to the south (Fig. 2.1; Fig. 2.3). To the west, the greenstone synclinorium widens and exposes maximum stratigraphic thickness of the lowermost volcano-sedimentary successions. The greenstone successions are overturned with southward- facing younging directions and have been ascribed to the ca. 3470 Ma Mount Ada Basalt and the overlying Duffer Formation of the Warrawoona Group (Coongan Subgroup) and the 3427-3350

18 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.4: Geological-structural map of the studied ‘1 Mile Creek’ granite-greenstone traverse, western Doolena Gap greenstone belt; location indicated in Figure 2.3.

Ma Strelley Pool Formation (Van Kranendonk et al., 2004; 2006; Van Kranendonk, 2010). Stratigraphic affinities are solely based on regional lithological correlation within the East Pilbara Terrane (Thorpe, 1991, written comm. in Williams, 1999; McNaughton et al., 1993; Nelson, 1999; Beintema, 2003; de Vries et al., 2006; Fig. 2.3). The Strelley Pool Formation demarcates the southern extent of the study area and forms an up to 1 km-thick quartz-sandstone package that

19 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.5: Geological cross sections through the studied granite-greenstone traverse; locations of the

profiles A-B and C-D are indicated in the map of Figure 2.4; the profiles are perpendicular to the main S2 foliation and show the sub-parallel alignment of the margin of the Muccan Granitic Complex (MGC) with the foliation and (overturned) bedding in the greenstone belt; part of profile C-D in the Central Fold Belt runs sub-parallel to the local strike orientation, where the area is affected by ductile faulting and

associated drag folding (F4).

20 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______overlies the Warrawoona Group across an angular erosional unconformity (Van Kranendonk, 2010). South of the study area, the Euro Basalt (Salgash Subgroup) overlies the Strelley Pool Formation and is in turn unconformably overlain by the De Grey Superbasin and Fortescue Group (e.g. Hickman et al., 2010; Fig. 2.3). Both the Muccan Granitic Complex and the greenstone successions, up to the Fortescue Group, are cut by NE-SW-trending mafic dykes that have been ascribed to the East Pilbara Terrane-wide 2772 Ma Black Range Dyke swarm (U-Pb SHRIMP, baddeleyite; Wingate, 1999). Intrusion ages from within the Muccan Granitic Complex are derived from regional sampling and reveal contributions of each of the four supersuites between ca. 3470 and 3240 Ma (U-Pb SHRIMP, zircon; Nelson, 1998; Van Kranendonk et al., 2006; Van Kranendonk, 2010).

2.5 Lithostratigraphy and Structure

Detailed geological-structural mapping was conducted along the ~5x3 km ‘1 Mile Creek’ traverse from the south-westerly fringe of the Muccan Granitic Complex into the western Doolena Gap greenstone belt (Fig. 2.3; Fig. 2.4). The studied granite-greenstone traverse displays an overall transition from amphibolite to greenschist and finally sub-greenschist metamorphic facies to the south. Based on the presence of distinct lithological components and the recognition of abrupt changes in deformational style and intensity across key structural features, I subdivided the area, from north to south, into four structural domains: i) the Muccan Granitic Complex, ii) the South Muccan Shear Zone, iii) the Central Fold Belt, and iv) the southern Low-Strain Belt (Fig. 4). Regional foliation and overturned bedding planes dip steeply to sub-vertically to the north, striking parallel to the Muccan Granitic Complex interface (Fig. 2.5). Relative timing of structural events is inferred from overprinting relationships, and their presence and absence, respectively, within the domains. I recognized a total of five deformation phases (D1-D5). In the following sections I will describe the geological-structural inventories of each domain. For descriptive purposes I use systematic numeration for the deformation phases (D1-n) within each of the structural domains first, and later discuss the most likely inter-domain correlations of the structural features.

2.5.1 Muccan Granitic Complex

2.5.1.1 Lithological components and structures within the Muccan Granitic Complex

The margin of the Muccan Granitic Complex comprises various tonalitic, granodioritic and granitic orthogneisses. Migmatized and strongly foliated tonalitic and a minor granodioritic components are intruded by younger, less deformed granitic rocks. However, the overall poor exposure of outcrops within the Muccan Granitic Complex hampers the establishment of contact relationships between the rock types. The older components display a gneissic foliation (S1) and are intruded by multiple generations of tightly folded leucosomes (F2; Fig. 2.6a), granitic gneisses and later potassic pegmatite bodies. Syn-deformational in-situ reworking of the older gneiss

21 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.6: a) Field photograph of a migmatitic tonalitic gneiss from the Muccan Granitic Complex,

displaying syn-deformational tightly folded (F2) leucocratic schlieren and associated melanosomic restite; b) stereoprojections (lower hemisphere) of S1/2 foliations and L2 lineations within the Muccan Granitic Complex. components is inferred from the gradation of gneissic banding into areas of intensive migmatization, and from the association of foliation-parallel and folded leucocratic patches and schlieren that are found adjacent to melanosomic/restitic zones (Fig. 2.6a). Based on the following observations, I describe the overall foliation within the Muccan

Granitic Complex as a composite S1/2 foliation: i) the gneissic banding (S1) of the older components is crosscut by leucosomes, ii) the leucosomes display a foliation (S2) that cuts parallel into the older gneissic banding, and iii) in plan-view, the F2 fold limbs of the leucosomes are aligned parallel to the gneissic banding of their host rock. This dominant E-W-trending S1/2 foliation is defined through the gneissic banding and the alignment of biotite and/or amphibole, and dips sub-vertically to the north (Fig. 2.6b). Mineral stretching on the S1/2 foliation is observed in biotite, amphibole and retrogressive chlorite, and defines steeply to sub-vertically NNW- plunging lineations (L2; Fig. 2.4; Fig. 2.6b). The contact between the Muccan Granitic Complex and the South Muccan Shear Zone to the south has been mapped as the first appearance of mafic rocks belonging to the greenstone belt (Fig. 2.4).

22 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.7: Petrographic image (cross-polarized light) of representative deformational micro-textures within a tonalitic gneiss (Muccan Granitic Complex) showing randomly oriented magmatic plagioclase (Pl) porphyroclasts and stages of dynamic quartz (Qz) recrystallization: bulging recrystallization at grain boundaries and within interstices of Pl porphyroclasts (1); undulose extinction within quartz subgrains (2); elongate quartz subgrains with straight grain boundaries due to subgrain rotation (3); sutured boundaries indicate grain boundary migration (4).

2.5.1.2 Deformational micro-textures and metamorphism within the Muccan Granitic Complex

At the micro-scale, the gneissic foliation within both the ‘older’ and the younger Muccan Granitic Complex components is defined by the sub-parallel alignment of tabular biotite and/or hornblende, and elongated ribbons of recrystallized quartz sub-grains, and the shape-preferred orientation of strained quartz grains (Fig. 2.7). Primary magmatic sub- to anhedral plagioclase forms randomly oriented porphyroclasts (Fig. 2.7). In quartz-rich zones sub-grain rotation and grain boundary migration is observed, while fine-grained bulging recrystallization of quartz dominates within the interstices of plagioclase porphyroclasts (Fig. 2.7). In zones of higher deformational intensity, quartz forms medium-grained elongated crystals with both straight and lobate grain boundaries. Undulose extinction of quartz grains is a common feature (Fig. 2.7). Some of the older granitic rocks display higher deformational intensity and can be classified as protomylonites, in which primary plagioclase is recrystallized at grain margins and forms slightly elongated porphyroclastic augen. The plagioclase porphyroclasts display both brittle fracturing and ductile deformation. The latter is inferred from the presence of rare undulose extinction. Plagioclase porphyroclasts are surrounded by a matrix of ribbons of elongated, fine- to medium- grained recrystallized quartz ± plagioclase sub-grains. Very fine-grained recrystallized mafic material forms schlieren and shear bands within the matrix. Some of the late-stage pegmatite intrusions are undeformed and comprise sub- to anhedral megacrysts of plagioclase, K-feldspar and quartz. Greenschist-facies retrogression of the gneisses is evident in partial replacement of biotite and hornblende by chlorite ± epidote, and in the seritization of plagioclase. Both foliation-parallel and

23 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

-crosscutting microscale calcite veins are observed occasionally, most commonly within zones of highest deformational intensity.

Fig. 2.8: Detailed structural map (1) and stereoprojections representative of the South Muccan Shear Zone - Central Fold Belt transition; location of the map is indicated in Figure 2.4.

24 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

2.5.2 South Muccan Shear Zone

2.5.2.1 Lithological components and structures within the South Muccan Shear Zone

South of the Muccan Granitic Complex, the ~1 km-wide high-strain South Muccan Shear Zone represents a lithological transition zone, incorporating both granitic components from the Muccan Granitic Complex, and mafic rocks and rare chert/banded-iron formation (BIF) from the greenstone belt, as meter- to decameter-sized elongated lozenges (Fig. 2.5; Fig. 2.8). Amphibolite- facies conditions are evident in plagioclase + hornblende assemblages in the mafic rocks. Both the granitic and the amphibolite lozenges display highly schistose to mylonitic textures. Subordinate massive meta-basalt, -gabbro and -pyroxenite form more competent, less deformed lozenges that are aligned with the regional foliation. The amphibolite-facies lozenges are surrounded by a matrix of foliated mafic schists comprising greenschist-facies assemblages (chlorite-actinolite-epidote- quartz-calcite) and often show intensive carbonate alteration. Both foliation-parallel and crosscutting microscale- to decimeter-wide calcite and quartz veins intrude the South Muccan Shear Zone. The foliation planes within the South Muccan Shear Zone are parallel to the gneissic foliation and inferred margin of the Muccan Granitic Complex (Fig. 2.4; Fig. 2.8). Lineations display similar orientations as observed in the Muccan Granitic Complex and are defined by amphibole and/or chlorite mineral stretching. Further to the south, granitic components from the Muccan Granitic Complex become absent, and greenschist-facies (chlorite-quartz-actinolite-epidote) mafic rocks dominate the Central Fold Belt (Fig. 2.8).

2.5.2.2 Deformational micro-textures and metamorphism within the South Muccan Shear Zone

Within the South Muccan Shear Zone granitic lozenges show progressive mylonitization (Fig. 2.9). Few of the primary magmatic porphyroclasts are preserved, and plagioclase is often recrystallized (Fig. 2.9). Preserved plagioclase porphyroclasts (15-40 Vol.%) form elongated augen that are surrounded by a matrix of recrystallized feldspar, quartz and chlorite. Within some quartzo-feldspathic bands, primary clasts are almost absent, and both quartz and plagioclase form strained and elongated medium-size grains. In contrast, within adjacent lozenges, randomly oriented coarse-grained euhedral plagioclase porphyroclasts (~40 Vol. %) show magmatic concentric zoning, and simple and polysynthetic twinning. The primary magmatic plagioclase clasts swim in a matrix of fine-grained recrystallized quartz and chlorite ± feldspar. Ribbons of chlorite-dominated matrix indicate a slight foliation. Likewise, different deformational intensities are observed within incorporated mafic lozenges belonging to the greenstone belt. Mylonitic amphibolites comprise coarse-grained, pre-tectonic plagioclase and light- to dark-green hornblende porphyroblasts within a highly sheared matrix of syn-kinematic, fine-grained, green- to bluish-green actinolite laths, alternating with bands of fine- grained chlorite, quartz, and/or calcite (Fig. 2.9). A late-stage reactivation of the shear foliation is

25 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.9: Petrographic images of representative micro-textures from within the South Muccan Shear Zone: a) tonalitic protomylonite (cross-polarized light) displaying progressive augen formation of plagioclase (Pl) porphyroclasts, which show evidence for bulging recrystallization at their grain boundaries, and are surrounded by finely recrystallized mafic shear bands (1); recrystallized quartz (Qz) forms strained and elongated subgrains; Bt, biotite; b) mylonitic amphibolite (plane-polarized light); pre-

tectonic hornblende (Hbl, M1) porphyroblasts within a highly sheared matrix of syn-kinematic actinolite (Act, M2) and sub-parallel, finely recrystallized bands of quartz (Qz), chlorite (Chl) and calcite (Cc); the cross-cutting calcite vein in the center of the image is affected by later reactivation of the shear foliation, indicated by slight dextral displacements and ingrowth of newly formed calcite laths (1). evident in foliation-parallel calcite laths growing into crosscutting calcite veins that display slight displacements along the primary foliation (Fig. 2.9). Other lozenges are formed by both greenschist- and amphibolite-facies mafic schists. Greenschist-facies mafic schists comprise chlorite-quartz-calcite ± feldspar ± actinolite assemblages, and amphibolite schists are dominated by hornblende and plagioclase. Within some

26 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______rare, larger (10-100m) (ultra-) mafic lozenges coarse-grained magmatic textures are preserved, in which hornblende replaces former pyroxenes.

2.5.3 Central Fold Belt

2.5.3.1 Lithological components within the Central Fold Belt

The Central Fold Belt consists predominantly of greenschist-facies mafic rocks that have been ascribed to the Mount Ada Basalt (Van Kranendonk, 2010). Metabasalts, subordinate mafic intrusions (gabbro/dolerite, anorthosite, pyroxenite) and bands of interspersed chert, siltstone, and BIF form foliation-parallel lozenges, which crop out as EW-trending ridges that are separated by carbonate-altered deformation zones (Fig. 2.4). The northern part of the Central Fold Belt is characterized by massive meta-mafic material and medium- to coarse-grained mafic intrusions with limited indicators of emplacement style. Rare vesicles and deformed pillow shapes have been observed in the otherwise massive mafic rocks. Mafic intrusions are occasionally associated with crosscutting dykes. The southern Central Fold Belt is dominated by ocelli-bearing pillow basalt flows. A rare ~ 50 m-sized body of anorthosite preserving primary magmatic plagioclase has been observed in the southern Central Fold Belt. The anorthosite likely forms a sill within the mafic flows, however contact relationships are obscured by intense deformation. Primary magmatic textures and mineralogy are occasionally preserved within lozenges in the Central Fold Belt, and are characterized by dominantly plagioclase and clinopyroxene laths within a fine-grained groundmass in metabasalts, and coarse-grained plagioclase and clinopyroxene in meta-gabbro/- dolerite. The overall mafic succession is frequently interlayered by m-thick chert/BIF horizons that are often associated with layers and lenses of metasiltstone (Fig. 2.4).

2.5.3.2 Structures and deformational events within the Central Fold Belt

The dominant EW-trending foliation (S1/2) within the Central Fold Belt dips sub-vertically to the north, parallel to the margin of the Muccan Granitic Complex (Fig. 2.4; Fig. 2.8). The entire Central Fold Belt is extensively penetrated by foliation-parallel, EW-trending, m- to 100 m-wide carbonate-altered shear zones, which separate more competent lozenges and ridges of mafic rocks and chert/BIF. These altered shear zones become increasingly abundant and wider towards the margin of the Muccan Granitic Complex (Fig. 2.4; Fig. 2.8). They are associated with multiple generations of carbonate veins that overprint each other, and are often associated with quartz. Where exposure is limited, quartz often dominates the regolith and helps to infer the extents of these shear zones. Some of the coarser-grained mafic material escaped the development of penetrative S1/2 foliation and forms relatively undeformed elongated lozenges within a matrix of foliated carbonate-altered mafic counterparts. Albeit metamorphosed to greenschist-facies (chlorite-epidote), the more competent lozenges preserve primary magmatic textures. In contrast, mafic lozenges in the vicinity of the deformation zones display internal tight to isoclinal folding

(F2) of an earlier foliation (S1). Similar to the regional S2 foliation, the S1 foliation is not always

27 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.10: Field photographs of deformational features in the Central Fold Belt; a) isoclinally folded (F2) mafic boudin (S1) affected by F3 folding within a highly sheared and altered, mafic host rock, incorporating boudinaged (S1/2) quartz veins (Qz); see inlet for interpretative drawing; b) S1 parallel calcite vein (v1), affected by isoclinal folds (F2) and later undulating folds (F3) within highly altered and foliated (S1/2) mafic host rock; see inlet for interpretative drawing; c) L-tectonites: lineated (L2) ocelli- bearing pillow basalts from the southern part of the Central Fold Belt. penetrative, incorporating cm- to dm-scale boudins and sigma-clasts of the sheared host rock (Fig.

2.10a). Cm-wide carbonate veins (v1) sit within S1 foliation (Fig. 2.10b).

F2 folds affect the entire Central Fold Belt, from cm- to map-scale (Fig. 2.4; Fig. 2.11). 100-m- scale F2 folds incorporate both the little deformed lozenges and the internally F2-folded lozenges within their sheared limbs (Fig. 2.4; Fig. 2.11), and sheared-off F2 fold hinges (Fig. 2.10a). Indeed, the global S2 foliation mainly consists of highly transposed, sheared-apart remnants of F2 folds.

28 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.11: Detailed structural map (2) and stereoprojections representative of the Central Fold Belt; location of the map is indicated in Figure 2.4.

Mineral stretching lineations (L2) on S2 foliation planes are defined by metamorphic chlorite and actinolite and stretched ocelli in the basaltic units, and plunge sub-vertically to the NNE (Fig. 2.4; Fig. 2.8). Within the Central Fold Belt, a transition from S~L fabrics in the north to L>S tectonites in the south is observed (Fig. 2.10c). S2-parallel ductile shear zones (f2) separate m-scale lozenges representing limbs and hinges of the map-scale F2 folds. Overall the average size of lozenges is slightly larger in the Central Fold Belt (~50 m) than in the South Muccan Shear Zone (~10 m)

(Fig. 2.8). The fold axes of F2 plunge sub-vertically in EW direction (Fig. 2.11). S2 also features cm- to dm-wide syn-deformational, foliation-parallel carbonate-quartz veins (v2). S2 foliation- parallel carbonate-alteration-deformation-zones are ascribed to this D2 deformation, based on the observation of subsequent D3 overprinting relationships.

D3 deformation is evident in micro- to m-scale F3 refolding of S2, F2 folds, and associated lozenges and carbonate zones (Fig. 2.4; Fig. 2.10a and b; Fig. 2.11). The F3 folds exhibit

29 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______asymmetric, steeply plunging, steeply inclined to sub-vertically plunging, upright geometry with dextral dome-up kinematics. The sub-vertical fold axes point towards the margin of the Muccan

Granitic Complex to the NE (Fig. 2.11). Locally, D3 cm-scale crenulation cleavages and kink-folds are observed in meta-basalts, predominantly in the hinge zones of m-scale F3 folds in the southern

Central Fold Belt. Occasionally, dextral D3 brittle, axial-planar-parallel faults (f3) are observed in the core of F3 folds, associated with cm- to m-wide v3 carbonate veins.

An array of sinistral NE-SW-trending ductile faults (f4), which developed local S4 foliations and associated drag-folds (F4) along strike (Fig. 2.4; Fig. 2.11), characterizes D4 deformation in the Central Fold Belt. The D4-related deformation shows the highest intensity where the faults impinged on former carbonate zones, resulting in lateral displacements of up to 100 m. The ductile flow within the fault-affected carbonate zones incorporated clasts, which inherit deformational structures ascribed to the D1 to D3 events.

Late-stage D5 deformation is evident in dextral EW-oriented displacements of the NNE-SSW- trending gabbroic mega-dyke within the Central Fold Belt, and focused primarily along some of the pre-existing D2 carbonate shear zones. The local reactivation of S2 foliation during D5 deformation resulted in EW offsets of the prior developed f4 faults and associated drag-folds. The most prominent of the f5 displacements demarcates the southern boundary of the Central Fold Belt with a lateral dextral offset of ~1.2 km (Fig. 2.4).

2.5.3.3 Micro-textures, deformational intensity and metasomatism within the Central Fold Belt

Spatial variability in the deformational intensity and mechanism within metabasalts of the Central Fold Belt, at the microscale, can be related to the map-scale structural features described above. Most importantly, the various observed effects of secondary deformational textures and greenschist-facies metamorphic equilibration seem not to be dependent on primary mineralogy alone, but on the heterogeneous strain distribution, often associated with localized fluid infiltration. In analogy to the map-scale observations, microscope studies reveal that often recrystallization and grain-size reduction within metabasalts are focused along discrete shear bands, separating micro-lozenges of more intact, undeformed minerals. This is evident for example, in the occurrence of areas of well-preserved magmatic textures separated by micro- fractures and shear bands, now filled with chlorite or calcite. Highly lineated ocelli-bearing rocks are characterized by a homogeneous grain-size reduction by dynamic recrystallization of both feldspar and mafic (i.e., chlorite) components. Polyphase deformation is indicated by the presence of coarse-grained calcite veins within former ocelli that are cut by quartz veins and finely recrystallized mafic shear bands surrounding the elongated ocelli (Fig. 2.12a). Within some specimen, the early S1 foliation is preserved in boudinaged and isoclinally folded (F2) micro-veins of quartz (Fig. 2.12b). The effect of D3 deformation is recorded in folded S2 foliation and occurrence of S-C fabrics (Fig. 2.12b). Within metasiltstones, chlorite flakes define the S1/2 foliation and are partly altered to clays (e.g., Illite).

30 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.12: Petrographic images of representative micro-textures from within the Central Fold Belt: a) polyphase deformation within a lineated ocelli-bearing metabasalt (plane-polarized light): the lower part of the image represents a stretched former ocelli that is finely recrystallized to feldspar (Fsp), chlorite, quartz and opaque minerals; the ocelli is penetrated by coarse-grained calcite (Cc) veins, which in turn are cut by a thin band of finely recrystallized quartz (Qz) subgrains, marking the sheared boundary of the ocelli; the E-W-trending shear band in the center of the image is dominated by alternating finely recrystallized mafic material (1) and ribbons of a second generation of coarser grained chlorite flakes (Chl-2) and associated fine-grained opaque minerals; syn-kinematic rutile forms elongated grains associated with quartz in its pressure shadows; a late-stage calcite vein cross-cuts the mafic shear band; 2, epoxy-filled fracture; b) mafic schist (plane-polarized light), in which alternating bands of finely

recrystallized quartz (Qz) and chlorite (Chl-1) define the S1/2 foliation; F2 isoclinally folded micro- boudins comprising quartz subgrains indicate the earlier S1 foliation (1); F3 folding is evident in crenulations and fold-axial-parallel dextral f3 shear bands associated with renewed chlorite (Chl-2) growth.

Intensive metasomatism by aqueous/silicious and carbonaceous fluids is not only apparent in the formation of hydrous greenschist-facies assemblages, but in both foliation-parallel and

31 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______crosscutting calcite veins, secondary overgrowth by quartz aggregates, and syn-kinematic ribbons and bands of recrystallized quartz sub-grains incorporated in the foliation. Secondary medium- grained euhedral pyrite crystals overgrow the fine-grained chlorite-calcite matrix of some basaltic rocks. In other metabasalts, post-kinematic rutile is observed overgrowing chlorite-rich shear bands. Where primary magmatic textures are preserved, finely recrystallized calcite dominantly forms pseudomorphs after plagioclase laths, and chlorite after former pyroxene. Both plagioclase and clinopyroxene are occasionally preserved. Within the most intensively deformed, fluid- induced shear zones, the mafic protoliths are completely replaced by carbonate minerals ± quartz.

2.5.4 Southern Low-Strain Belt

2.5.4.1 Lithological components and structures within the southern Low-Strain Belt

The Low-Strain Belt is weakly metamorphosed to greenschist/sub-greenschist-facies. It represents a well-preserved and less deformed stratigraphic section through the upper Mount Ada Basalt, the Duffer Formation, and the unconformably overlying Strelley Pool Formation (Fig. 2.4). Throughout the Low-Strain Belt, bedding directions are overturned and facing is to the south, as indicated by the orientation of sedimentary and pillow structures. Bedding planes of the younger Strelley Pool Formation dip more steeply and at a slightly oblique strike angle to the underlying Warrawoona Group, reflecting the angular nature of the unconformity (Fig. 2.4, inset; Fig. 2.5). Stratigraphic conformity between the Mount Ada Basalt and the Duffer Formation could be demonstrated locally. However, lithological boundaries are often sheared and tectonically reactivated. EW-trending anastomosing shear zone networks pervade the Low-Strain Belt, enveloping large (100-m scale), occasionally sigmoidal lozenges in the basaltic rocks of the Mount

Ada Basalt (Fig. 2.4; Fig. 2.10). The shear zones show foliations generally parallel to S2 in the Central Fold Belt (Fig. 2.4; Fig. 2.13). They are often associated with carbonate alteration within the basaltic rocks of the Mount Ada Basalt, and silicification within the felsic rocks of the Duffer Formation and the Strelley Pool Formation, respectively.

The earliest observed deformational structures (D1) in the Low-Strain Belt are an array of ~NS- trending brittle faults (f1A), which are found only within a ca. 10-30 m thick succession of siltstone and mudstone, partly of volcanogenic origin (crystal shards, pumice fragments), interlayered with the upper pillow basalts of the Mount Ada Basalt (Fig. 2.14a). The faults are aligned perpendicular to the bedding in the upper succession and gradually curve towards W, until they become parallel to the bedding planes. In between the faults, smaller antithetic faults are developed (Fig. 2.14a). The extent of displacement along these faults is unclear, but cannot be significant due to the relatively small scale of the faults (~20 m height). The faults are associated with cm-scale en échelon tension gashes.

In the SE corner of the study area, an array of 100-m-scale NE-SW-trending faults (f1B) dextrally offsets the Duffer Formation and parts of the underlying Mount Ada Basalt up to 300 m

32 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.13: Detailed structural map (3) and stereoprojections representative of the Low-Strain Belt; location of the map is indicated in Figure 2.4.

(Fig. 2.13), but does not affect the unconformably overlying Strelley Pool Formation (Fig. 2.4). To the north, the faults are cut by EW-trending anastomosing shear zones (D2).

A later array of brittle, up to 10 m-wide NNE-SSW-trending faults (f3) affects the entire Low- Strain Belt (Fig. 2.4) and sinistrally offsets lithological boundaries by up to 200 m. The brittle nature of the faults is characterized by silicified fault breccia containing clasts of quartzitic sand- and siltstone of the surrounding Strelley Pool Formation and Duffer Formation. Within the underlying Mount Ada Basalt, the f3 faults bend into an associated NE-trending fault, which separates a wedge-shaped block of greenstone material in the SW of the study area (Fig. 2.4). The latter fault is inferred from carbonate-altered damage zones, but can be traced to the west of the study area (satellite images), where it curves into parallel alignment with the NNE-trending fault array and caused significant sinistral displacements within the Strelley Pool Formation. In the SE corner of the study area the f3 faults merge with earlier dextral f1B faults within the Duffer Formation, where a mega-breccia is observed (Fig. 2.4; Fig. 2.13). A prominent EW-trending dextral f4 fault (f5 in Central Fold Belt) displaces the NNE-trending gabbroic mega-dyke by ~1.2 km at the northern domain boundary (Fig. 2.4).

33 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.14: Field photographs from the Low-Strain Belt: a) Syn-depositional dextral faults (f1A) and associated antithetic faults in a sedimentary succession within the Mount Ada Basalt; b) well- preserved ocelli-bearing pillow basalts with inter-pillow hyaloclastite, Mount Ada Basalt; see inlet for detail; c) stromatolitic structures in chert on top of a regressive sequence within the Duffer Formation.

34 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

2.5.4.2 Micro-textures, deformation and metamorphism within the Low-Strain Belt

The non-foliated pillow- metabasalts of the Mount Ada Basalt comprise greenschist to sub- greenschist-facies mineral assemblages dominated by chlorite, calcite and clay minerals. Preservation of magmatic textures (Fig. 2.15a) and relic primary mineralogy (i.e., plagioclase, clinopyroxene, opaque minerals), however, is widespread. The matrix of the basalts is mostly finely recrystallized to assemblages of chlorite, albite, quartz and c1ay minerals, and occasional epidote. Some rocks contain finely recrystallized areas of associated calcite, quartz and white mica. Albitized plagioclase forms laths and former sub- to euhedral phenocryst, and often shows preserved twinning. Pyroxenes are mostly replaced by fine-grained aggregates of clay minerals. Vesicles, where present, are filled with microcrystalline quartz and/or radial calcite laths. Fractures and veins are mostly filled by medium-grained chlorite. Euhedral pyrite and other opaque phases often overgrow the matrix or occur within chlorite-filled fractures. Overall, the basaltic rocks of the Low-Strain Belt resemble the little deformed mafic lozenges in the southern part of the Central Fold Belt. A metamorphic foliation in the Low-Strain Belt basalts is only observed within local shear zones, and defined by the alignment of chlorite flakes in metabasalts. Fold microstructures as observed within the Central Fold Belt are completely absent within the Low-Strain Belt. Where pillow basalts are affected by shear, strain is mostly localized within the fine-grained inter-pillow material, surrounding the slightly flattened pillows and exaggerating the anastomosing nature of the shear zones. Within shear zones, an increased silicification, inferred from secondary overgrowth by quartz aggregates and abundance of microscale quartz vein networks, is observed. Similar to the basalts, the volcaniclastic sedimentary and felsic volcanic rocks of the Duffer Formation often preserve primary features. In the volcanic rocks remnants of magmatic textures, such as crystal laths and euhedral phenocrysts of former feldspars and quartz are present. The felsic volcanic rocks are often intensively penetrated by networks of quartz-filled micro-fractures. Within the clastic sediments, primary lamination and shape and size variations of grains are well- preserved. The volcanogenic sedimentary layers preserve internal textures of pumice, accretionary lapilli and rock fragments (Fig. 2.15b). The quartz sandstone of the Strelley Pool Formation comprises undeformed, well-rounded coarse-grained quartz clasts surrounded by recrystallized microcrystalline silica cement (Fig. 2.15c). Some conglomeratic layers with larger pebble-sized grains grading into sandstone are observed. The sandstone is often crosscut by mm- to cm-wide veins filled by coarse-grained, comb-textured quartz. Within m-wide fault zones, the sandstone is more intensively recrystallized and original grain boundaries are not preserved.

2.5.4.3 Lithostratigraphy of the Low-Strain Belt

a) Mount Ada Basalt

South of the prominent f4 fault, a succession of dominantly ocelli-bearing pillow basalt flows of the upper Mount Ada Basalt is exposed almost continuously over a stratigraphic thickness of

35 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.15: Petrographic images of well-preserved primary micro-textures from within the Low-Strain Belt: a) lower greenschist-facies massive mafic rock with primary magmatic spinifex texture, Mount Ada Basalt (plane-polarized light): randomly oriented bladed crystal laths (up to 2 cm) of former olivine or pyroxene are replaced by an inner core of chlorite (Chl) and quartz (Qz) and an outer mantle of brown clay minerals (1); the matrix is formed by chlorite pseudomorphs after former plagioclase microlaths, and interstitial secondary quartz (Qz); b) volcaniclastic tuff (plane-polarized light) from the Mount Ada Basalt comprising altered pumice fragments with quartz-filled vesicles (1), and rock fragments with preserved magmatic textures (2); c) quartzitic sandstone from the Strelley Pool Formation (cross- polarized light) displaying well-rounded coarse-grained quartz (Qz) within a matrix of recrystallized microcrystalline silica (quartz) cement (1).

~550m (Fig. 2.4; Fig. 2.14b; Fig. 2.16). South-facing younging directions are indicated by pillow shapes. Flow interfaces within the succession cannot be readily recognized in the field. However, increased abundance of vesicles and larger-sized ocelli (up to 3 cm diameter) are observed directly

36 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______above interlayered chert-siltstone horizons. Overall, vesiculated flows are rare. Both ocelli and matrix include quartz and/or calcite filled amygdules. Chilled pillow margins and inter-pillow hyaloclastites are relatively well preserved (Fig. 2.14b). The basalts contain greenschist- to sub- greenschist-facies mineral assemblages of chlorite-actinolite-albite ± epidote ± quartz ± calcite and clay minerals. Partial preservation of relic primary magmatic mineralogy and texture is widespread and indicates that the majority of basalts can be classified as tholeiites. Particularly in the vicinity of intercalated chert bands, the basalts are affected by secondary silicification, manifested in penetrating sub-millimeter-wide quartz veins and secondary fine- to medium-grained subhedral quartz aggregates. In zones of higher deformational intensity (EW-trending shear zones, faults), the basalts are foliated and often show intensive carbonate alteration, equivalent to observations in the Central Fold Belt. The relatively homogeneous succession of pillow basalts is interlayered by subordinate more massive, spinifex-textured mafic flows/sills and intruded by massive gabbro/dolerite sills. Hiatus in basaltic volcanism are evident in interlayered thin (0.5-3 m) chert bands. Siltstone lenses directly above the chert bands contain fragments and shards of crystals and pumice fragments, indicating a volcanic origin. Another volcaniclastic sedimentary layer, associated with chert horizons, forms a ca. 10-30 m sequence of laminated mudstone beds grading into siltstone, and is affected by the f1A faults (Fig. 2.16).

b) Felsic volcanic - sedimentary succession (Duffer Formation)

In the study area felsic volcanic flows of the Duffer Formation are found to intercalate with volcaniclastic sediments previously mapped as Mount Ada Basalt (Van Kranendonk, 2010). Because the intercalated felsic volcanic - sedimentary rocks rest conformably on top of basaltic flows of the Mount Ada Basalt as a well-defined package, I ascribe this entire succession to the Duffer Formation (Fig. 2.4). In the central and western part of the study area, the base of the Duffer Formation consists of up to 200 m of siltstones that conformably overly the pillow basalts of the Mount Ada Basalt (Fig. 2.16). At the base of the siltstones, a ca. 1 m-thick unit of repeating cm-thick, upward-fining conglomerate to siltstone beds is observed locally (Fig. 2.16, III). The base succession comprises crystal shards, clasts of mafic volcanic rock and fragments of pumice (i.e. pumiceous tuff). The overlying felsic siltstone layers comprise accretionary lapilli and agglomerations. The entire siltstone sequence is weakly recrystallized, but primary lamination is mostly preserved. The cm- to dm-thick siltstone beds occasionally alternate with thinner, fine- laminated mudstone beds. The sediments are often penetrated by a quartz vein network. Rare medium-grained ~50 m-wide mafic bodies (dolerite) intruded the siltstones as sills or dykes (Fig. 2.16, IV). To the south, the siltstone transitions into a sequence of repeated 1 to 2 m-thick conglomerate- siltstone-mudstone cycles (Fig. 2.16, section III) with consistent upward-fining stratigraphy. A ca. 2 m-thick felsic pumiceous tuff horizon rests on top of this turbiditic sequence (Fig. 2.16, IV). Above the felsic tuff, fine-laminated silt- and mudstones display diffuse, mm-scale internal folding

37 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.16: Stratigraphic columns from the Low-Strain Belt; positions of columns I to V are indicated in Figure 2.4; refer to text for description. and decimeter-scale mudstone lenses are folded. The siltstones and mudstones gradually coarsen up into sandstones, which become increasingly interlayered with m-thick, recrystallized blue chert bands (Fig. 2.16). The uppermost chert layer reaches a thickness of up to 30 m. Its base is intensively recrystallized and contains multiple chert/silica veins that vertically crosscut the underlying sandstones (Fig. 2.16, III, IV). The chert/silica veins carry brecciated clasts of previously fractured, as well as primary laminated chert fragments. The preservation of primary

38 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______laminated chert increases towards the top of the unit, where abundant stromatolitic structures are observed (Fig. 2.14c; Fig. 2.16, III, IV). The stromatolite chert can be traced laterally over up to 2 km across the study area (Fig. 2.4). An up to 50 m-thick, red/purple-colored (ultra-) mafic volcanic flow lies conformably above the stromatolite chert (Fig. 2.16, III, IV). The flow is highly altered but both relic magmatic spinifex textures and pillow structures are preserved. To the east, the volcanic flow is cut-off by a dextral D1 fault. In the vicinity of the fault the flow becomes highly vesiculated and associated with a gossanous crust. Between the volcanic flow and the Strelley Pool Formation unconformity, intensively silicified and sheared sandstones crop out (Fig. 2.4; Fig. 2.16). In the eastern part of the study area the stratigraphy overlying the Mount Ada Basalt is different and characterized by an increasing volume of felsic volcanic flows and sills (Fig. 2.16, I, II). Above the contact with the Mount Ada Basalt, felsic volcaniclastic layers and felsic volcanic rocks (quartz-feldspar-pyrite) with preserved relic primary magmatic textures occur. These units are overlain by silicified pumiceous sandstone, interlayered with massive recrystallized rhyolite flows with quartz-filled vesicles and an uppermost medium-grained felsic volcanic unit with preserved relic magmatic textures (Fig. 2.16, I, II). Up to the unconformity marked by the overlying Strelley Pool Formation, the Duffer Formation reaches a stratigraphic thickness of ~580 m in the central part of the study area (Fig. 2.16, III). To the west, the Duffer Formation pinches out at a low angle beneath the unconformity (Fig. 2.4).

c) Strelley Pool Formation

The unconformable contact between the Duffer Formation and the Strelley Pool Formation is either not exposed, or disturbed by bedding-parallel shear-deformation. The angular nature of the unconformity is evident in the difference of the orientation of the two formations (Fig. 2.4, inset; Fig. 2.5). In the far eastern part of the study area, a silicified ~15 m-thick conglomerate of poorly rounded and unsorted clasts of quartz, feldspar and abundant red-banded jasperous chert fragments marks the base of the Strelley Pool Formation (Fig. 2.16, II). The base conglomerate grades into an up to ~1 km-thick quartzitic sandstone package of more well-rounded quartz and minor feldspar clasts. Occasional cross-bedding structures and upward fining within rare ca. 20 cm conglomerate- sand-siltstone beds indicate consistent younging to the south. The entire Strelley Pool Formation is recrystallized to varying degrees, and penetrated by bedding-parallel and crosscutting chert/silica vein networks on the scale of tens of meters.

2.6 Discussion

Based on my lithostratigraphic and structural observations, I propose a model of protracted structural development of the studied western Doolena Gap greenstone belt that spans the entire mid- to upper crustal evolution of the East Pilbara Terrane and later Pilbara Craton-wide-events during the Palaeo- to Neoarchaean (Fig. 2.17).

39 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

2.6.1 Deformation history

A combined total of five deformational events (D1-D5) are recognized within the structural domains. Structural data is evaluated geometrically and kinematically with hemispherical projections (Stereonet 9; Allmendinger, 2015). The reconstruction of geometry and kinematics through time requires back-rotations accounting for progressive deformation. I follow the actual back-rotation workflow and work my way backwards in time in the following discussion, starting with the youngest deformation features. A synthesis of the reconstructed deformational events within each domain, including my interpretation of inter-domain correlations between deformation phases, is given in Figure 2.17.

The youngest observed structures are the EW-trending f5 faults (f4 in the Low-Strain Belt) that affect the entire greenstone belt and offset the mafic mega dyke with a dextral strike-separation of ca. 1.2 km (Fig. 2.4). Because the dyke is ascribed to the East Pilbara Terrane-wide ~2772 Ma

Black Range Dolerite Suite (Wingate, 1999), D5 deformation must have occurred in the

Neoarchaean, or later. Structures that predate D5 faulting show ductile behavior within the Central Fold Belt and brittle deformation in the Low-Strain Belt. The marked difference in deformation style and intensity across the most prominent f5 fault leads us to conclude that it most likely accommodated substantial subvertical displacement, juxtaposing the Central Fold Belt with the

Low-Strain Belt during D5. From both field and microscale observations it is evident that f5 faulting locally reactivated earlier foliations (S1/2) in the Central Fold Belt and the South Muccan Shear Zone (Fig. 2.17). It is therefore possible that some relative vertical movement occurred during an earlier deformation event.

The similar orientation and kinematics of the sinistral NNE- to NE-trending ductile f4 faults and associated drag-folds (F4) within the Central Fold Belt and the brittle f3 faults within the Low-

Strain Belt suggest contemporaneous faulting that I ascribe to an overall D4 event (Fig. 2.17). The

Central Fold Belt represented deeper crustal levels during the D4 event (Fig. 2.17). As evident in satellite images, the D4 faults within the Low-Strain Belt penetrate the entire greenstone succession to the south of the study area, including rocks of the De Grey Superbasin, but do not affect rocks of the Fortescue Group. The f4 faulting can thus be constraint in time to the Mesoarchaean (Fig. 2.17).

The correlation between D3 structures in the Central Fold Belt and structures in the Low-Strain

Belt is ambiguous. Within the Central Fold Belt, D3 structures include the F3 dextral dome-up folds and associated f3 hinge faults, as well as the carbonate-fluid influx (v3) and deformation- reactivation that affected previously developed D2 shear zones. Within the Low-Strain Belt, EW- trending silica veins and deformation zones of the Strelley Pool Formation resemble the high- strain zones and anastomosing shear zones within the underlying Warrawoona Group. However, the Strelley Pool Formation shear zones are predominantly bedding parallel and thus oriented at a slight angle to the anastomosing shear zones within the Warrawoona Group. It is possible that e.g. bedding-perpendicular grain-size variations in the Strelley Pool Formation provided rheological anisotropies to develop faults at a slightly different angle than within the Warrawoona Group

40 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.17: Synoptic timeline and interpretative synthesis of major structural events and kinematics with inter-domain correlation; structural data is displayed in lower-hemisphere stereoprojections; refer to text for description and constraints on the relative timing of deformational events.

41 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

(e.g. Yin and Ranalli, 1992). In turn, the dome-up dextral block-rotation as observed in the ‘Brain

Hill’ area (Fig. 2.13) suggests contemporaneous D3 deformation (or reactivation) in the

Warrawoona Group of the Low-Strain Belt and D3 deformation in the Central Fold Belt. Since the Low-Strain Belt exhibits similar anastomosing shear zones below and above the Strelley Pool

Formation unconformity, I ascribe the D3 event to post-Strelley Pool Formation deposition.

Inspecting the kinematics of D3 back-rotated to a horizontal orientation of the Strelley Pool Formation, the tilting did not change the kinematics and orientation significantly. Back-rotated F3 fold axes within the Central Fold Belt plunged sub-vertically to the SW, without a change in the overall dextral, dome-up kinematic movement (Fig. 2.17).

D2 deformation represents the most intense deformation event in the study area. It includes the tight to isoclinal F2 folds in the Muccan Granitic Complex and the Central Fold Belt, the global S2 foliation and mylonitization in the South Muccan Shear Zone, and the L2 mineral stretching lineations in the Muccan Granitic Complex, the South Muccan Shear Zone and the Central Fold Belt, and associated L-tectonite development in the southern Central Fold Belt (Fig. 2.17). Stereo- projections unambiguously demonstrate a geometrical correlation between S1/2 foliations in the Muccan Granitic Complex, the South Muccan Shear Zone, the Central Fold Belt, and locally in the

Warrawoona Group of the Low-Strain Belt (Fig. 2.17). Within the Strelley Pool Formation, D2 structures are absent, and considering the intensity of D2, it is unlikely that the Strelley Pool Formation could have escaped this event. Back-rotation of the Strelley Pool Formation demonstrates that at this stage, the Warrawoona Group shows a SSW-facing tilt of approx. 45° (Fig. 2.17). The extent of the tilt implies significant tectonic activity that I interpret as the result of initial doming of the Muccan Granitic Complex to the north. This interpretation is supported by the sub-parallel alignment of the intra-domain foliations (S1/2), the Warrawoona Group bedding orientations, and the dome-keel interface. I argue that the overall geometry of D2-related structures was developed at this stage, and therefore is best represented through back-rotation to the horizontal layering of the Strelley Pool Formation (Fig. 2.17), keeping in mind that it is the result of progressive deformation up to large finite strain. The limited nature of exposure within the

Muccan Granitic Complex hampers conclusive correlation of F2 folds between the Muccan Granitic Complex and the Central Fold Belt. However, observable fold shapes and orientations indicate similar deformational mechanisms and geometries and thus a contemporaneous development. The partly syn-anatectic nature of F2 folds within the Muccan Granitic Complex provides evidence for high temperatures and possibly elevated fluid pressures at the dome margin during F2 formation, as expected near large intrusions. The thermal and fluid input of the dome could also explain the large intensity of ductile deformation observed at its interface. Effective viscosity of rocks is highly sensitive to temperature and water fugacity (e.g. Burlini and Bruhn, 2005; Rybacki et al., 2006; Bürgmann and Dresen, 2008).

S1 foliations within the Muccan Granitic Complex and the Central Fold Belt could not be correlated conclusively, and it remains unclear if the development of early S1 gneissic foliation in the Muccan Granitic Complex corresponds to the S1 foliation and foliation-parallel v1 carbonate veins in the keel rocks of the Central Fold Belt. Within the Central Fold Belt, the composite nature

42 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______of S1 and S2 and evidence for a shear component in S1 development suggest a similar deformational mechanism of S1 and S2. It is possible that the early foliation developed during the initiation of the major D2 event along the margin of the rising Muccan Granitic Complex.

How the timing of f1 faults within the Low-Strain Belt relates to deformation phases in the

Central Fold Belt and Muccan Granitic Complex also remains somewhat unclear. The fact that f1A faults are restricted to a sedimentary succession within the upper Mount Ada Basalt suggests a syn-depositional origin. Therefore, f1A faults likely developed before the tilt of the Warrawoona Group. I performed two-fold back-rotation to the palaeo-horizontal Warrawoona Group layering for evaluation of the palaeo-orientation and kinematics. The rotation reveals a listric shape of the faults with normal west-block down kinematic movement in an extensional regime (Fig. 2.17). It remains unclear if early ductile D1 deformation already initiated the subsequent tilting within lower crustal levels (Central Fold Belt, Muccan Granitic Complex) at this time (Fig. 2.17). A more or less contemporaneous nature of D1 in the Low-Strain Belt and D1 in the Central Fold Belt is supported by the regressive clastic sequence in the Duffer Formation, indicating uplift and thus possibly tectonic activity as a response to early Muccan Granitic Complex doming (Fig. 2.17).

2.6.2 Structural evolution and tectonically controlled deposystems

2.6.2.1 Depositional environment and early extension

The earliest exposed rocks within the studied western Doolena Gap greenstone belt are ascribed to the Mount Ada Basalt (e.g. Van Kranendonk, 2010). The present study shows that the bulk of the Mount Ada Basalt forms a pile of tholeiitic, ocelli-bearing pillow flows, episodically interlayered with thin chert or BIF horizons. The widespread presence of pillow basalts and chert, combined with the scarcity of clastic sediments and vesicles within the basaltic flows is consistent with a low-energy, deep marine depositional environment. The dominance of volcanic style with expected short flow periphery, and absence of volcanic sheet flows suggest relative proximity to sources of volcanic outpourings (e.g. Gregg and Fink, 1995). I interpret the observed numerous layer-parallel and discordant mafic bodies, which increase in abundance stratigraphically downwards within the Mount Ada Basalt, as sills and dykes that represent intracrustal feeder systems to the supracrustal pillow flows. Considering the wide distribution of the Mount Ada Basalt within greenstone belts throughout the East Pilbara Terrane (Van Kranendonk et al., 2006; Hickman, 2008), the absence of sheet flows and dominance of pillow basalts could be explained by volcanic eruptions fed by extensive systems of intracrustal magma chambers and feeder systems, possibly through syn-depositional crustal fissures and/or faults. Deposition of the Mount

Ada Basalt was accompanied, at least locally, by development of listric normal faults (f1A), indicating an extensional regime in the upper crust that could support the interpretation of intracrustal feeder systems and explain the eruption of predominantly pillow basalts through extension related fractures.

43 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Overlying the Mount Ada Basalt, the onset of the Duffer Formation within the study area is characterized by an almost abrupt cessation of basaltic volcanism, and the change to a more diverse depositional environment dominated by intercalated felsic volcanic rocks and volcaniclastic sediments. The spatial characteristics of the Duffer Formation show that felsic volcanic flows are more prevalent to the east and virtually absent to the west, where authigenic tuffs and volcaniclastic sediments are observed. I interpret the large dextral f1B faults within the Duffer Formation and upper Mount Ada Basalt as of syn-depositional origin. The poor preservation of the faults due to later overprinting and reactivation prevents a conclusive interpretation. However, similar structures have been reported to the east, where preservation is better (deVries et al., 2006). The faults display west-block down normal faulting after back- rotation. This is in accordance with the observed depositional change from felsic volcanic flows with interlayered volcaniclastic sandstones in the east, possibly forming a topographic high at the marginal basin, towards more finer-grained laminated silt- and mudstones with occasional soft- sediment disturbance interlayered with turbiditic sequences in the west, where the basin deepened and was fed by marginal and fluviatile input from the east. The lower felsic succession in the western part of the study area comprises pumice- and accretionary lapilli-bearing tuff beds likely presenting authigenic ejecta supplied from the felsic volcanism to the east into the basin to the west.

2.6.2.2 Main phase of granitic doming and associated greenstone sagduction within the western Doolena Gap greenstone belt

Within the Duffer Formation, ongoing deposition in the central-western part of the study area displays a transition from silt- and mudstones to sandstones and finally stromatolitic chert, reflecting an overall regressive sequence towards progressively shallower subaqueous sedimentary facies. I interpret this observation as the result of crustal uplift that predated subsequent tilting of the Warrawoona Group within the study area. The SSW-facing ~45° tilt of the Warrawoona Group beddings prior to Strelley Pool Formation deposition indicates significant tectonic activity that I associate with initial Muccan Granitic Complex dome component emplacement to the north. Uplift and subsequent tilting as a response to dome emplacement resulted in non-deposition and/or erosion of the upper Warrawoona Group in the western Doolena Gap greenstone belt. A vesicular, spinifex-textured, but highly altered (ultra-) mafic flow that directly overlies the stromatolitic chert within the Duffer Formation resembles part of a laterite profile, preserving saprolite and concretions of gossanous crust. If this in fact is so, subaerial weathering could have occurred during uplift and erosion of the Warrawoona Group. Previously proposed stratigraphically upward increase in subaerial deposition within the Warrawoona Group (Barley et al., 1979), and the interpretation that felsic volcanic centres formed within a larger basin (e.g. DiMarco and Lowe, 1989) supports the above interpretation of uplift as a result of dome emplacement. Subsequent tilting, after deposition of the Duffer Formation, can be correlated with emplacement of the Tambina Supersuite. 3410 Ma metamorphic zircon within

44 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

Fig. 2.18: Cartoon illustration of major structural events; refer to text for descriptions (a-d). domes (U-Pb SHRIMP; Nelson, 1999), and 3440 Ma metamorphic monazite within greenstone metasedimentary rocks (U-Th-Pb SHRIMP; François et al., 2014) document additional evidence for this early tectono-thermal activity in the East Pilbara Terrane. The oldest records of metamorphism and deformation associated with this early tilting, are found in slightly deeper levels of the upper crust that are now closest to the Muccan Granitic

Complex (Fig. 2.18a). I ascribe an early amphibolite-facies metamorphism (M1), evident in pre-

45 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______tectonic hornblende porphyroblasts, of mafic rocks along the margin of the Muccan Granitic Complex as a result of a thermal contact aureole that developed during emplacement of granitoid magmas within the dome. The early S1 foliation associated with early carbonate-vein injection (v1) within the Central Fold Belt, is characterized by deformational features (e.g. asymmetric boudins, sigma clasts) that indicate non-coaxial shear deformation, which may have developed along the margin of the ascending dome (e.g. Dixon, 1975). Due to the presence of a S1 foliation and rather heterogeneous protolith composition (providing rheological contrast and anisotropy in a horizontally stratified mid-lower crust), the tight F2 folds in the Central Fold Belt likely initiated as upright buckle folds before being sheared passively to high finite strain during ongoing tilting and shear of the greenstones. Tightening of F2 folds to parallel alignment of the fold limbs and fluid- induced hinge faults defined the new S2 foliation. The protracted S2-parallel ~dip-slip stretching in the Central Fold Belt led to transposition of the F2 folds, resulting in rootless folds surrounded by anastomosing shear zones. Such a transposition mechanism has been described by e.g. Callender (1983) for deformed stratigraphic packages within the Precambrian basement of northern New Mexico (USA). In the Doolena Gap greenstone belt, syn-deformational influx of carbonaceous fluids caused syn-kinematic veins (v2) and is also seen as alteration of shear zones separating limbs and hinges of F2 folds. In map view, this process produced the disconnected wavy competent lozenges interlayered with wide, highly deformed carbonate-altered schistose shear zones (Fig. 2.4). The process caused imbrication of different structural levels by sub-vertical dome-up thrusts (Fig. 2.18a). Strain within the interface shear zone, the mylonitic South Muccan Shear Zone, is generally large, distributed heterogeneously and non-coaxial. The South Muccan Shear Zone accommodated protracted differential movement between the dome and greenstone keel. Folded anatectic schlieren within the dome margin exhibit geometry similar to F2 folds in the Central Fold Belt and may thus have formed synchronously (Fig. 2.18a). Within the Central Fold Belt the transition from L~S to L>S tectonites towards the south implies increasing sub-vertical constrictional flow towards a zone of sinking to the SW. The interpretation is similar to observations within the Warrawoona Syncline, where a zone of sinking characterized by intensive vertical L-tectonite development has been identified to the west of the greenstone synclinorium (e.g. Collins et al., 1998; François et al., 2014). It is interesting to note that the direction of constrictional flow of the greenstone keel occurred towards a triple junction between ascending domes (Muccan, Mount Edgar, North Pole; Fig. 2.1), rather than towards synclinal centres between two adjacent domes. Similar observations have been made within the Western Dharwar Craton (South India), where vertical foliation-trajectories define sagduction related triple points between adjacent granitoid domes within deeper levels of the crust (e.g. Bouhallier et al., 1995).

46 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

2.6.2.3 Progressive dome-and-keel formation: partial convective overturn versus horizontal tectonics

The younger D3 event mainly affected the keel rocks, but may have caused renewed bulging recrystallization of quartz within grain interstices and distinct bands of the dome gneisses. The F3 folds within the Central Fold Belt are characterized by wider inter-limb angles (90°-50°) and associated widespread kink-folding, suggesting lower finite strain compared to the earlier D2 deformation. Within the Low-Strain Belt, anastomosing shear zones sub-parallel to the bedding are ascribed to D3 deformation that resulted in the rotation of 100 m-scale blocks, in accordance with the dextral dome-up kinematics indicated by F3 folds in the Central Fold Belt (Fig. 2.18b). I ascribe the timing of D3 to after Strelley Pool Formation deposition. My structural analysis indicates a shallow SW plunge of F3 fold axes (Fig. 2.17). In conjunction with the open to close, upright and asymmetric fold shape with dextral dome-up kinematic indications, I propose buckling in combination with a shear component in the development of F3 folds. The dextral dome-up shear-sense is further indicated along reactivated S2 planes and f3 hinge faults. D3 also brought renewed or reactivated carbonate-fluid influx along the fold-axial-plane parallel f3 faults within the Central Fold Belt (Fig. 2.18b).

Overall, I interpret the D3 structures as the result of combined tightening of the keel syncline (buckling) and further dome-up dextral movement relative to the Muccan Granitic Complex. The sub-vertical flow of the keel rocks is roughly in agreement with the zone of sinking to the SW that is defined by the earlier D2 structures. However, the more westward flow orientation inferred from D3 structures likely describes a progressive radial rotation from initial one-sided tilting towards a triple junction at the side of the syncline. The observation is in agreement with the angular rotation across the Strelley Pool Formation unconformity, as well as the younger angular unconformities and related depocenter shifts in the De Grey rocks (Nijman et al., 1999) that ultimately define the late-stage zone of sinking as evident in regional geological-structural maps (Van Kranendonk, 2010).

The timing of D3 deformation after Strelley Pool Formation deposition is in agreement with previous models that ascribe the main phase of partial convective overturn and associated dome- and-keel development within the East Pilbara Terrane to after the deposition of the uppermost Kelly Group, and contemporaneous to major thermal activity/emplacements within the domes, collectively known as the Emu Pool Supersuite (e.g. Collins et al., 1998; Van Kranendonk et al., 2002; Sandiford et al., 2004). Van Kranendonk et al. (2002) reported that folds ascribed to this deformational event affected the 3335 Ma Euro Basalt and overlying Wyman Formation within greenstone belts adjacent to the Corunna Downs Complex, but not the <3230 Ma unconformably overlying Budjan Creek Formation (Soanesville Group; Hickman, 2012). The timing of this deformational episode is supported by ~3310 Ma metamorphic zircon ages (U-Pb SHRIMP; Nelson, 1999; François et al., 2014), and possibly ~3330 Ma hydrothermal zircon from the Duffer Formation (U-Pb SHRIMP; McNaughton et al., 1993).

47 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

The D4 sinistral faulting in NW direction affected the entire greenstone succession within the study area and presumably occurred in the mid-Archaean, as evident in the southward continuation, where rocks of the De Grey Superbasin are cut by the faults (satellite images). Within the study area, faulting was brittle in the Low-Strain Belt and ductile in the Central Fold

Belt (Fig. 2.18c), suggesting that during D4 the Central Fold Belt was still part of a deeper crustal level. The domains were then vertically juxtaposed during oblique strike-slip thrusting (D5) along reactivated S2 foliations and associated carbonate-altered shear zones, after emplacement of the gabbroic mega dyke (Fig. 2.18d). Hickman et al. (2010) ascribed the timing of reactivation of greenstone synclines within the East Pilbara Terrane to deposition of the Neoarchaean lower Fortescue Group. The Black Range Dyke swarm, ascribed to the early Fortescue Group, is interpreted to have formed during a phase of rifting, - after cratonization of the Pilbara Craton

(Hickman et al., 2010). Due to the absence of isoclinal F2 folds within the Low-Strain Belt of the study area, I infer that significant stratigraphy was lost during D5 thrusting. In contrast to partial convective overturn, horizontal plate tectonic models, including thrust stacking of tectonic slices, core complex formation, and fold interference, have been proposed to explain the development of dome-and-keel crustal geometry (e.g. van Haaften and White, 1998; Zegers et al., 1996; Blewett, 2002). Based on this study, I argue that the juxtaposing of different crustal levels comprising different lithological-structural inventories might be misleadingly interpreted as a result of plate tectonic related thrust stacking, as adjacent structural domains might appear as being ‘exotic’ (see van Haaften and White, 1998). However, thrusting within the East Pilbara Terrane, including loss of stratigraphy, as well as fold interference patterns, are mostly ascribed to late-stages in the East Pilbara Terrane evolution, when compressional boundary forces were very well dominant, particularly within the tightening greenstone synclines. As mentioned by Van Kranendonk et al. (2001), the extent of thrust displacement is important to imply horizontal plate tectonic processes in this regard. In this study, I conclusively demonstrate that the most extensive displacement due to thrusting with significant loss of stratigraphy occurred in the Neoarchaean, when plate tectonic boundary forces were indeed impacting on the ancient East Pilbara Terrane nucleus (e.g. Hickman et al., 2010). Moreover, 3D analogue models of Rayleigh- Taylor instabilities show that linear fold-and-thrust belts develop in the upper crust even above circular drips (Pysklywec and Cruden, 2004).

2.6.3 Implications for early Archaean tectonics

The local lithostratigraphic and structural interpretations can be accommodated within previous regional models of overall partial convective overturn as the main geodynamic mode responsible for dome-and-keel development (e.g. Collins et al., 1998). Firstly, the study supports the existence of an early phase of extension during deposition of greenstone rocks of the Warrawoona Group. Previous observations of early extensional structures within the Warrawoona Group elsewhere in the East Pilbara Terrane have been interpreted as the result of either: i) core complex formation (e.g. Zegers et al., 1996), ii) local caldera collapse

48 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

(Nijman et al., 1999), or iii) growth faulting during basin development (e.g. Nijman and deVries, 2001). As proposed in this study, syn-kinematic reworking of granitic rocks occurred during or subsequent to deposition of the Coongan Subgroup (i.e. Mount Ada Basalt, Duffer Formation), in agreement with previous geochronological data for rocks of the Callina and Tambina Supersuites, which have been identified as main components within the Muccan Granitic Complex (e.g. Van Kranendonk et al., 2006). The coeval nature of dome inflation and greenstone deposition precludes a core complex type setting (i). The main argument in favor of the caldera collapse (ii), or growth fault (iii) models has been the lack of geometrical relationship between the observed extensional structures and the current position of granitic domes (e.g. Nijman et al., 1999). However, this study demonstrates that the extensional faulting within the Doolena Gap greenstone belt occurred just prior to (f1A, Mount Ada Basalt) and possibly during (f1B, Duffer Formation) crustal uplift, followed by greenstone tilting and related deformation that I ascribe to early doming of the Muccan Granitic Complex. I therefore argue that upper crustal extension represented an early response to magmatic activity and associated doming within lower crustal levels. Secondly, this study demonstrates that within the Doolena Gap greenstone belt the most intense deformation and related tilting of stratigraphy occurred between deposition of the Duffer Formation and the Strelley Pool Formation. This observation contains two implications: i) major dome initiation within the Doolena Gap greenstone belt occurred prior to the main East Pilbara Terrane dome forming event at 3310 Ma (e.g. Collins et al., 1998; Van Kranendonk et al., 2002; Sandiford et al., 2004). The nature of the present study is very localized, but consistent with a previous regional study by Hickman (1984), in which early uplift of the most northern East Pilbara Terrane compared to the central and SE part of the East Pilbara Terrane has been proposed based on lateral facies and thickness variations. The cause of such differences in the craton-scale spatiotemporal pattern of crustal uplift, if in fact true, remains speculative: a possible scenario may be the lateral shift of mantle heat/plume sources. On the other hand, doming as late as 3020 Ma occurred also in the areas of proposed early uplift (Hickman and Van Kranendonk, 2004); ii) previous geochronological data constrains the deposition of the Duffer Formation to 3470 Ma (deVries et al., 2006), and the maximum age of Strelley Pool Formation deposition is 3427 Ma (Nelson, 1998). Thus, the early stage of dome-and-keel development in the Doolena Gap greenstone belt was accomplished within a time frame of <40 Ma. This relatively fast rate of gravity-driven crustal reorganization is in agreement with numerical models (e.g. Robin and Bailey, 2009; Thébaud and Rey, 2013). Robin and Bailey (2009) found that diapiric overturn of up to 60% of crustal volcanics can occur in as little as 10 Ma. Thébaud and Rey (2013) simulated a thermal anomaly due to insulation of radiogenic crust, buried beneath the 7.5 km-thick Kelly Group, promoting partial melting at its base. The partial melting initiated gravitational instabilities leading to rapid (~10 Ma) partial convective overturn. Johnson’s et al. (2014) simulations suggest overturn times as short as 5 Ma. This study also supports the notion that hydrochemical weakening of shear zones is critical for accommodating strain within the greenstone keel (e.g. Thébaud and Rey, 2013). de Bremond d’Ars et al. (1999) estimated numerically that the viscosity of greenstones must be reduced in

49 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______order to derive viscosity contrasts (dome-keel), which would result in the observed wavelengths (~50-100 km) of naturally occurring dome-and-keel structures. The authors showed that the required reduction of the viscosity of the keel compared to an average tonalitic-trondhjemitic- granodioritic basement could be achieved through secondary alteration. However, the greenstones must not be completely altered, otherwise the viscosity ratio to the basement would be too low, and could in turn only be produced through thermal anomalies, and thus softening, of the basement (de Bremond d’Ars et al., 1999; Sandiford et al., 2004; Thébaud and Rey, 2013). In the case of the Doolena Gap greenstone belt, I argue that the localization of alteration weakening provided an overall accommodation of strain, while inter-shear zone domains and lozenges retained sufficient density and strength to maintain the required viscosity ratio and gravitational instability across the dome-keel interface. In addition, softening of the basement is evident in my observations of migmatization and syn-kinematic leucosome formation within the Muccan Granitic Complex during the main D2 event, further promoting Rayleigh-Taylor flow. In contrast to models of thermal blanketing of radiogenic crust (Sandiford et al., 2004), or heating from below through upwelling mantle during lithospheric extension, the softening of the felsic mid-crust could also have been the result of volatile influx from a de-hydrating/-carbonating lower crust/sinking keel. Finally, the present observations agree with a subsequent, incremental development of the dome-and-keel structures, controlled by interplay between ascent of partial melts within domes and sagduction of greenstones. The progressive expansion of the domes resulted in increasingly complex strain patterns within the greenstone synclinoria.

2.7 Conclusions

This study reveals a complex multiphase structural history of the early Archaean Doolena Gap greenstone belt and adjacent Muccan Granitic Complex margin. I demonstrate that the western part of the Doolena Gap greenstone belt exhibits a deformational history (D1-D5) that spans much of the Archaean East Pilbara Terrane evolution and can be correlated with previously recognized East Pilbara Terrane-wide deformation events associated with partial convective overturn of the mid- to upper crust. The well-preserved stratigraphy in the Low-Strain Belt and the overall deformational overprinting relationships represent an exceptional opportunity to study some of the earliest structures and deformational mechanisms involved in the development of distinct dome- and-keel geometry in the East Pilbara Terrane. The evolution of the study area is summarized below: 1) Deposition of predominantly basaltic pillow flows and interlayered chert/BIF (Mount Ada Basalt), accompanied by west-block down listric normal faults, indicate volcanic outpourings in a deep marine, possibly E-W extensional basin. The abundance of mafic intrusions and sills, and evidence of proximal volcanic vents suggest intracrustal feeder systems.

50 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

2) Cessation of mafic volcanism, onset of felsic volcanism to the east and sedimentation to the west (Duffer Formation). Sedimentary facies and putative growth-faults indicate progressive E-W basin development. 3) Ongoing Duffer Formation sedimentation displays an overall regression towards shallower aqueous deposition (i.e. sandstone, stromatolites), implying local tectonic uplift, possibly related to dome component emplacement and associated amphibolite-facies contact metamorphism. 4) Subsequent southward tilting of Warrawoona Group beddings, and development of

transposed shear folds (F2) and localized fluid-induced normal faults within deeper crustal

levels (D2); increasing constrictional flow of keel rocks (sagduction) towards SW; mylonitic shear zone at dome interface; syn-anatectic shear folding of partial melts in dome. 5) Massive sandstone filling the sub-basin (Strelley Pool Formation), during erosion/non- deposition of the uppermost Warrawoona Group, providing evidence for sub-aerial weathering. 6) Further tilting and rotation towards dome triple point to the SW (after deposition of Kelly

Group?); associated D3 dextral dome-up shear deformation in keel. 7) NE-SW sinistral faulting, brittle in Low-Strain Belt, ductile in Central Fold Belt 8) Vertical juxtaposing of different crustal levels through dextral oblique slip along locally

possibly reactivated D2 foliations/faults, after dyke emplacement (2772 Ma).

51 Chapter 2 Lithostratigraphy and structure of the early Archaean Doolena Gap greenstone belt ______

52 “…Stones crumble and decay, Faiths grow old and they are forgotten, But new beliefs are born…” (Chief Joseph/Hinmatóowyalahtq̓ it, Nez Perce, 1840-1904)

Chapter 3 ______

FORMATION, DIFFERENTIATION AND REWORKING OF EARLY (ca. 3500-3590 Ma) EAST PILBARA TERRANE CONTINENTAL MATERIAL AND IMPLICATIONS FOR DOME-AND-KEEL INITIATION – EVIDENCE FROM THE WESTERN DOOLENA GAP GREENSTONE BELT

3.1 Abstract

New petrological, geochemical, and U-Pb zircon (LA-ICP-MS) geochronological data from a potentially petrogenetically related suite of intermediate to felsic gneisses along the southwestern margin of the domal Muccan Granitic Complex, Doolena Gap greenstone belt (East Pilbara Terrane) are presented. Upper Concordia intercepts yield U-Pb zircon ages of 3499 ± 22 Ma, 3576 ± 22 Ma, and 3591 ± 36 Ma, which are interpreted as the timing of emplacement of the gneisses’ protoliths. These Doolena Suite gneisses represent the oldest components of continental material discovered in the Muccan Granitic Complex and in the East Pilbara Terrane so far. Two out of 14 samples represent proper tonalitic-trondhjemitic-granodioritic (TTG) rock types that can be classified as medium- to high-Al TTG with typical garnet signatures of highly fractionated heavy rare earth elements (HREE; La/Yb ~35.2 and 51.0). Other samples comprise dioritic and granodioritic gneisses with higher HREE concentration levels and lower light rare earth element (LREE) fractionation (La/Yb ~6.2-13.5), and evolved potassic granitic gneisses (La/Yb(N) ~11.2-25.3). In combination with petrological and geochemical trace element modeling, a petrogenetic model is discussed, in which parental TTG magma derived from partial melting of deeply buried, hydrous garnet amphibolite, leaving a dense residue of garnet, clinopyroxene and minor ilmenite. Partial melting is explained due to over-thickening of a mafic proto-crust, along a minimum Archaean geothermal gradient, indicated by a modeled hydrous solidus at ~700 °C and ~13 kbar. Based on geochemical trace element modeling (Rayleigh fractionation), the dioritic and granodioritic components are suggested to represent cumulate fractions of a parental TTG magma that were emplaced at moderately great crustal depths (>8 kbar), as indicated by the presence of magmatic epidote.

Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Petrological and geochemical data from in situ syn-deformational leucosomes and late-stage pegmatites are evidence for eutectic granitic mineralogy and very low REE concentration levels, suggesting low-temperature partial anatexis of the host gneisses. It is argued that reworking of the 3499-3591 Ma Doolena Suite gneiss components along the margin of the Muccan Granitic Complex was the consequence of volatile influx from granitic magmatism during granitic doming. Additional U-Pb zircon provenance ages from a quartz-sandstone that overlies an angular unconformity within the adjacent greenstone belt are presented. Apparent 207Pb/206Pb ages of the detrital zircon confirm the stratigraphic correlation of the sandstone with the ~3427-3350 Ma Strelley Pool Formation. The data indicate that part of the pre-3500 Ma continental rocks were exposed and eroded until 3427 Ma. This implies that early dome-and-keel formation affected the study area until 3427 Ma, prior to the major regional dome-and-keel formation within the East Pilbara Terrane generally thought to have occurred at 3310 Ma. A model is discussed, in which the early Archaean onset and progressive increase in juvenile continental TTG magma production in the East Pilbara Terrane was a direct consequence of over- thickening of mafic proto-crust, as the lower crustal base was pushed down into a ‘zone of partial melting’ (ca. >45 km depth). Increasing infiltration of buoyant continental magma into vertically thickening stagnant proto-crust may have caused a change in the tectonic regime of the ancient East Pilbara Terrane crust to a ‘dynamic lid’ mode characterized by gravitational re-organization and dome-and-keel formation.

Keywords: East Pilbara Terrane, early Archaean, TTG, dome-and-keel tectonics, U-Pb zircon geochronology, Muccan Granitic Complex, Strelley Pool Formation;

54 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

3.2 Introduction

Juvenile continental crustal growth during the early Archaean (3850-3200 Ma) supposedly was both voluminous and rapid (e.g. review in Kamber, 2015). Preserved early Archaean terranes comprise characteristic granite dome – greenstone keel crustal structures that represent the manifestations of distinct crustal growth within a hotter early Earth (e.g. Moore and Webb, 2013; Kamber, 2015; Van Kranendonk et al., 2015). Importantly, the processes of formation of the preserved early Archaean crust must be distinct to that which occurred both in the Hadean (>3850 Ma), which has little to no preservation, and in the later Archean to the present day when definitive plate tectonic process dominated crust formation (e.g. Van Kranendonk et al., 2010; Kamber, 2015). The dome-and-keel structures of early Archaean terranes are interpreted to be the result of partial convective overturn of the mid- to upper crust, as a result of gravitational instabilities arising from a buoyant felsic mid-crust, overlain by dense mafic volcanic rocks (e.g. Collins et al., 1998). As mentioned by Sizova et al. (2015), previous models of partial convective overturn suppose the pre-existence of a felsic mid-crustal layer, without explaining the generation of the early felsic melts (e.g. Sandiford et al., 2004; Thébaud and Rey, 2013). It has been suggested that mafic proto-crust rapidly thickened through massive, mantle-derived volcanic resurfacing events resulting in Archaean-type oceanic plateau (e.g. Kamber et al., 2005; Kamber, 2015; Van Kranendonk et al., 2015). Models for the formation of felsic melts from over-thickened mafic crust include: i) polybaric infra-crustal differentiation of the mafic crust, (e.g. Smithies et al., 2009), ii) partial melting at the base of the lower crust (e.g. Clemens et al., 2006; Reimink et al., 2016), and iii) partial melt production as a result of interaction of delaminating lower crustal eclogite and upwelling astenospheric upper mantle (e.g. Bédard, 2006; Johnson et al., 2014). In the 3530-3225 Ma East Pilbara Terrane of Western Australia, voluminous continental material that forms granitic dome complexes, surrounded by synclinal greenstone keels, records the production of tonalitic-trondhjemitic-granodioritic (TTG) and associated granitic magma of the well-preserved 300 Ma history of the East Pilbara Terrane evolution (e.g. Van Kranendonk et al., 2006; 2015; Hickman and Van Kranendonk, 2012). Continental rocks older than 3500 Ma, however, are extremely sparse and the presence of extensive pre-East Pilbara Terrane continental basement is largely inferred from 3800-3500 Ma inherited and detrital zircon (e.g. Thorpe et al., 1992; Nelson, 2000; Kemp et al., 2015), and very rare ~3660 and 3580 Ma gneissic xenoliths (e.g. McNaughton et al., 1988; Hickman and Van Kranendonk, 2012). Geochemical and isotopic studies have shown that within the 3530-3225 Ma East Pilbara Terrane record, only the oldest ~3490-3460 Ma Callina and 3450-3420 Ma Tambina supersuite TTG represent juvenile crustal magma, while younger TTG and granitic rocks were formed through reworking of older TTG (e.g. Collins, 1993; Smithies et al., 2003; Van Kranendonk et al., 2015). In order to establish a more advanced model of the proposed rapid early Archaean-style crustal growth of the East Pilbara Terrane, two fundamental questions need to be addressed: i) How did the earliest granitoid melts form? and ii) when and why did the evolving crust ultimately collapsed in a gravitational re-

55 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.1: a) Simplified geological map of the East Pilbara Terrane showing the distribution of granitic supersuites within the domal complexes (modified after Van Kranendonk et al., 2006); NP, North Pole; WA, Warrawagine; YG, Yilgalong; study area indicated; b) Location of the East Pilbara Terrane in Western Australia. organization of up-doming granite and sagduction of mafic greenstone-keel rocks? Addressing these questions will contribute to understanding the tectono-thermal processes and timeframes of early Archaean crust formation and evolution in the East Pilbara Terrane. Solutions to both questions are highly controversial because of the following: i) It is generally accepted that Archaean TTGs were derived through partial melting of deeply buried hydrated basalt (i.e. amphibolite; e.g. Foley et al., 2002; Rapp et al., 2003; Smithies et al., 2009; Moyen and Martin, 2012; Nagel et al., 2012; Polat, 2012; Martin et al., 2014). Based on the presence of both high-Al (i.e. high pressure) and low-Al (i.e. low-pressure; e.g. Moyen and Martin, 2012) 3490-3420 Ma TTG components in the East Pilbara Terrane, it has been argued that TTG melt production in the East Pilbara Terrane occurred over various depths and were due to various degrees of partial melting within the crust (Smithies et al., 2009). The assumption that the mafic proto-crust had to develop sufficient thickness until voluminous high- to medium-pressure TTG could be produced, led to the suggestion that pre-3500 Ma East Pilbara Terrane TTG were likely characterized by low-Al compositions and formed at lower crustal depth (e.g. Smithies et al., 2009; Van Kranendonk et al., 2015). The model stands in contrast to models of TTG derivation

56 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______where geochemical trace element signatures of TTG require the presence of rutile-bearing eclogite residue, implying partial melting at greater depth (e.g. Rapp et al., 2003; Bédard, 2006). Also, the derivation of Archaean TTG compositional varieties has been explained either by fractional crystallization (Martin, 1987; Bédard, 2006), or by partial melting at different depths (e.g. Smithies et al., 2009; Moyen and Martin, 2012). ii) The findings in Chapter 2 suggest the presence of a suite of spatially associated granitoid gneisses along the margin of the Muccan Granitic Complex (Doolena Gap greenstone belt, East Pilbara Terrane) that were possibly partially reworked during a partial convective overturn event that occurred prior to the formation of the unconformity below the quartz-sandstone. If the sandstone can be conclusively ascribed to the regional 3427-3350 Ma Strelley Pool Formation (Hickman, 2008; Van Kranendonk, 2010), these observations suggest that i) the protoliths of the host gneisses must be older than 3427 Ma, making them some of the oldest continental rocks in the East Pilbara Terrane and ii) the gneisses were affected by partial reworking prior to 3427 Ma and therefore prior to the regional main event of partial convective overturn (~3310 Ma; e.g. Collins, 1993; Smithies et al., 2009; Van Kranendonk et al., 2015). Petrological-geochemical understanding of these rocks will have implications both for early continental crust formation in the East Pilbara Terrane, and for reworking during the earliest recognized partial convective overturn event and associated dome-and-keel development. Here, I present U-Pb zircon (LA-ICP-MS) geochronological and petrological-geochemical data from this potentially ancient gneiss suite (Muccan Granitic Complex, Doolena Gap greenstone belt) to [1] determine the timing of protolith emplacement, and [2] discuss their petrogenesis by testing proposed contrasting models of Archaean TTG formation. [3] Geochemical and petrological data from selected syn-deformational leucosomes is aimed to constrain the mechanism of partial reworking with implications for early dome-and-keel initiation. Finally, [4] I present U-Pb data of detrital zircon from the quartz-sandstone to conclusively test the stratigraphic correlation with the Strelley Pool Formation, and hence the antiquity of the early partial convective overturn event (Chapter 2). The detrital zircon ages will have implications for provenance and un-roofing of ancient East Pilbara Terrane continental material during dome-and- keel formation.

3.3 Brief Geology and Tectonic Model of the East Pilbara Terrane

The 3530-3225 Ma East Pilbara Terrane forms the ancient granite dome - greenstone keel nucleus of the Archaean Pilbara Craton (e.g. Hickman and Van Kranendonk, 2012). The ca. 60 km diameter dome-forming granitic complexes of the East Pilbara Terrane each comprise multiple magmatic components that are ascribed to four major granitic supersuites: the 3490-3460 Ma Callina, the 3450-3420 Ma Tambina, the 3324-3290 Ma Emu Pool, and the 3275-3225 Ma Cleveland supersuites (Fig. 3.1; Van Kranendonk et al., 2006). Within the domes, older magmatic components are mostly found along the margins, while younger components form the dome centers (e.g. Van Kranendonk et al., 2006). The granitic supersuites mostly comprise TTG, but

57 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.2: Simplified geological map of the study area showing locations of samples used in this study; all samples including brief description and GPS coordinates are listed in Table 3.1. potassic granites are also observed, even in the oldest Callina supersuite (Van Kranendonk et al., 2006). The granitic domes are interleaved by synclinal volcano-sedimentary successions (i.e. greenstone keels) of the 3530-3225 Ma Pilbara Supergroup (Fig. 3.1). The Pilbara Supergroup comprises numerous autochthonous ultra-mafic – mafic to felsic volcanic cycles dominated by tholeiitic pillow basalts (e.g. Smithies et al., 2007). Major felsic volcanic formations at the top of the volcanic cycles coincide with the granitic supersuites (Van Kranendonk et al., 2006). Direct petrogenetic relationships between the felsic volcanics and the TTGs, however, are not established until ~3420 Ma, and most of the volcanic cycles are interpreted as tholeiitic differentiation series (e.g. Smithies et al., 2007). The volcanic cycles of the Pilbara Supergroup are interpreted as the result of successive plume volcanism (e.g. Smithies et al., 2007; Hickman and Van Kranendonk, 2012). Major regional dome-and-keel formation (partial convective overturn) at 3310 Ma supposedly initiated by partial melting of mid-crustal granitic/TTG material, causing buoyant rise of the partial melts as magmatic domes, and providing a soft, ductile substrate into which the overlying dense supracrustal rocks sagducted to form the greenstone keels (Collins et al., 1998; Sandiford et

58 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.3: Field photographs of gneisses from the Muccan Granitic Complex; a) typical outcrop of granitoid gneisses along the southwestern margin of the Muccan Granitic Complex (location DW-327; view to partial convective overturn, outcrop ~6 m across from left to right of image); b) migmatitic tonalite-trondhjemite gneiss boulder showing tightly folded leucosomes and leucosome injections, parallel to the host gneiss foliation; it is shown that a foliated leucosome vein terminates in a larger partial melt pocket (front of image); the melt pocket is less deformed and more coarse-grained (i.e. pegmatitic), and cross-cuts the host gneiss foliation; this possibly indicates a relationship between incipient in-situ syn-tectonic partial melt production, melt migration and injection, and late-stage pegmatite intrusion (ca. 50 m WNW of location DW-309; thickness of hammer head ~2.5 cm); c)

migmatitic tonalite-trondhjemite gneiss boulder showing evidence for syn-tectonic (i.e. syn-D2-folding; Chapter 2) in-situ leucosome formation; the darker segregates between the leucocratic layers are interpreted to represent restites/melanosomes (ca. 50 m WNW of location DW-309; thickest part of hammer head ~2.5 cm); d) syn-tectonic generations of tightly folded leucosomes cross-cutting tonalitic- trondhjemitic host gneiss; high volumes of leucosomes indicate increasing accumulation of partial melts, possibly causing buoyant magmatic ascend of the Muccan Granitic Complex dome (location DW-327; hammer length ~33 cm); refer to text in Discussion section. al., 2004; Van Kranendonk et al., 2015). The existence of pre-East Pilbara Terrane (>3530 Ma) crust of partly continental nature is based on following evidence: i) detrital and inherited zircon with ages up to 3800 Ma (e.g. Thorpe et al., 1992; Nelson, 2000; Kemp et al., 2015), ii) bulk-rock and zircon isotopic data suggesting crustal extraction as early as 4000 Ma (Tessalina et al., 2010), but mostly between ca. 3500 and 3700 Ma (e.g. Hf-zircon, Kemp et al., 2015; Nd-bulk-rock, Smithies et al., 2003, 2009), iii) crustal contamination of even the oldest (3530 Ma) greenstone volcanic rocks (e.g. Green et al., 2000), and iv) rare xenoliths of tonalitic gneiss and anorthositic gabbro up to 3660 Ma within younger granitic gneiss (McNaughton et al., 1988; Hickman and Van Kranendonk, 2012). This set of evidence led to the interpretation that the East Pilbara Terrane formed as a plume-derived volcanic plateau on top of a precursor (>3500 Ma) partly continental basement (e.g. Hickman and Van Kranendonk, 2012; Van Kranendonk et al., 2015).

59 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

3.4 Study Area and Sample Selection

The samples investigated in the present contribution come from the 1 Mile Creek granite dome - greenstone keel traverse that extends from the southwestern margin of the Muccan Granitic Complex to the north into the southerly adjacent Doolena Gap greenstone belt (Fig. 3.1; Fig. 3.2). Within the Muccan Granitic Complex rock exposure is poor and contact relationships between different gneissic components largely unclear. However, at least two relatively older components are identified within larger outcrops (Fig. 3.3a), including a hornblende-biotite dioritic to granodioritic gneiss, and a biotite tonalitic-trondhjemitic gneiss (Fig. 3.3b,c). The gneisses show evidence for migmatization and subsequent in situ syn-deformational partial anatexis (Chapter 2; Fig. 3.3b,c). Relatively younger components, beside the in situ leucosomes, form both foliated and undeformed leucocratic veins and intrusions, and pegmatite bodies (Fig. 3.3a,d). The tight folding of the leucosomes (Fig. 3.3c) resembles tight to isoclinal folds within the mafic schists of the greenstone belt (Central Fold Belt; Fig. 3.2). In Chapter 2 it is suggested that the partial anatexis of the gneisses and the deformation in the greenstones are associated with early dome-and-keel development. Due to the absence of related structural features in the quartz-sandstone (Fig. 3.2), the latter event occurred prior to the formation of the unconformity, at 3427 Ma if the quartz- sandstone can be correlated with the Strelley Pool Formation. From within the Muccan Granitic Complex and adjacent South Muccan Shear Zone, 21 samples were collected for petrological and geochemical characterization (Fig. 3.2; Table 3.1). Only least altered samples were taken and secondary quartz and carbonate veins avoided. The samples were preliminarily grouped based on their mineralogy into i) diorite and granodiorite, in which hornblende is the dominant mafic mineral, including three gneisses from the proper Muccan Granitic Complex and two mylonites from the South Muccan Shear Zone, ii) tonalitic- trondhjemitic and granitic gneisses from the Muccan Granitic Complex (nine samples), in which biotite is the dominant mafic mineral, and iii) both in situ leucosomes and other leucocratic veins and patches, and pegmatite bodies (seven samples). Restitic domains and schlieren (i.e. melanosomes) are observed in the field adjacent to the in situ leucocratic patches (Fig. 3.3c), but could not be readily sampled due to their relatively small size. Sample locations are indicated in the map (Fig. 3.2). Leucosomes that were directly separated from corresponding host gneisses are indicated with ‘B’ following the sample number in Table 3.1. A number of gneisses were selected for U-Pb zircon geochronology, but insufficient numbers of zircon were extracted for dating from most of the TTG and granites, particularly the leucosomes. Nevertheless, two dioritic and granodioritic gneisses and two TTG-granite gneisses did contain sufficient zircon (Table 3.1). Finally, two quartzitic sandstone samples from the presumed Strelley Pool Formation were collected for the purpose of U-Pb zircon dating (Fig. 3.2; Table 3.1). One sample represents an poorly sorted immature quartz-jasper conglomerate and comes from the lowermost part of the sandstone package, and the other sample represents a well- sorted mature quartz-sandstone from higher up in the stratigraphy (Fig. 3.2).

60 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

3.5 Analytical Methods

Thin sections of all samples were prepared, where present perpendicular to the foliation plane. The samples were petrographically examined with a Leica DM750P polarizing microscope, fitted with a camera system (Leica ICC50HD) for image acquisition. Where given, modal abundances of

61 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______mineral phases represent average modes estimated by the eye from domains of preserved magmatic textures. Because of the gneissic and/or migmatitic layering (excluding leucosomes), sufficient amounts of samples that are representative of the bulk composite rock were used for geochemical sample preparation. Distinct leucosomes and pegmatites were separated from their host rock. The samples were broken down with a sledgehammer to gravel size within heavy-duty sample bags to avoid contamination. The samples were then crushed to sand size (~500 µm) with a tungsten-free rock crusher, and powdered using an agate mill. For major element X-ray fluorescence (XRF) analyses, the powdered samples were first ignited for ca. 3 hours in ceramic crucibles at 950 °C to measure loss on ignition (LOI) by weight. XRF discs were fused with a multi-sample Claisse OX fuser from a mixture of 0.87 g of ignited sample powder and 6.67 g of a 1:1 lithium metaborate and lithium bromide flux. WD-XRF analyses were performed with a Pananalytical Axois 1 kW system fitted with a Rh anode X-ray tube, LIF200, PE20mm, PX1 and LIF220 crystals, a Duplex and a Flow detector, at the Central Analytical Research Facility, Queensland University of Technology. Bulk-rock USGS (AGV-2 281, BCR-2; Wilson, 1997; 1998) and SARM (1, 2, and 6; Mintek, SA, Johannesburg, SARM certificates, www.mintek.co.za) standards were analyzed repeatedly to correct for instrumental drift throughout the run. XRF major element detection limits are provided in Table 3.A-I (Appendix 7). For trace element analyses 50 mg of sample were dissolved in a mixture of 1.5 ml double distilled HCl, 0.5 ml double distilled HNO3 and 0.5 ml double distilled HF in PTFE vials using a Milestone bench-top Ultra-wave single reaction chamber microwave digestion system. The samples were then evaporated to incipient dryness in Teflon beakers on a hot plate at 80 °C, followed by adding of 2 ml double distilled HCl and 1 ml MilliQ water and leaving the capped samples on a hot plate at 100 °C overnight, and in open beakers until dry. Further evaporation at

100 °C was done after adding 2 ml double distilled HNO3, and again after adding 1 ml MilliQ and

1 ml double distilled HNO3. The samples were then refluxed in 0.2 ml HNO3 and 4.8 ml MilliQ at 100 °C overnight. After repeated washing with MilliQ, the solution was centrifuged in capped tubes at 35 RPM for 15 min. The samples were weighed to calculate a dilution factor. Aliquots of the solution were transferred to clean tubes, internal standards added, and weighed to determine a second dilution factor. Analyses were performed with an Agilent 8800 triple quadrupole ICP-MS at the Central Analytical Research Facility, Queensland University of Technology. Detection limits for ICP-MS trace elements are provided in the Table 3.A-II (Appendix 8). Zircon separation was performed by commonly used techniques including sledgehammer and tungsten disc mill (<500 µm), washing, drying, Frantz magnetic barrier separator (>1.5 or 1.8 A, depending on mafic content), and heavy-liquid separation with methylene iodide at the University of Queensland rock lab. Single grains were handpicked under a binocular light microscope and mounted in epoxy. Zircon was characterized by cathodoluminescence (CL) imaging using a ZEISS Sigma field emission scanning electron microscope (SEM) in variable pressure (VP) operation mode at 20 kV and 1.2 nA with a scanning speed of 9. CL imaging of zircon from both the orthogneisses and the sedimentary rocks revealed faint concentric oscillatory growth zoning of

62 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.4: see next page for Figure caption.

alternating light- to medium and medium- to dark CL intensity, indicating a magmatic origin. Zircon from the dioritic gneiss 370-1 shows thin outer rims of medium CL intensity, truncating the inner growth structures. Only the oscillatory-zoned zircon cores were targeted for U-Pb and selected trace element isotopic analyses. The grains were ablated on selected spots of 30 µm- diameter size using a 193 nm excimer laser. Ablated material was transported in a He carrier gas and analyzed with an Agilent 8800 triple quadrupole ICP-MS at the Central Analytical research Facility, Queensland University of Technology. Zircon standards Temora-2 (416.78 ± 0.33 Ma; Black et al., 2004) and Plešovice (337.13 ± 0.37 Ma; Sláma et al., 2008) were used as primary and secondary reference material, respectively, and analyzed repeatedly after every sequence of 10 unknowns. The NIST 610 glass (Jochum et al., 2011) was used as a primary standard to determine trace element compositions. The LA-ICP-MS data reduction (i.e. baseline correction, integration of zircon standards and down-hole fractionation correction) and propagated error calculation was

63 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.4: Photomicrographs showing representative mineralogy and texture of various gneiss components (Muccan Granitic Complex); a) dioritic gneiss 370-1, showing magmatic textures in the upper right half of the image; eu- to subhedral coarse-grained hornblende displays a faint color change from brownish- green in the cores to green at the margins; the margins comprise fine-grained anhedral apatite inclusions; hornblende forms an interlocking texture with coarse-grained eu- to subhedral plagioclase and medium- grained titanite; areas of interstitial quartz are recrystallized to subgrains; a deformation band runs through the centre of the image (upper left to the bottom right), displaying a foliation defined by the sub- parallel alignment of fine- to medium-grained, elongated anhedral hornblende and biotite laths, and sheared and recrystallized elongated quartz and plagioclase; the deformation band is rich in finely recrystallized anhedral titanite (plane-polarized light); b) dioritic gneiss 370-1, showing medium-grained eu- to subhedral magmatic epidote aligned within the foliation defined by biotite laths; biotite is partly replaced by chlorite; plagioclase forms anhedral medium-grained porphyroclasts in zones of intensive deformation; apatite occurs as inclusions or intergrowths with patchy biotite, and forms fine- to medium- sized, eu- to subhedral grains; interstitial quartz is always recrystallized (cross-polarized light); c) tonalitic-trondhjemitic gneiss 327-3, showing coarse-grained plagioclase porphyroclasts clouded by abundant very fine-grained opaque micro-inclusions; where plagioclase is recrystallized at its rims the opaque clouding is not observed; some plagioclase porphyroclasts display intensive sericitization (e.g. bottom right of image, not visible in plane-polarized light); fine- to medium-grained anhedral biotite laths are almost entirely replaced by chlorite, forming elongated schlieren surrounding the plagioclase porphyroclasts; the sub-parallel alignment of the mafic schlieren with rotated plagioclase porphyroclasts and elongated patches of recrystallized quartz subgrains define the gneissic foliation; magmatic epidote forms interlocking textures withplagioclase; epidote is partly replaced by chlorite (plane-polarized light); d) granitic proto-mylonite 310-3, displaying medium-grained magmatic remnants of plagioclase and minor alkali feldspar porphyroclasts within a matrix of fine- to medium-grained quartz, alkali feldspar and plagioclase representing dynamic recrystallization during ductile shear deformation (see also Chapter 2); the more quartzofeldspathic layers are segregated from very thin schlieren of elongated fine-grained chlorite laths replacing former biotite (cross-polarized light); e) pegmatite 308-2, showing graphic intergrowth of quartz with very coarse-grained feldspar (Fsp); the feldspar is interpreted to be of an albite-rich ternary composition, as antiperthitic exsolution of microcline within plagioclase host are common (not shown); the lamella in the image either represent antiperthitic K-feldspar exsolution lamella, or albite-twin lamella (cross-polarized light; note the yellow interference color of quartz due to ca. 35-40 µm thin section thickness); f) leucocratic domain from granitic migmatite gneiss 327-1, showing quartz myrmekite in plagioclase in association with alkali feldspar; compared to the magmatic plagioclase clouded by opaque micro-inclusions found in the host gneisses (e.g. Fig. 3.4c), feldspars in the leucocratic domains of migmatites are clear, and free of opaque inclusions, but display tracks of other inclusions (quartz or fluid?; see arrow); the observations suggest syn-tectonic metasomatism and/or partial melting (cross-polarized light); mineral abbreviations are given below Table 3.1; sample numbers correspond to locations in Figure 3.2 and descriptions and GPS coordinates in Table 3.1. performed in the Igor Pro-based Iolite software (WaveMetrics). For U-Pb Concordia diagrams the Excel-based Isoplot add-in was used.

3.6 Petrography

Besides some of the pegmatites, all samples comprise domains characterized by a gneissic foliation. Often domains of varying deformational intensity are present within a single rock, as indicated by different degrees and mechanisms of dynamic recrystallization of quartz and feldspar (see Chapter 2). Most of the samples show greenschist-facies retrogression evident in the partial replacement of mafic phases (hornblende, biotite) by chlorite ± epidote, and the sericitization of plagioclase. In the following sections, the primary magmatic mineralogy and preserved magmatic textures are described to infer magmatic processes. Sample locations are shown in the map of Figure 3.2, and mineral assemblages and GPS locations are given in Table 3.1.

64 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

3.6.1 Dioritic to granodioritic gneisses

The dioritic and granodioritic gneisses are distinguished from tonalitic-trondhjemitic and granitic gneisses by the presence of abundant hornblende (Table 3.1). The samples from the Muccan Granitic Complex (Fig. 3.2) preserve domains characterized by a relatively undeformed coarse- grained magmatic texture, comprising a mineral assemblage of hornblende, plagioclase, biotite, quartz, minor epidote and titanite, and accessory, though relatively abundant apatite and zircon (Fig. 3.4a,b). The gneissic foliation is mainly defined by bands of biotite laths with shape- preferred orientation, and anhedral hornblende and titanite, incorporating recrystallized anhedral plagioclase and elongated patches of quartz sub-grains in the foliation (Fig. 3.4a,b). The magmatic domains comprise a phaneritic interlocking texture of dominantly coarse- grained eu- to subhedral hornblende and plagioclase (Fig. 3.4a). Both hornblende and plagioclase are present in similar abundance of approx. 35-40 Vol. %, each. Granodiorites are distinguished from diorites by relatively more modal abundance of biotite and slightly less abundant hornblende. The coarse-grained euhedral hornblende shows green to slightly brownish pleochroism, with a compositional zoning indicated by brownish-green cores towards green rims (Fig. 3.4a). Within the margins, hornblende often contains fine-grained inclusions of anhedral apatite (Fig. 3.4a). Along the grain boundaries, hornblende is often associated with titanite. Within the deformed bands, hornblende is anhedral and aligned in the gneissic foliation, associated with abundant foliation-defining biotite, and partly replaced by secondary chlorite. Within the interstices of interlocking hornblende and plagioclase, patches of quartz are recrystallized by bulging and subgrain formation (Fig. 3.4a,b). The coarse-grained eu- to subhedral plagioclase is characterized by abundant polysynthetic twinning, and shows areas of abundant micro-inclusions of opaque phases, resulting in a dark shading in plane-polarized light (Fig. 3.4a). Where plagioclase is partly recrystallized, it lacks the dark shading and the polysynthetic twinning and forms anhedral medium-sized grains (Fig. 3.4a). Apatite within the matrix is slightly larger (up to ~500 µm) and sub- to euhedral, compared to the inclusions in the hornblende margins. The matrix apatite is often overgrown by, and or intergrown with, either coarse-grained patchy biotite in undeformed domains, or medium-grained biotite laths within the foliation-defining deformation bands (Fig. 3.4b). Titanite forms medium-sized grains associated with hornblende in magmatic domains (Fig. 3.4a), and, either fine-grained or medium-sized elongated grains within the foliation-defining deformation bands (Fig. 3.4a,b). Magmatic eu- to subhedral epidote is observed (Fig. 3.4b). Zircon is abundant and visible in thin section (>100 µm; Fig. 3.4a,b), and occurs in the matrix or as inclusions in biotite and hornblende.

3.6.2 Tonalitic-trondhjemitic and granitic gneisses

The tonalitic-trondhjemitic and granitic gneisses comprise biotite as the primary mafic phase and show varying modal abundances of K-feldspar. K-feldspar is rare in the TTG gneisses (<5 Vol.%), but reaches minor to major abundances in some of the granites (ca. 5-15 Vol.%). The variation in

65 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

K-feldspar is transitional, and the two rock types cannot be readily distinguished by petrographic observations. The same seems to be true for biotite, which shows a transitional negative correlation of abundance with increasing K-feldspar content. Quantification of the bulk-rock mineral modal abundance based on petrographic observations, however, is hampered by the fact that the tonalitic-trondhjemitic and granitic gneisses are significantly affected by different degrees of deformation and recrystallization. The rocks often comprise a heterogeneous distribution of mineralogy and deformation intensity dependent grain-size and recrystallization mechanisms, which mesoscopically reflect the gneissic banding and incipient migmatization (Fig. 3.3b,c; Fig. 3.4c,d). Nevertheless, remnant primary magmatic mineralogy and textures are observed (Fig. 3.4c). Often, sub- to euhedral plagioclase porphyroclasts show simple or polysynthetic twinning (e.g. Fig. 3.4d). Within the South Muccan Shear Zone (see map in Fig. 3.2), granitic mylonites have been observed that comprise porphyroclastic plagioclase characterized by a chemical zoning (not shown). Mostly, plagioclase porphyroclasts form medium- to coarse-sized anhedral grains indicating dynamic recrystallization at the grain boundaries during deformation (Fig. 3.4c,d). In some samples, plagioclase porphyroclasts form elongated augen within a highly shear-deformed and recrystallized matrix of quartz, K-feldspar, biotite and minor plagioclase. Similar to the plagioclase in the diorites and granodiorites, preserved magmatic cores of the plagioclase porphyroclasts are clouded by very fine-grained opaque micro-inclusions (Fig. 3.4c). K-feldspar forms mostly anhedral medium-sized grains within matrix interstices formed by plagioclase, or as porphyroclasts within highly deformed quartzofeldspathic domains (Fig. 3.4d). K-feldspar often shows typical microcline cross-hatch twinning (Fig. 3.4d). Quartz is always anhedral and recrystallized (Fig. 3.4d), often forming subgrains within interstitial patches (Fig. 3.4c). Biotite occurs as fine- to medium grained elongated laths forming thin schlieren in between the quartzofeldspathic layers of rotated plagioclase porphyroclasts and elongated patches of quartz subgrains (Fig. 3.4c). The petrographic distinction between magmatic and metamorphic epidote is not clear. Some magmatic epidote in the tonalitic-trondhjemitic gneisses is inferred by the eu- to subhedral shape and preferred occurrence in association with magmatic plagioclase, in absence of replacement textures (e.g. Fig. 3.4c). On the other hand, the occurrence of anhedral epidote in association with plagioclase, and/or biotite and metamorphic chlorite within highly deformed granitic gneisses rather indicates a metamorphic or metasomatic growth of epidote. Accessory phases within the tonalitic-trondhjemitic and granitic gneisses are less abundant and more fine- grained compared to the diorites and granodiorites. However, significant amounts of apatite (Fig. 3.4c) and titanite, and few zircon have been observed in the tonalitie-trondhjemites, but overall the trace phases decrease in abundance in the K-feldspar enriched granites.

3.6.3 In situ leucosomes, late-stage pegmatites and leucocratic domains of migmatites

The in situ leucosomes and pegmatites are overall less deformed compared to their host gneisses, or not deformed at all, often preserving pristine magmatic textures and mineralogy. The pegmatites are very coarse-grained (Fig. 3.4e), while the leucosomes are medium- to coarse-

66 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

grained, and both rock types display interlocking magmatic textures. Both pegmatites and leucosomes comprise an assemblage consisting of dominantly quartz (~30-40 Vol.%) and feldspar (~60-70 Vol.%) ± minor biotite (Table 3.1). In fact, as inferred from field observations (Fig. 3.3b), leucosomes and pegmatites show mineralogical similarities indicating a petrogenetic relationship, with differences depending on host rock composition and the syn- (in-situ leucosomes) to post- (pegmatites) deformational intrusive history (Fig. 3.3b). The modal abundance of the primary major phases reflects a eutectic granitic composition. The feldspars are mostly aniperthitic, showing patchy exsolution textures of microcline K-feldspar within coarse-grained sodic plagioclase hosts. The modal abundance of exsolved alkali and sodic feldspars has not been

67 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______estimated. A eutectic composition is supported by the presence of graphic intergrowths of quartz rods with the composite ternary feldspar hosts (Fig. 3.4e). The leucosomes and pegmatites almost completely lack any accessory phases. Some samples contain fine- to medium-grained anhedral titanite, and other samples comprise accessory fine- to medium-grained euhedral opaque phases, possibly ilmenite or magnetite. Within quartzofeldspathic domains of the migmatitic host gneisses, quartz myrmekites are observed within medium- to coarse-grained plagioclase feldspars and associated with eu- to subhedral medium-grained K-feldspar (Fig. 3.4f). Besides the higher abundance of mafic schlieren, the overall mode of quartz and feldspar within these leucocratic domains resembles the eutectic mineralogy of the leucosomes and pegmatites. Medium-grained quartz, K-feldspar and plagioclase are surrounded by a fine- to medium-grained recrystallized matrix of interstitial quartz and feldspar. In contrast to the opaque clouding of magmatic plagioclase in less deformed dioritic, granodioritic and tonalitic-trondhjemitic host rock domains, plagioclase is clear (plane-polarized light), but contains tracks of other inclusions, likely represented by fluids or quartz (Fig. 3.4f).

3.7 Geochemistry

In order to infer petrogenetic processes and relationships between the studied rock types, the samples are geochemically characterized, first based on major element composition, and then based on trace element concentration. Major element XRF data are given in Table 3.2 and trace element ICP-MS data are given in Table 3.3. For trace element normalization (indicated by N in the text) the C1 chondrite reference values from McDonough and Sun (1995) were used.

3.7.1 Major elements

The major element data supports the preliminary subdivision of the samples based on their mineralogy into dioritic-granodioritic and more evolved tonalitic-trondhjemitic and granitic rocks by a conspicuous silica content gap between ca. 65 and 71 wt.% SiO2, shown in the Harker variation diagrams of Fig. 3.5a-h. SiO2 contents overall range from ~58 to 76 wt.% in the host gneisses, and reach up to 78 wt.% in the leucosomes/pegmatites. All samples plot along a calc- alkaline trend in the AFM diagram (Fig. 3.6a). However, in the modal Na-K-Ca ternary plot (Fig.

3.6b) only the more primitive diorites-granodiorites and tonalite-trondhjemites with SiO2 wt.% content below ca. 72 wt.% plot along a classical TTG evolution trend (e.g. Martin, 1993), while the more evolved granites show increasing K-enrichment with increasing SiO2 not typical for proper TTG-type rocks (Fig. 3.6b; e.g. Moyen and Martin, 2012). While some degree of alkali metal ion (e.g. K+, Na+) mobility during alteration cannot be excluded, the correlation between increasing K2O content with the petrographic observation of increasing modal abundance in K- feldspar, and minor scatter of decreasing Na2O and increasing K2O variation with increasing silica content (Fig. 3.5e,g) suggest relatively pristine compositions. This is supported by low to very low LOI values of 0.16-1.81 (Table 3.2). Therefore, it is argued that the transitional potassic calk-

68 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.5: Harker variation diagrams for major element oxides (a-h) using SiO2 as an index of differentiation. alkaline trend (Fig. 3.6b) represents the original chemical-mineralogical evolution of the granites, as minor alteration related mobility is not expected to affect the molecular Na-K-Ca plot significantly. In the following text, the more evolved granitic gneisses will be referred to as potassic granites to distinguish them from the tonalitic-trondhjemitic rocks.

The dioritic and granodioritic gneisses range in SiO2 content from ~58 to 64 wt.%. In the major element Harker variation diagrams (Fig. 3.5) the higher MgO, Fe2O3, CaO, P2O5 and TiO2 plot roughly on a linear trend line with the tonalite-trondhjemite and potassic granites. Al2O3, K2O and

Na2O, on the other hand, display reverse trends, i.e. very slightly decreasing K2O and slightly increasing Na2O and Al2O3 with increasing SiO2 content, compared to the tonalite-trondhjemite and potassic granites. The drastic increase in K2O (Fig. 3.5e) with corresponding decrease in Na2O (Fig. 3.5g) within the tonalite-trondhjemite - granite series mark the observed divergence from the

69 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

sodic trondhjemitic trend towards the potassic calc-alkaline trend in the Na-K-Ca ternary plot (Fig. 3.6b). MgO content and Mg# for the diorites and granodiorites range from 2.36 to 3.26 wt.% and 33 to 31, respectively, showing a systematic decrease with increasing SiO2. The most primitive tonalite-trondhjemite rocks have Mg# of 14 to 19, which systematically decrease to Mg# ~8 in the potassic granites. The leucosomes and pegmatites mostly preserve the Mg# of their host gneiss, with the exception of one leucosome, characterized by higher Mg# (13) than its host (8). According to the classification defined in Moyen and Martin (2012), only two of the samples can be classified as proper TTG (310-4 and 327-3; Table 3.2). These samples have Al2O3 contents slightly below 15 wt.%, hence just at the mark suggested to separate high- and low-Al TTG rocks

(e.g. Moyen and Martin, 2012). Their Al2O3/TiO2 ratios are 54.46 and 53.82, respectively. The

70 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.6: a) AFM (Na2O+K2O – FeO+Fe2O3 – MgO) wt.% ternary diagram displaying a calk-alkaline series trend for the gneisses from the Muccan Granitic Complex and South Muccan Shear Zone (boundary between tholeiitic and calk-alkaline series after Irvine and Baragar, 1971); b) molecular Na-K- Ca triangular diagram (after Barker and Arth, 1976), showing that the two proper TTG samples plot within the field for trondhjemitic (i.e. sodic calk-alkaline) differentiation, while the granitic samples diverge towards a transitional more potassic calk-alkaline trend. samples have low Mg# at 14 and 19, and Na2O contents and Na2O/K2O ratios of 5.00 and 2.58

(310-4), and 5.62 and 6.00 (327-3) at SiO2 at 71.79 and 72.16, respectively.

Within the leucosomes and pegmatites SiO2 content is generally higher than in the host gneisses, and ranges from 73.83 to 75.89 wt.% and 75.34 to 78.41 wt.%, respectively. The Al2O3 content is similar or slightly lower in the leucosomes and pegmatites, compared to the host gneisses. Only one sample shows higher Al2O3 and Na2O, and lower K2O and slightly lower SiO2 than its host rock. TiO2, Fe2O3, CaO, Na2O contents and Mg# are similar or slightly lower in the leucosomes compared to host gneisses, but K2O content is generally enriched in the leucosomes (Fig. 3.5e).

3.7.2 Trace elements

3.7.2.1 Rare earth elements (REE)

The diorites and granodiorites display relatively flat chondite-normalized HREE pattern (Fig. 3.7a) with Gd/Yb(N) of 2.9 and 2.3, respectively, at relatively high HREE concentration levels, compared to typical Archaean TTG (e.g. Martin, 2005). The HREE concentrations increase with decreasing SiO2 from granodiorite to diorite. The overall L/HREE fractionation is also relatively low compared to typical TTG, with La/Yb(N) ~13.5 and ~6.2, with decreasing SiO2. La/Sm(N) ratios yield values of 3.2 and 1.9. Ce anomalies are not observed, while Eu shows a slightly negative anomaly in the more primitive diorite (Fig. 3.7a). The two proper TTG samples show similar concentration levels of the most incompatible LREE, compared to the diorites/granodiorites, but display much higher overall L/HREE fractionation (La/Yb(N) ~35.2 and 51.0), characterized by similar HREE concentrations to the

71 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.7: Chondrite normalized (McDonough and Sun, 1995) rare earth element (REE) pattern for host gneisses (a) and in situ leucosomes and pegmatites (b); rock type symbols are from Figure 3.6. potassic granites, but significantly higher LREE concentrations (Fig. 3.7a). La/Sm(N) ratios of the two TTGs are 9.3 and 7.6, while their HREE pattern are almost flat with Gd/Yb(N) of 4.32 and 3.69. Both Ce and Eu anomalies are negligible, however the more silica-rich sample (327-4) shows a slightly negative Eu anomaly.

The potassic granitic gneisses (SiO2 ~74.97-76.10 wt.%) show progressively lower L/HREE fractionation compared to the TTG with increasing SiO2 and La/Yb(N) ratios of 11.2 – 25.3, and display overall decreasing REE concentration levels (Fig. 3.7a). The amount of L/HREE fractionation varies within the potassic granitic gneisses as a result of different flattening of the HREE patterns (Gd/Yb(N) of 2.33 – 4.46). Eu anomalies within the potassic granites are mostly absent. Only in one sample (308-3) a visible positive Eu anomaly is observed. The latter sample also displays the highest Ce anomaly. The other potassic granites show no to slightly positive Ce anomalies. Where present, the Ce and Eu anomalies increase with increasing SiO2.

72 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.8: Chondrite normalized (McDonough and Sun, 1995) multi-element spider-diagrams of selected trace elements for host gneisses (a) and in situ leucosomes and pegmatites (b); rock type symbols are from Figure 3.6.

The leucosomes and pegmatites display overall much lower concentration levels in REE. Their LREE fractionation is characterized by La/Sm(N) ranging from 2.02 to 4.21, but with relatively flat HREE pattern (Gd/Yb(N) 0.89 – 2.12). Eu anomalies systematically increase with decreasing overall REE concentration levels. Three of the leucosomes display positive Ce anomalies, which are absent or negligible within the two samples with the lowest overall REE concentration (Fig. 3.7b).

3.7.2.2 High field strength (HFSE), large ion lithophile (LILE) and other trace elements

Overall a systematic decrease in HFSE (Ti, Hf, Zr, Nd, Sm, Nb, Ta), as well as Sr and partial convective overturn, with increasing SiO2 and decreasing REE content, from diorite-granodiorite to tonalite-trondhjemite to potassic granite, is observed, while incompatible Pb, Th and LILE (e.g. Ba, K, Rb) tend to increase (Fig. 3.8a). The leucosomes and pegmatites have HFSE that mostly

73 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______overlap with the host gneisses, albeit with relative depletions in Th, and a larger range in Zr and Hf (Fig. 3.8b). It should be noted that particularly the more incompatible LILE could have been mobilized during alteration. However, the general LILE enrichment, and concentrations for Rb and K in the proper TTG are in good agreement with reported values for average Archaean TTG (e.g. Moyen and Martin, 2012). Only Ba shows a higher scatter, possibly due to some degree of post-emplacement alteration modification. Alteration may also have affected Ce concentrations to some degree, resulting in the observed variation in Ce anomalies. The increasing enrichment of LILE in the potassic granites may correspond to the overall increasing differentiation indicated by increasing silica content (Fig. 3.5). The more compatible HFSE are generally considered immobile, and the observed concentration levels and normalized signatures (e.g. negative Ti, Nb, Ta anomalies; Fig. 3.8a) are typical characteristics for Archaean TTG (Table 3.3; Moyen and Martin, 2012). Yttrium concentrations are very high in the diorite (36.3 ppm) and granodiorite (15.8 ppm), moderate in the TTG and potassic granites (9.9 to 3.4 ppm), and low in the leucosomes and pegmatites (0.4-1.8 ppm). According to Moyen and Martin (2012) the partial convective overturn concentrations of the two proper TTG samples fall in the range of medium-/high-Al (i.e. medium-

/high-pressure) TTG. As expected from the Al2O3 content, plotting in-between high- and low-Al type TTG, Sr content at corresponding SiO2, and Nb/Ta ratios (~14) identify the two proper TTG samples as medium-pressure TTG (see Moyen, 2011; Moyen and Martin, 2012). This is also consistent with Sr/partial convective overturn ratios, particularly from the more primitive proper TTG (310-4) with Sr/partial convective overturn of ~67.5. Niobium shows decreasing negative anomalies compared to neighboring La (LREE) with increasing SiO2 and decreasing REE concentration levels (Fig. 3.8a,b). In the potassic granitic gneisses characterized by the overall lowest REE concentration levels, Nb does not show any anomaly. However, Ta displays increasingly lower values that correlate with increasing Ti negative anomalies (Fig. 3.8a). The diorite, granodiorite and the proper TTG show higher normalized Ta relative to Nb, with Nb/Ta ratios of 6.6 to 14.7. A similar Nb-Ta fractionation (Nb/Ta of 13.1) is observed in the most primitive end-member of the potassic granites, while with decreasing SiO2 and REE, the other potassic granites display increasingly higher Nb/Ta ratios of 18.2 to 32.3. The leucosomes and pegmatites display Nb/Ta ratios of 11.2 to 18.4, likely inherited from their host gneisses.

3.8 U-Pb Zircon LA-ICP-MS Geochronology

In order to constrain the timing of protolith emplacement of the gneisses, zircon from a dioritic, a granodioritic, a tonalitic-trondhjemitic, and a granitic gneiss (Table 3.1) were analyzed for U-Pb LA-ICP-MS age determination. Additionally, detrital zircon from the sandstone (Fig. 3.2) that overlies an angular unconformity within the greenstone belt, were analyzed to test if the sandstone correlates with the regional 3427-3350 Ma Strelley Pool Formation, and if zircon provenance ages

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correspond to the crystallization ages within the Muccan Granitic Complex. U-Pb zircon data are given in Table 3.4 and selected zircon chemical data is given in Table 3.A-III (Appendix 9).

3.8.1 Muccan Granitic Complex gneiss components

3.8.1.1 Dioritic gneiss 370-1

A total of 19 spots on zircon grains from the dioritic gneiss 370-1 were selected for U-Pb age determination. The zircon grains are prismatic with well-developed pyramidal terminations, and are relatively large with an average c-axis elongation of ca. 400 µm (max. up to 600 µm), and medium to high aspect ratios (1:2 to 1:6). Most grains display concentric oscillatory internal structures, and few grains display convolute zoning, as revealed by CL imaging. One spot (spot#1, Table 3.4) was excluded from the evaluation due to outstandingly high Ti and low U, Th and Pb concentrations, indicating accidental sampling of a Ti-rich micro-inclusion. Two further spots (#11, 12) were excluded due to outstandingly high non-formula Fe contents. The remaining 16 spots yield 207Pb/206Pb ages ranging from 3475 ± 39 Ma to 3521 ± 30 Ma, with U-Pb discordance of -2.6 % (i.e. reverse discordant) to 3.4 % (normal discordant), where discordance is defined as the difference in 207Pb/206Pb age and 206Pb/238U age divided by the 207Pb/206Pb age multiplied by 100. A 207Pb/206Pb weighted mean age of 3493 ± 21 Ma is calculated. The 16 spots define a Discordia with an upper intercept age of 3499 ± 22 Ma (Fig. 3.9a; MSWD = 5.3), and a poorly-

77 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______constraint lower intercept age of 149 ± 2500 Ma. The upper intercept age of 3499 ± 22 Ma is interpreted to represent the timing of zircon crystallization during magmatic emplacement of the protolith of the dioritic gneiss. The age is in agreement with a better constraint age of 3503 ± 14 Ma for thermally annealed and chemically abraded core domains of zircon from the same sample (Wiemer et al., 2017; abstr.; MSWD = 1). The lower intercept indicates Pb loss in the Phanerozoic. Ti-in-zircon thermometry yields a weighted mean temperature of 680 ± 32 °C (Watson et al., 2006).

3.8.1.2 Granodioritic gneiss 309-2

From the granodioritic gneiss 309-2, 15 zircon spot analyses yielded 207Pb/206Pb apparent ages in the range of 3483 ± 54 Ma to 3581 ± 75 Ma (0.3 to 17.3 % discordance). Exclusion of one spot of outstandingly high Fe concentration (#4, Table 3.4), and two spots of high Ti content (#5 and 6), results in an upper Concordia intercept of 3576 ± 22 Ma (MSWD=1.9), defined by 12 spots (Fig. 3.9b). The age is interpreted as the timing of magmatic emplacement of the protolith of the gneiss. Ti-in-zircon thermometry yields a weighted mean of 766 ± 83 °C.

3.8.1.3 Tonalitic-trondhjemitic gneiss 310-4

A total of 45 spots were analyzed for U-Pb from zircon of the tonalitic-trondhjemitic gneiss 310-4. The zircon grains are pyramidal-prismatic and display concentric oscillatory CL zoning. The grain sizes range from ca. 50 to 200 µm, with aspect ratios of ca. 1:3 to 1:5. The sample shows a wide range of 207Pb/206Pb ages (2831-3549 Ma), characterized by a wide range of discordance (6.9- 65.4%). Excluding seven spots of outstandingly high Ti content, the remaining 38 spots define a Discordia yielding an upper intercept age of 3591 ± 36 Ma (MSWD=10.5), and a lower intercept age of 899 ± 87 Ma (Fig. 3.9c). The age of 3591 Ma is accepted as the timing of protolith emplacement. Ti-in-zircon thermometry yields a mean of 698 ± 55 °C.

3.8.1.4 Granitic gneiss 308-1

From the granitic gneiss 308-1, only eight zircon grains (of 50 to 100 µm size) could be extracted, and each was analyzed. Zircon from 308-1 is highly discordant (5.9-71.9 %) with 207Pb/206Pb ages ranging from 1587 to 3442 Ma (Table 3.4). The zircons show high HfO2 concentrations (>1.1 wt.%), and relatively high concentrations of non-formula Ca (80 to 6400 ppm), corresponding to high estimated alpha-doses (Table 3.4). Two spots that plot far off from Discordia and show elevated Ti contents are excluded. The remaining six spots define an upper intercept age of 3457 ± 110 Ma (MSWD=9.3) and a lower intercept age of 517 ± 260 Ma (Fig. 3.9d). Due to the low number of analytical spots and associated large error, the age cannot be considered reliable. However, the upper intercept of 3457 Ma indicates emplacement during either the Callina or the Tambina supersuite. Ti-in-zircon thermometry yields an average temperature of 704 ± 122 °C.

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Fig. 3.9: U-Pb zircon Concordia diagrams (Wetherill, 1956); a) dioritic gneiss 370-1; b) granodioritic gneiss 309-2; c) tonalitic gneiss 310-4; d) granitic gneiss 308-1; refer to text for discussion.

3.8.2 Detrital zircon from the quartz-sandstone

3.8.2.1 Quartz-jasper base conglomerate 332-1 and quartz-sandstone 335-1

A total of 69 detrital zircons (14 grains from the quartz-jasper base conglomerate 332-1, and 55 grains from the quartz-sandstone 335-1) were analyzed. The detrital grains range in size from ca. 30 to 150 µm, with aspect ratios of 1:2 to 1:5. All grains are characterized by oscillatory concentric zoning revealed by CL imaging. Both rounded and sub- to euhedral external morphologies are observed. All grains are affected by some degree of mechanical brakeage, particularly around the rims. The entire detrital zircon population yielded 207Pb/206Pb apparent ages ranging from 2930 to 3784 Ma for sample 335-1, and 3369 to 4638 Ma for sample 332-1 (Table 3.4). Given these data, the two samples are evaluated together. Data from selected trace elements (Appendix Table 3.A-III) revealed accidental sampling of micro-inclusions characterized by unusually high Ti contents, and unusually low Zr and Hf contents for three zircon grains (spot#2, 11 and 30). Furthermore, two analyses yielded extremely high 207Pb/206Pb ages (4638 and 4470 Ma). The high Pb contents are interpreted as representing the incorporation of unsupported excess Pb (1700 and 4230 ppm). One of these spots is characterized

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Fig. 3.10: U-Pb Concordia diagram for detrital zircon grains; a) “old” group; b) “young” group; refer to text for discussion. by extremely high U-Pb reverse discordance (-36; spot#71). The above ages were not considered reliable, and excluded from the following discussion. The remaining analyses display a correlation between increasing discordance and increasingly younger 207Pb/206Pb ages, with the least discordant (5.9 %) grain yielding a 207Pb/206Pb age of 3560 Ma, and the most discordant spot (230.8 %) yielding a 207Pb/206Pb age of 3184 Ma. In total, there are only five spots of <10 % discordance (commonly accepted threshold; Spencer et al., 2016), with 207Pb/206Pb ages of 3505, 3510, 3520, 3560, and 3423 Ma (Table 3.4). These ages are in agreement with the accepted minimum and maximum depositional ages of the Strelley Pool Formation based on the geochronological compilation of the Geological Survey of Western Australia (SHRIMP U-Pb zircon, conventional U-Pb zircon; references in Hickman, 2008; Kemp et al., 2015). Due to the high discordance it is deemed problematic to examine the probability/density – age distribution of the studied detrital grains. However, based on the 207Pb/206Pb ages, distinct age groups are recognized. A group of 28 spots of old apparent 207Pb/206Pb ages (3484-3605 Ma) yields a weighted mean 207Pb/206Pb age of 3538 ± 32 Ma, or a relatively well-defined Discordia with upper and lower intercepts of 3563 ± 19 Ma and 8 ± 88 Ma, respectively (MSWD = 2.9; Fig. 3.10a). Another 15 spots yield a weighted mean 207Pb/206Pb age of 3430 ± 20 Ma, or Concordia intercepts at 3454 ± 28 Ma and 12 ± 74 Ma (MSWD = 3.8; Fig. 3.10b). The upper intercept and weighted mean ages of both these groups are represented by the <10 % discordant spots given above. Furthermore, both age groups correlate with the upper intercept ages of the gneiss components from the Muccan Granitic Complex (i.e. 3457 Ma and 3576 Ma). A last group of 19 young 207Pb/206Pb very discordant spots defines a Discordia with intercepts at 3392 ± 22 Ma and 52 ± 37 Ma (MSWD = 2.9), respectively. Ti-in-zircon thermometry yields an average of 712 °C for the three most concordant spots from sample 332-1, and 697 °C for the three most concordant spots from sample 335-1. These temperatures are common Ti-in-zircon temperatures found in many plutonic rocks, about 50 to 100 °C more than minimum melt temperatures (Harrison et al., 2007; Fu et al., 2008). Note that the most discordant grains have higher Ti contents, which translates to spurious high estimated magmatic temperatures.

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Similarly, other trace element concentrations in the detrital zircon population show a positive correlation with increasing U-Pb discordance. It remains unclear if primary magmatic compositions controlled the degree of subsequent U-Pb disturbance, or vice versa. A correlation between increasing P, and REE3+ with decreasing measured Si (counts per second), for example, may suggest complex xenotime substitution (P5+ + REE3+ = Si4+ + Zr4+; e.g. Hoskin and Schaltegger, 2003) associated with elevated REE3+ content in the source magma (e.g. Hanchar et al., 2001). Accompanying increase in U and Th in the zircon may have resulted in radiation damage induced metamictization of trace element enriched magmatic zircon. On the other hand, metamictization of initially high U and Th zircon may have altered the zircon susceptible to later chemical modification by low-temperature hydrothermal fluids (e.g. Geisler et al., 2003; Allen et al., 2016, abstr.; Wiemer et al., 2017, abstr.). The latter is supported by the fact that particularly LREE and non-formula elements (e.g. Ca) increase with increasing discordance.

3.9 Discussion

The presented geochronological data demonstrates the antiquity of the studied gneisses along the southwestern margin of the Muccan Granitic Complex, with upper Concordia zircon U-Pb ages of 3499-3591 Ma. In order to distinguish these rocks from the younger regional granitic supersuites within the Muccan Granitic Complex, I will refer to the 3591-3499 Ma components, which represent some of the oldest continental rocks so far detected in the East Pilbara Terrane, as the Doolena Suite gneisses. Based on the petrological and geochemical data, the magmatic formation and differentiation of the protolith of the gneisses will be discussed below, supported by thermodynamic and geochemical trace element modeling to constrain depth and degree of partial melting of a suitable source rock composition. Magmatic reworking of the gneisses will be discussed based on the petrological-geochemical data from the in situ leucosomes and pegmatites, with implications for dome-and-keel development. U-Pb zircon ages from the quartz-sandstone formation will be discussed in regards to detrital provenance and timing of dome un-roofing. Finally, a tectono-magmatic crustal growth model for the Doolena Gap – Muccan Granitic Complex is discussed to comment on the early Archaean East Pilbara Terrane crustal evolution.

3.9.1 Petrogenetic constraints on the formation of the 3499-3591 Ma Doolena Suite gneisses

In order to constrain the petrogenesis of the protoliths of the early Archaean Doolena Suite gneisses, a combination of thermodynamic and geochemical modeling has been implemented. The stability of metamorphic mineral assemblages of a suitable source rock composition over a range of mid- to lower crustal pressure (P) and temperature (T) conditions is computed. The thermodynamic modeling aims to distinguish P-T conditions and corresponding mineral assemblages between three different modes of juvenile Archaean granite/TTG formation put forth in previous geochemical and experimental studies: i) low-pressure partial melting of garnet-free amphibolite, ii) medium-pressure partial melting of garnet amphibolite, and iii) high-pressure

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Fig. 3.11: Theriak-Domino pressure (P) - temperature (T) pseudosection of stable mineral assemblages computed for the average C-F2 basalt (n=8; data from Smithies et al., 2007) in the model system

NCFMASHT (Na2O-CaO-Fe2O3-MgO-Al2O3-SiO2-H2O-TiO2) with excess H2O (5 mol%); present day continental and estimated minimum Archaean shield geotherms from Condie (1984) and references therein; light grey area indicates garnet-free assemblages, dark grey area indicates stability of garnet; three end-member models for TTG formation are shown: i, garnet-free amphibolite (low-P); ii, plagioclase-free garnet amphibolite (medium-P); iii, rutile eclogite (high-P); the yellow arrow indicates the preferred mode of parental melt production; refer to text for description and discussion. partial melting of rutile-bearing eclogite (e.g. Foley et al., 2002; Rapp et al., 2003; Moyen and Martin, 2012). Computed P-T conditions, modal abundances of stable phases, and inferred melt- producing metamorphic reactions help to constrain the mechanism of continental melt production for the studied Doolena Suite gneisses. The output will be used to perform geochemical trace element modeling with the attempt to match model results with the observed rock compositions. By no means is the purpose of the thermodynamic-geochemical modeling to re-visit general aspects of the petrogenesis of continental melt production in the Archaean (e.g. review in Moyen and Martin, 2012), but to constrain the mode of formation of the Doolena Suite gneisses based on the previously established models of low-, medium-, and high-pressure TTG melt derivation. The model assumptions are: Firstly, based on geochemical and experimental evidence, it is generally accepted that in order to produce TTG-like trace element signatures, the source of the partial melts must be an enriched metabasalt characterized by slightly negative Ti, Nb and Ta anomalies, and enrichment in LILE and LREE, compared to modern MORB (e.g. Martin et al., 2014; Nagel et al., 2012; Polat, 2012; Smithies et al., 2009). Over-thickened oceanic island arc basalts (e.g. Polat, 2012; Nagel et al.,

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2012), enriched tholeiites extracted from a less depleted (i.e. primitive) early Archaean mantle (e.g. Smithies et al., 2009), or contamination of basaltic magma with pre-existing felsic material, either as a recycled mantle component, or during crustal ascent (e.g. Smithies et al., 2009; Bédard, 2006), have been proposed as likely melt sources. Such distinction is beyond the scope of this study. However, it should be noted that for the East Pilbara Terrane a non-uniformitarian model in absence of evidence for arc-style subduction is favored (e.g. Van Kranendonk et al., 2015). Secondly, the source metabasalt must be hydrous, consistent with the overall calk-alkaline trend that requires elevated water content (>2 wt.%; e.g. Zimmer et al., 2010; Condie, 2015). A resulting hydrous magma causes the suppression of plagioclase crystallization relative to mafic phases, and the suppression of crystallization of silica phases in favor of Fe-oxides, and stabilization of amphibole (i.e. hornblende), leading to Fe depletion during magma evolution (Zimmer et al., 2010). Furthermore, a hydrous source shifts the solidus towards lower temperatures. Hydration of the source metabasalt could have resulted from the interaction with magmatic or metamorphic fluids in the crust, or from the deep burial of mafic volcanic rocks that have been altered by interaction with surface seawater (e.g. Thébaud et al., 2006, 2008). Furthermore, it is assumed that the protolithic compositions represent first extraction melts from the hydrous basaltic source. Besides the more potassic granitic compositions, it is unlikely that the relatively mafic diorite-granodiorite and proper TTG samples are products of reworking of pre-existing continental (e.g. TTG) crust, as this would require almost 100 % melt fractions. Finally, geochemical data from this study suggest possible differentiation trends (Fig. 3.5) starting from the more mafic end-members of the suite. Hence, I consider the most mafic TTG (310-4) and/or the diorite (370-1) or granodiorite (309-2) as being closest to a potential parental melt composition.

3.9.1.1 Formation of parental continental magma through partial melting of hydrous metabasalt

a) Constraints from thermodynamic modeling (Theriak-Domino)

Thermodynamic modeling of stable mineral assemblages in mid- to lower crustal P-T space was performed using the Theriak-Domino software (de Capitani and Petrakakis, 2010; internally consistent thermodynamic database from Holland and Powell, 1998) with excess H2O (0.1-5.0 mol%) in the NCFMASHT model system (Fig. 3.11). An average Archaean enriched tholeiite composition (3525 Ma Coonterunah C-F2 type; Smithies et al., 2007) was used (Fig. 3.11). It has been previously demonstrated that compositions of the C-F2 type basalts represent a suitable source for post-3500 Ma TTGs from the East Pilbara Terrane (Smithies et al., 2009). Modeling of water-absent compositions does not predict the stability of amphibole, and the solidus is at very high temperatures of ca. 1200 °C. As mentioned above, a hydrous source is required to produce calc-alkaline melts. At low H2O of 0.1 mol%, only ~1% modal hornblende is computed to produce melt through de-hydration – not enough to derive sufficient partial melt

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volumes. Therefore, more realistic water contents of H2O 5 mol.% were modeled (Fig. 3.11), in agreement with constraints on the stabilization of amphibole (i.e. hornblende), requiring source rock H2O in excess of ~3 wt.% (Zimmer et al., 2010). At these conditions the pseudosection (Fig. 3.11) shows that the primary mode of partial melting is through metamorphic de-hydration breakdown of hornblende, defining the solidus in the computed system. Based on the background of previously proposed TTG melt derivation through partial melting within over-thickened crust, minimum melting conditions are explored, following P-T paths defined by estimated minimum geothermal gradients (Fig. 3.11; present day continental and Archaean shield geotherms from Condie, 1984, and references therein). The three end-member TTG formation scenarios are shown in the pseudosection (Fig. 3.11): i) low-P melting in absence of garnet falls in the range of elevated Archaean geotherms (e.g. Condie, 1984), or requires additional heat input to produce melts at ca. 5-6 kbar, >770 °C (Fig. 3.11-i); ii) medium-P melting of garnet amphibolite is possible at estimated minimum geotherms, at ca. 13.5 kbar, >700 °C (Fig. 3.11-ii); and iii) the stability field for high-P rutile-bearing eclogite lies along the estimated minimum Archaean shield geotherm, however, ca. 300 °C higher than the computed solidus at ca. 24 kbar, 850 °C (Fig. 3.11-iii). The latter scenario of high-P rutile-eclogite melting seems unlikely based on the presented thermodynamic computation. Partial melting would occur at much lower temperatures, before rutile as the primary Ti-bearing accessory phase is stabilized at ca. 1150 °C. Using other thermodynamic databases, it has been shown that rutile may be stable at lower pressures (e.g. ~11 kbar, Nagel et al., 2012). Therefore, the presence of rutile as a residual phase in TTG melt production requires further evaluation by geochemical modeling (see next section). At medium-P, the pseudosection indicates that partial melting of garnet-amphibolite through the breakdown of hornblende would occur at relatively low temperatures (~700 °C; Fig. 3.11-ii), leaving an eclogitic residue of mostly garnet, clinopyroxene and ilmenite above ~ 13.5 kbar, 750 °C (Fig. 3.11-ii). Theriak calculations yield a sub-solidus assemblage of garnet (33%) – clinopyroxene (31%) – hornblende (29%) – ilmenite (3.5%) – quartz (2.5%) – plagioclase (1%) at ~600 °C, 11 kbar. Finally, partial melting at low-P (Fig. 3.11-i) would result in a residue rich in plagioclase and comprising olivine, but with low to no modal abundance of garnet (Fig. 3.11-i). Due to the low HREE content of the studied samples, and their classification as medium- to high-Al (i.e. low- to high-P) TTG (after Moyen and Martin, 2012), melting at rather elevated pressures (residual garnet) is favored. This will be discussed further based on geochemical modeling in the next section.

b) Constrains from geochemical trace element modeling

In order to model trace element variations through de-hydration breakdown of hornblende as the primary melt producing reaction, the equation for non-modal batch melting (Shaw, 1970) is used:

퐶0 퐶퐿 = [1] 퐷0+퐹(1−푃)

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Fig. 3.12: Chondrite normalized multi-element spider-diagram for geochemical modeling of C-F2

metabasalt (C0) source melting to derive a melt (CL) that resembles naturally observed tonalite- trondhjemite (target 310-4) trace element concentrations and signatures implementing non-modal batch

melting; the modeled melt (CL) represents 25 % partial melting of 25 % clino-pyroxene – 35 % hornblende – 30 % garnet – 10 % ilmenite for 100 % hornblende melting; also shown are variations in Nb-Ta by using 10 % rutile instead of ilmenite in the residue, and Sr variations by melting of plagioclase (10 % plagioclase – 90 % hornblende); refer to text for description. with CL, concentration of trace element in the melt; C0, initial concentration of trace element in the solid; D0, bulk partition coefficient for minerals in the initial solid; P, bulk partition coefficient for minerals entering the melt fraction; F, partial melt fraction. Bulk partition coefficients D0 and P were calculated as:

푖 푗 퐷0 = ∑ 퐾퐷 ∗ 푥 [2]

j i with x , proportion of mineral (wt.%) in the initial solid; KD , partition coefficient of element i for mineral j. The used partition coefficients are from Bédard (2006). In an attempt to fit the trace element concentration between the modeled melt and the most basic TTG end-member (310-4) through disequilibrium partial melting, the average metabasalt C-F2 source was used. A good match was achieved by 25 % partial melting of a garnet–amphibolite (i.e., wt.%: 25% clinopyroxene-35% hornblende-30% garnet-10% ilmenite) by non-modal melt extraction of 100% hornblende as the melt-entering phase (Fig. 3.12). About 30% garnet is required in the residue to match the HREE pattern of the TTG 310-4, which is in agreement with the mode of garnet in the Theriak calculation for a medium-P melting path along a minimum geotherm (Fig. 3.11-ii). Modeling rutile as the critical phase to distinguish high-P melting residues would yield negative Nb-Ta anomalies that are too large (Fig. 3.12). Furthermore, using the partition coefficients from Bédard (2006), rutile fractionates Nb > Ta (chondrite normalized), but most of the here observed TTG and diorites have Ta > or >> Nb (chondrite normalized; Fig. 3.8; Fig. 3.12). Ilmenite, on the other hand, yields a better match (Fig. 12), and is in agreement with partial

85 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.13: a) Nb/Ta versus Zr/Sm plot testing the mode and source composition of Archaean TTG formation; rutile-eclogite, eclogite, and amphibolite fields are from Foley et al. (2002) for melting of a chondritic source; the low Nb/Ta field (light grey) indicates melting of a typical Archaean basalt source transformed to eclogite (Rapp et al., 2003); the proper TTG from the Doolena Suite plot within or close to the amphibolite field (after Moyen and Martin, 2012). b) La/Yb(N) versus Yb(N) plot testing previous models of TTG petrogenesis; trends for partial melting of i: a garnet-free source (pink; grt0), ii: a garnet amphibolite source (green; shown for 10 and 25 % garnet), and iii) an eclogite source (orange; ecl), are shown (after Moyen and Martin, 2012); it is shown that the proper TTG rocks from the studied Doolena Suite plot at elevated La/Yb(N) ratios at medium to low Yb(N), requiring the presence of garnet in the source; it is discussed in the text that the proposed trends are not recommended to solely define TTG petrogenesis without considering fractional crystallization to modify the observed REE concentrations. melting at ~13.5 kbar, as indicated by the pseudosection (Fig. 3.11-ii). This is demonstrated in the Nb/Ta versus Zr/Sm plot of Figure 3.13a, in which the studied TTG plot closer to the amphibolite source field compared to the fields for eclogite sources (Foley et al. 2002; Rapp et al., 2003; Moyen and Martin, 2012). Partial melting at low-P conditions was investigated by modeling the effects of plagioclase entering the melt, and as a residual phase. Fig. 3.12 shows that melting of plagioclase would be required to achieve positive Sr relative to neighboring Nd in a chondrite-normalized plot. However, at above ca 10 % plagioclase in the source, the normalized Sr concentrations become too high in the modeled melt. The absence of significant Eu anomalies, and the La/Yb(N) ratios (Fig. 3.13b), further support the interpretation that garnet, and not plagioclase, dominated the source of the parental TTG melts. For the medium-P source garnet amphibolite assemblage modeled melt compositions are not perfect either, with calculated melt MREE being too high. To achieve a better match, less hornblende as a melt-entering phase is required. This implies either that some hornblende represents a residual phase or that almost pure hornblende cumulate was extracted through fractional crystallization early in the magma evolution. Both, hornblende as a residual phase (e.g. Martin, 1987, 2005; Kramers, 1988), or the extraction of a hornblende-rich cumulate (ca. 45 % fractionation of >80 % hornblende; Bédard, 2006) have been previously proposed. Archaean hornblendites and hornblende pyroxenites have been discovered, for example, in the Fiskenæsset anothosite complex, SW Greenland (e.g. Polat et al., 2012; Huang, 2012).

86 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Summing up, two end-member models for the petrogenesis of the diorites/granodiorites need to be investigated: i) the diorites-granodiorites represent fractionated cumulate portions that are reciprocal to the more evolved TTG and granites, or ii) these rocks also represent juvenile melt compositions derived from partial melting of hydrous metabasalt. The diorites and granodiorites have relatively high HREE and show overall L/HREE fractionation much lower than the TTG (e.g. 310-4; Fig. 3.8, Fig. 3.12) of this study. In order to produce such trace element signatures from melting of the same Coonterunah C-F2 source, garnet-free partial melting is required to reproduce the observed HREE pattern at similar melt fractions (~20-30 %; Fig. 3.13b). This would imply that these rocks were formed at low-P conditions (Fig. 3.11-i). If this were true, infra-crustal melting over a wide range of pressures, as proposed for the derivation of post-3500 Ma East Pilbara Terrane TTGs, is substantiated (Smithies et al., 2009). However, for the modeled C-F2 basalt composition, partial melting of garnet-free amphibolite is restricted to pressures less than 6- 8 kbar in the temperature range 750-1000 °C (Fig. 3.11-i). Based on the observation of magmatic epidote in the diorites, which indicates crystallization in excess of 8 kbar (e.g. Zen, 1984), fractional crystallization is more likely. In synthesis, it is argued that the parental TTG magma was produced by ~25% partial melting of a hydrous garnet amphibolite at ca. 45-50 km depth, along a minimum geothermal gradient (Fig. 3.11-ii), at minimum temperatures of ca. 750 °C (wet solidus). Modeling indicates that parental melts likely evolved through fractional crystallization into the diorites-granodiorites, which make up the field-sampled hornblende, plagioclase and titanite cumulates, and more evolved granitic melts.

3.9.1.2 Derivation of compositional types: fractional crystallization versus depth and degree of partial melting

Based on the assumption that mafic proto-crust had to reach sufficient thickness to produce the post-3500 Ma medium- to high-P TTG rocks in the East Pilbara Terrane, it has been proposed that pre-3500 Ma granitic/TTG melts may have formed as low-P partial melting products (e.g. Smithies et al., 2009; Van Kranendonk et al., 2015). The petrological-geochemical modeling presented above indicates derivation of a parental partial melt that resembles the observed tonalite- trondhjemite 310-4 with a U-Pb zircon crystallization age of 3591 Ma. The oldest granodiorite dated in this study is represented by sample 309-2 with an upper Concordia zircon age of 3576 Ma. As discussed above, the lower L/HREE fractionation, with relatively high HREE concentration levels of this rock, may be produced by low-P partial melting of a garnet-free source assemblage (Fig. 3.11-i) except that the presence of magmatic epidote (>8 kbar), indicate magmatic emplacement in elevated crustal depths, where garnet is stable. Therefore, proposed models of hornblende fractionation (e.g. Bédard, 2006) are considered here. Such a petrogenetic model would imply the formation of granodioritic and dioritic rocks as fractionated portions from a parental melt that was derived at medium-P. Following Bédard (2006), who demonstrated that various TTG compositions might be derived by fractional crystallization with extraction of hornblende-plagioclase dominated (diorite) cumulate, it is here considered that the studied dioritic

87 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.14: Chondrite normalized multi-element spider-diagram for geochemical modeling of Rayleigh

fractional crystallization of a tonalite-trondhjemite melt (310-4, C0), targeting granodiorite 309-1 as the fractionating solid (CS); the modeling demonstrates a relatively good match by 10 % extraction of 72 % hornblende – 10 % plagioclase – 10 % ilmenite – 5 % titanite – 1.5 % rutile – 1.3 % apatite – 0.2 % zircon for the target fractionated solid 309-2; refer to text for discussion of the remaining liquid composition. and granodioritic rocks represent cumulate-type fractions of a parental melt. Cumulate fractionation could explain the observed SiO2 gap between the diorite-granodiorite and the TTG- evolved granite series (Fig. 3.5). Fractional crystallization of a parental TTG melt is geochemically modeled using the equation for Rayleigh fractional crystallization (Neuman et al., 1954):

퐶 퐶 = 0 [3] 퐿 퐹(퐷−1) for the remaining melt, and:

퐶 퐶 = 0 [4] 푆 퐷∗퐹(퐷−1)

for the solid residue, with CS, concentration of trace element in the remaining solid. The TTG 310- 4 composition is used to model a parental melt. The geochemical trace element modeling (Fig. 3.14) demonstrates that ca. 10 % hornblende dominated fractionation (72% hornblende - 10% plagioclase - 10% ilmenite - 5% titanite - 1.5% rutile - 1.3% apatite - 0.2% zircon) produces a modeled cumulate (CS) that resembles the naturally observed granodiorite composition (309-2;

Fig. 3.14), and a remaining slightly evolved melt (CL). Differences between modeled residue (CS) and the granodiorite are a more negative Pb anomaly, and much lower Zr/Hf. Pb content difference is minor and could be explained by mobility of Pb in the natural sample, or by

88 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______adjustment of modeling minor plagioclase to fractionate. The natural samples and (granitic) melt all have similar Zr/Hf, but the modeled residua are richer in Hf, relatively. The reason for a difference in Zr/Hf is unclear. It is possible that the used partition coefficients are not applicable to this natural example. At the modeled proportions of ~10 % solid fractionation, the remaining melt is overall compositionally similar to the source (i.e. 310-4), but it is shown that lower levels in HREE and Ti are achieved, as observed in some of the natural granites (Fig. 3.14). It is argued therefore that fractional crystallization played a major role in the derivation of the compositional spectrum. Particularly the granodiorite 309-2 resembles the modeled extraction of a fractional crystallization portion from a medium-P parental melt. This builds upon the case that at 3591 Ma parental TTG melts were produced at elevated pressures (~13 kbars), indicating crustal thickness of at least 45- 50 km. It should be noted that the modeled Rayleigh fractional crystallization represents a simplified assumption, and both magmatic recharge through continuous melt production at the lower crustal base, as well as possible assimilation with previously crystallized or pre-existing wall-rock and/or magma mixing may have altered the observed compositions to some degree. However, the model that different TTG compositions require different degrees of partial melting at different crustal depth and that pre-3500 Ma TTG were formed at low-P (Smithies et al., 2009; Van Kranendonk et al., 2015) is not confirmed. A model is proposed, in which 3499-3591 Ma parental TTG magma was produced within a specific partial melt zone at the base of 45-50 km-thick hydrous mafic crust, and infiltrated the crust, where it fractionated to produce dioritic-granodioritic portions.

3.9.2 Reworking and un-roofing of the Muccan Granitic Complex - implications for early (>3427 Ma) dome-and-keel initiation

3.9.2.1 Time constraints and detrital provenance

As outlined in Chapter 2, intensive deformation (D2), geometrically associated with doming of the Muccan Granitic Complex, affected the western Doolena Gap greenstone belt prior to formation of the angular unconformity beneath the quartz-sandstone formation (Fig. 3.2). The new detrital zircon data confirms the correlation of the quartz-sandstone with the regional 3427-3350 Ma Strelley Pool Formation (Hickman, 2008; Van Kranendonk, 2010). It is therefore suggested that dome-and-keel development due to partial convective overturn occurred between ca. 3460 and 3427 Ma (Chapter 2), hence prior to the regional main event of partial convective overturn in the East Pilbara Terrane at ca. 3310 Ma (e.g. Sandiford et al., 2004; Van Kranendonk et al., 2015). Considering only the most concordant detrital zircon ages, two igneous provenances are inferred: an early source with crystallization ages between 3505 and 3560 Ma, and ii) a younger provenance represented by a crystallization age of ca. 3423 Ma. The older provenance coincides with the emplacement ages of the newly discovered 3499-3591 Ma Doolena Suite gneisses (this study), while the younger detrital provenance coincides with emplacement ages of the regional

89 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

3450-3420 Ma Tambina Supersuite (e.g. Van Kranendonk et al., 2006). The Tambina episode of granitic magmatism was contemporaneous with the proposed early partial convective overturn event (3460-3427 Ma), and parts of the Muccan Granitic Complex have been previously assigned to the Tambina Supersuite (Fig. 3.1; Van Kranendonk et al., 2006). The coincidence of the detrital ages with the emplacement ages within the Muccan Granitic Complex suggest that doming and associated tilting of the greenstone package resulted in a topography that led to the un-roofing and erosion of the Muccan Granitic Complex to disperse proximal sedimentary detritus into the newly developed sub-basin that is now the Strelley Pool Formation. The poorly sorted base conglomerate (332-1) of the Strelley Pool Formation comprises immature, angular detrital material (e.g. jasper chert, quartz, minor feldspar) supporting a proximal sediment source. The well-sorted quartz- sandstone (335-1) higher up in the Strelley Pool Formation stratigraphy is interpreted to have matured through in situ reworking within a shallow marine marginal to fluviatile environment (Chapter 2). The sedimentary reworking and maturation possibly caused a concentration of detrital zircon, as reflected in the higher abundance of detrital zircon grains in the quartz-sandstone 335-1 compared to the immature base conglomerate 332-1. The interpretation of proximal sediment sources is consistent with a recent study by Kemp et al. (2015) who discussed the provenance of >3500 Ma detrital zircon from various sedimentary formations in the East Pilbara Terrane. The authors concluded that the provenance was likely proximal and derived from erosion of local granitic domes, as age groupings of post-3500 Ma zircon grains coincide with proposed uplift and erosion of granitic domes in the East Pilbara Terrane (e.g. Hickman and Van Kranendonk, 2004; Kemp et al., 2015). The absence of concordant ages younger than 3427 Ma in the presented detrital record is not surprising given the interpretation that in the study area partial convective overturn occurred between 3460 and 3427 Ma (Chapter 2). If the maximum depositional age of the Strelley Pool Formation in the Doolena Gap greenstone belt is accepted at 3427 Ma (e.g. Hickman, 2008), magmatic quiescence (no zircon crystallization) would be expected during 3427 and 3350 Ma, while the exposed pre-3427 Ma material was eroded. At this point it should be noted that the presence of thermally softened mid-crustal granitic substrate and associated buoyant rise of granitic partial melts are considered a driving mechanism of the 3310 Ma partial convective overturn event in the East Pilbara Terrane (e.g. Collins et al., 1998; Sandiford et al., 2004). The data and interpretations presented here suggest that the early 3460-3427 Ma partial convective overturn occurred during the final stages of juvenile TTG production (e.g. Smithies et al., 2009), which in turn may have caused reworking of some of the more ancient continental mid-crustal components. Based on field evidence (Chapter 2), it is argued that tight D2-folding of syn- deformational leucosomes within the 3499-3591 Ma Doolena Suite gneisses (Fig. 3.3b,c) formed during this early partial convective overturn event. In the next section, the mechanism of partial melting to form the observed syn-deformational in situ leucosomes will be discussed.

90 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.15: Chondrite normalized rare earth element pattern for geochemical modeling of disequilibrium

partial melting of host granitic gneiss 327-4 (C0) to match the spatially associated in situ leucosome 327- 4 B (CL); the range in REE concentration of other leucosomes and pegmatites is indicated in the grey- shaded area; a good match in overall REE concentration level is achieved for 5-10 % partial melt fractions by 50 % alkali feldspar – 50 % plagioclase melting of an initial solid modeled as 45 % plagioclase – 24 % biotite – 15 % apatite – 10 % titanite – 5 % alkali feldspar – 1 % allanite; to match the observed positive Eu anomaly in the leucosome, ca. 50 % partial melting is required, or the used partition coefficients are incorrect; refer to text for discussion.

3.9.2.2 Constraints on reworking mechanism and dome initiation

Other workers have suggested that the major (3310 Ma) dome-and-keel forming event in the East Pilbara Terrane was triggered by partial melting of pre-existing felsic continental mid-crust (e.g. Sandiford et al., 2004; Van Kranendonk et al., 2015). The dominant cause of this partial melting is debated. Van Kranendonk et al. (2015) recently reviewed the cause of partial melting as a combination of: i) burial of granitic upper crust into warm mid-crust as a result of continuous supracrustal volcanism, ii) thermal blanketing and heating from radioactive decay of U-, Th- and K-enriched granitic mid-crust, and iii) heat from a mantle plume source. This 3310 Ma partial convective overturn occurred after a ca. 80 Ma hiatus in tectono-magmatic activity, indicated by the deposition of the regional 3427-3350 Ma Strelley Pool Formation (e.g. Hickman, 2008), and most of the post 3350 Ma TTG and granites are interpreted to have formed through re-melting of the pre-existing >3420 Ma TTGs (e.g. Smithies et al., 2009). In this work an earlier partial convective overturn event is proposed. Its age (3460-3427 Ma) coincides with the final stages of juvenile TTG production (e.g. Smithies et al., 2009), and thus this is a new mechanism for reworking of pre-existing TTG and granite producing the syn-deformational leucosomes. This requires discussion. In order to investigate the formation of the in situ leucosomes, geochemical modeling was performed. Due to the petrographic observations of antiperthitic feldspars and graphic quartz lamella textures within feldspar (Fig. 3.4e), eutectic compositions of the leucosomes are inferred. The low REE concentrations, and flat HREE pattern of the leucosomes compared to their host

91 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______gneisses (Fig. 3.7b), indicate retention of REE. It is likely that accessory phases in the host gneisses (i.e. apatite, zircon) did not melt and retained significant HREE concentrations. This is consistent with the observed lack of zircon in the leucosomes. Previous studies argued that melt- segregation rates are higher than chemical equilibration rates during mid-crustal partial anatexis, resulting in disequilibrium melting (e.g. Sawyer, 1991). In support, unpublished data shows that zircon within the host gneisses comprise a thin outer recrystallization rim, formed at ~3447 Ma, indicating temperatures were high enough to recrystallize part of the zircon, but not high enough to melt zircon (Wiemer, unpublished). Therefore, disequilibrium (i.e. non-modal) partial melting was modeled using the equation given above (Shaw, 1970), with additional KD values for alkali feldspar from Leeman and Phelps (1981). The granitic host gneiss 327-4 was used as a starting composition (C0), and the spatially associated (i.e. in situ) leucosome 327-4B was targeted as a partial melt (CL). The modeling resulted in a good match in REE pattern through ca. 5-10 % partial melting (50 % alkali feldspar - 50 % plagioclase), to produce the low REE concentrations in the leucosome (Fig. 3.15), which is in agreement with the observed eutectic composition, not accounting for REE partitioning into quartz. For the initial solid (C0), a modal assemblage of 45% plagioclase – 24 % biotite – 15 % apatite – 10 % titanite – 5 % alkali feldspar – 1 % allanite yields a good match. Significant amounts of apatite and minor allanite are required in the residue to match the observed LREE pattern. Although the modeled modal abundance of apatite in the residue is higher than observed in the granitic gneiss 327-4, the overall lack of REE-rich accessory phases (e.g. monazite, apatite, zircon, allanite) in the leucosomes supports the retention of REE and particularly LREE during melt segregation, favoring rather low temperatures, insufficient to melt accessory phases (see above). The overall REE concentration levels can be modeled by very low melt fractions (5 %). However, either melt-fractions of ca. 50 % are required to match the observed positive Eu anomalies in the leucosomes (Fig. 3.15), or the used partition coefficients for Eu (feldspar-melt) are incorrect. Considering the syn-deformational nature of the in situ leucosomes, high melt fractions may still be in agreement with rather low-temperature partial anatexis. Although some migration of the partial melts cannot be precluded (possibly involving further fractionation of plagioclase), the overall volume of leucosomes compared to the volume of adjacent melanosomic layers is relatively high, based on field observations (Fig. 3.3b,c). As evident in the pre-D2 folding migmatization (Chapter 2; i.e. segregation of quartzofeldspathic layers) of the host gneisses, it may be proposed that during progressive ductile shear deformation sub-solidus segregation of more felsic recrystallized domains occurred (e.g. Fig. 3.4e). Such layers would have been more fertile for partial anatexis under low-temperature hydrous conditions due to more felsic compositions and smaller grain sizes (e.g. Johannes and Gupta, 1982). It has been demonstrated that flow of fluids and melts is more effective in smaller grain-size domains (Wark and Watson, 2000). Grain size reduction through dynamic recrystallization of quartzofeldspathic segregates resulted in a susceptibility for fluid (volatile) influx, supporting partial melting at low temperatures across a wet solidus.

92 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

Fig. 3.16: Final a) Schematic histogram of early East Pilbara Terrane crust showing the frequency of pre- 3500 Ma detrital zircon from Kemp et al. (2015) and of the entire detrital zircon population from this study; the age range of the Callina (Cal.) and Tambina (Tamb.) Supersuites and the newly proposed Doolena Suite (*) are indicated above; b) pseudo - time-sliced cartoon illustration of the accumulation of granitic/TTG material during early vertical growth of the East Pilbara Terrane crust; the model envisages a rapid vertical growth of mafic proto-crust through mantle-derived volcanic resurfacing events in a stagnant lid tectonic regime; increasing accumulation of felsic mid-crust from a ‘zone of partial melting’ at the over-thickening lower crustal base ultimately leads to a critical Rayleigh-Taylor instability (within ca. 200-300 Ma) causing early partial convective overturn at around 3460-3427 Ma, hence the transition towards a dynamic stagnant lid regime.

The studied pre-3500 Ma Doolena Suite gneisses come from the very margin of the Muccan Granitic Complex dome, while most of the Muccan Granitic Complex components are interpreted to have formed between 3490 and 3420 Ma, or later (e.g. Van Kranendonk et al., 2006). Due to the in situ syn-deformational nature of the leucosomes, in combination with the fact that even the diorites and granodiorites, which were emplaced into the lower mid- to shallow lower crust (magmatic epidote), show evidence for syn-deformational low-temperature partial anatexis, it is suggested that the here observed partial melting was the result of the continuous TTG magma infiltration evident in the center of the dome (i.e. Tambina Supersuite; Fig. 3.1; Van Kranendonk et al., 2006). The continued voluminous magmatism would have supplied heat and likely H2O (i.e. volatiles) into the marginal areas of the developing dome to cause partial anatexis during ductile deformation at relatively low temperatures. This model stands in contrast to the partial melting

93 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______mechanism associated with major reworking that affected the East Pilbara Terrane mid-crust around 3310 Ma (e.g. Collins, 1993; Van Kranendonk et al., 2015). The early partial convective overturn likely occurred during the final stages of juvenile continental melt production, and was associated with infiltration of volatile (H2O)-rich fluids liberated during crystallization of TTG and granite in the central parts of the Muccan Granitic Complex at 3460-3420 Ma. Based on the presented field observations and petrography, and the constraints on the petrogenesis of the partial melts, following model is proposed: i) development of an early foliation

(S1; Chapter 2) during dome expansion and shear deformation along the margins of the initiating dome due to magmatic infiltration in its centre; ii) heat and volatile (e.g. H2O) influx caused progressive metasomatism, likely sourced from crystallizing magma in the deeper dome centre, infiltrating developing migmatitic segregates during ongoing deformation; iii) crossing of a wet low-temperature solidus resulted in syn-deformational eutectic partial melts during D2-folding (Fig. 3.3c; Chapter 2); iv) ascend into shallower crustal levels due to accumulating buoyant partial melts (and dome-up shear deformation at the Muccan Granitic Complex margin; Chapter 2) into a regime were deformational strain was increasingly accommodated by the partial melts, ultimately forming late-stage melt pockets that intruded as pegmatites subsequent to cessation of the bulk deformation episode (D2).

3.9.3 Implications for early continental crust formation in the East Pilbara Terrane

Combining the new emplacement ages with previous and new detrital zircon data from the East Pilbara Terrane (e.g. Kemp et al., 2015; this study) suggests that the 3499-3591 Ma protoliths of the newly discovered Doolena Suite gneisses from the SW margin of the Muccan Granitic Complex represent part of a major episode of continental crust formation in the East Pilbara Terrane (Fig. 3.16a). The record of pre-3500 Ma continental rocks in the East Pilbara Terrane has previously been inferred from findings of rare gneissic xenoliths, and from the occurrence of 3700 to 3500 Ma detrital zircon grains in younger sediments (e.g. Nelson, 2000; Hickman and Van Kranendonk, 2012; Kemp et al, 2015). The pre-3500 Ma detrital zircon grains are widespread in the East Pilbara Terrane, and interpreted as derived from local granitic complexes (Kemp et al., 2015). Accepting a chondritic model of the early Archaean East Pilbara Terrane upper mantle, crustal sources from which the detrital zircon host rocks were derived, were extracted from the mantle between ca. 3700 to 3600 Ma based on zircon Hf isotopic data (Fig. 3.16a; Kemp et al., 2015). Although different mantle models are used, both Hf in zircon and Nd bulk-rock model ages suggest contemporaneous formation of mantle-derived mafic crust and occurrence of early TTG melts between ca. 3700 and 3500 Ma (Smithies et al., 2003, 2009; Kemp et al., 2015). If the age- frequency distribution of the >3500 Ma detrital zircon directly reflects the volume of continental crust formation (Fig. 3.16a,b), it can be inferred that granitic melt production started at ca. ≥3700 Ma and progressively increased culminating in a first major peak between ca. 3500-3600 Ma that coincided with the emplacement of the Doolena Suite gneisses. As inferred from the mantle model ages, the increasing production of continental material occurred contemporaneous, or directly

94 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______subsequent to long-lasting mantle-derived magmatism that produced source rocks for the TTG and granitic gneisses. As demonstrated in the petrological-geochemical modeling, the parental TTG melts of the Doolena Suite gneisses formed at ~45-50 km depth. It is therefore proposed that vertical thickening of mantle-derived mafic crust directly controlled the generation of granitoid melts from the lower crustal base (Fig. 3.16b). The model is consistent with models of early Archaean TTG generation in the Barberton granite-greenstone belt (e.g. Clemens et al., 2006). Here, a model is envisaged, in which early continental magma did not form as single events, but rather continuously from ca. 3700 to 3450 Ma, as thickening of the crust proceeded through mantle-derived volcanism, and progressively more material reached the critical depth to allow partial melting and TTG parental melt production (‘zone of partial melting’; Fig. 3.16b). The possibility of continuous juvenile magma production in the early East Pilbara Terrane evolution has been recently put forth (Van Kranendonk et al., 2015). The timing of emplacement of the 3591-3499 Ma Doolena Suite (this study), and the established emplacements of the 3490-3460 Ma Callina and the 3450-3420 Ma Tambina Supersuites (e.g. Van Kranendonk et al., 2006), suggest that the early Archaean mid-crust of the East Pilbara Terrane has been almost continuously infiltrated with juvenile continental magma, generated from the lower crustal base due to over- thickening. The confirmed Strelley Pool Formation stratigraphic correlation with rocks in this study area constrains previously recognized dome-and-keel formation related deformation to between ca. 3460 and 3427 Ma (Chapter 2), interpreted as the earliest partial convective overturn in the East Pilbara Terrane. If continental melts were produced from ca. 3700-3600 Ma onwards, as reflected in the zircon record, this means that the crust gravitationally re-organized after reaching critical volume of mid-crustal granite and mafic volcanic supracrustal loading after ca. 200-300 My (Fig. 3.16b). As recently mentioned by Sizova et al. (2015), previous models of partial convective overturn have not taken into account the formation of granitic melts, but supposed the pre-existence of a felsic mid-crust in order to develop crustal-scale instabilities (e.g. Collins et al., 1998; Sandiford et al., 2004; Thébaud and Rey, 2013). The lowermost supracrustal Warrawoona Group greenstone succession in the East Pilbara Terrane is interpreted as representing ultra-mafic – mafic – felsic volcanic cycles, separated by sediments and unconformities. Considering the interpretation that the felsic volcanics are products of differentiation of underlying tholeiitic basalts and not associated with TTG, the Warrawoona Group may be interpreted as representing pulses of mafic mantle- derived magmatism, separated by hiatus, during which tholeiitic differentiation within the over- thickened crust occurred (e.g. Smithies et al., 2007; Van Kranendonk et al., 2015). The absence of early TTG-like surficial edifice is in agreement with recent numerical models that demonstrate that intermediate- to felsic layers may remain in the mid crust for tens of millions of years (e.g. Sizova et al., 2015). This supports the idea that a critical gravitational (i.e. Rayleigh-Taylor) instability required the contemporaneous accumulation of granitic melts in the mid crust and supracrustal loading through mafic volcanism over ca. 200-300 Ma, until rapid partial convective overturn occurred within less than 40 Ma (Fig. 3.16b).

95 Chapter 3 Petrogenesis and reworking of the ca. 3500-3590 Ma Doolena Suite gneisses ______

3.10 Conclusions

Based on new U-Pb zircon geochronological data, 3499-3591 Ma intermediate to felsic gneisses (Doolena Suite) along the southwestern margin of the Muccan Granitic Complex, Doolena Gap greenstone belt are identified. Upper Concordia ages yield 3499 ± 22 Ma for a quartz dioritic gneiss, 3576 ± 22 Ma for a granodioritic gneiss, 3591 ± 36 Ma for a tonalitic-trondhjemitic gneiss, and 3457 ± 110 Ma for a potassic granitic gneiss. Petrological and geochemical data, in combination with petrological and geochemical modeling suggests that TTG parental magma was derived through partial melting of deeply buried hydrous garnet amphibolite along a relatively normal geothermal gradient at the base of over-thickened (~45-50 km) mafic proto-crust. Geochemical trace element modeling favors polybaric fractional crystallization of the parental TTG magma to form dioritic-granodioritic fractions. Zircon provenance ages confirm the stratigraphic correlation of a quartz-sandstone, overlying an angular unconformity in the adjacent greenstone belt, with the ~3427-3310 Ma Strelley Pool Formation. The data indicate that pre-3500 Ma granitic material from the Muccan Granitic Complex was exposed and eroded until ~3427 Ma. This provides time constraints for a previously recognized deformation event that affected the greenstones below the unconformity and was possibly associated with syn-deformational leucosome formation during doming of the Muccan Granitic Complex. Petrological and geochemical data from the leucosomes indicate low temperature partial reworking of the host gneisses, possibly as a result of heat and volatile (e.g.

H2O) supply from voluminous TTG magmatism between 3450 and 3420 Ma within the dome center.

96 “…Und wenn er seine Nase mit der Hand zuhielt, Dann nahm er keine Gerüche mehr wahr, Bis er seine Nase wieder öffnete.” (Abu Bakr Ibn Tufail, ca. 1105-1185)

Chapter 4 ______

GEOCHEMISTRY OF THE ca. 3470-3460 Ma MOUNT ADA BASALT AND DUFFER FORMATION IN THE WESTERN DOOLENA GAP GREENSTONE BELT (EAST PILBARA TERRANE, WESTERN AUSTRALIA): IMPLICATIONS FOR EARLY ARCHAEAN UPPER CRUSTAL GROWTH

4.1 Introduction

The lithological and geochemical make-up of Archaean greenstone belts is key to understanding the formation and evolution of Earth’s early crust (e.g. Furnes et al., 2014; Kamber, 2015; Herzberg et al., 2010). Numerous late Archaean (3200-2500 Ma) greenstone belts show evidence for their formation as ophiolite-like or accretionary complexes in convergent plate-tectonic settings similar to those found on modern-day Earth (e.g. Polat and Kerrich, 1999; Shervais, 2005; Manikyamba et al., 2005). In contrast, early Archaean (3850-3200 Ma) greenstone associations are characterized by extensive craton-wide lateral correlation of volcanic formations, and show evidence for autochthonous deposition onto pre-existing continental basement (e.g. Van Kranendonk, 2011; Hickman and Van Kranendonk, 2012; Kröner et al., 2013; Kamber, 2015). It is proposed that the early Archaean greenstone belts formed in a distinct tectonic regime as a result of higher heat flow from the mantle (Kamber, 2015). Rapid vertical growth of mantle-derived mafic proto-crust during the early Archaean may have resulted in over-thickening and subsequent partial re-melting of the mafic crust to produce differentiated mid-crustal layers (e.g. Bédard et al., 2006; Smithies et al., 2009; Johnson et al., 2014; Kamber, 2015; Chapter 3). Continuous predominantly dense, mafic greenstone volcanism on top of the partly differentiated buoyant basement is proposed to have led to partial convective overturn responsible for the development of characteristic granite dome - greenstone keel crustal structures (e.g. Chardon et al., 1996; Collins et al., 1998; Van Kranendonk 2011; Thébaud and Rey, 2013; François et al., 2014; Van Kranendonk et al., 2015). However, there is still no general consensus on the mechanisms of formation of early Archaean greenstone, and within individual terranes greenstone belts likely

Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______experienced unique evolutionary paths (e.g. de Wit and Ashwal, 1995; Furnes et al., 2014; Kamber, 2015). The occurrence of high-MgO basalts and komatiites indicate high degrees of partial mantle melting, which suggest crustal growth in association with hot mantle upwellings in the form of mantle plumes or heat-pipes (e.g. Kamber et al., 2005; Shervais, 2005; Smithies et al., 2005, 2007; Moore and Webb, 2012). On the other hand, it remains unclear if higher temperatures inferred from high-MgO basalts (e.g. Herzberg et al., 2010; Sobolev et al., 2016) are representative of overall elevated mantle temperatures during the early Archaean. Often, komatiites are absent or subordinate, and greenstone successions are dominated by low-MgO tholeiitic basalts in association with significant volumes of felsic volcanic rocks (Furnes et al., 2013; Condie, 2016). The predominance of low-MgO basalts may suggest partial melting from rather normal temperature ambient mantle (e.g. Murphy et al., 2015, abstr.; Bédard, 2006). The prevalence of low-MgO basalts and the minor contribution from high temperature melts such as komatiites indicates variable thermal conditions in the mantle, which could represent fluctuating thermal regimes or mixing of hotter upwellings with cooler upper mantle domains (e.g. Smithies et al., 2007; Kamber, 2015). On the other hand, intra-crustal magma chambers may have played a significant role in the derivation of low-MgO basalt from komatiitic/high-MgO magma through fractional crystallization (e.g. Smithies et al., 2007; Van Kranendonk et al., 2015), or may have supported the interaction, and hence contamination, with pre-existing crust to derive compositional types (e.g. Green et al., 2000). Furthermore, early Archaean greenstone deposition coincided with spatially associated tonalitic-trondhjemitic-granodioritic (TTG) plutonism within granitic domal complexes (Van Kranendonk et al., 2006; Kamber, 2015). The classical view of Archaean greenstone volcanism invokes a bimodal mafic-felsic system, in which mafic volcanics represent mantle derived rocks, and felsic volcanics represent eruptions of the contemporaneous crustal-derived TTG magmatism (Kröner et al., 2013; review Kamber, 2015). Contrastingly, felsic volcanic rocks from the most ancient greenstone volcanic cycles of the early Archaean East Pilbara Terrane (Warrawoona Group, Western Australia) have been suggested to be geochemically distinct to contemporaneous TTG plutonic rocks, until within the uppermost 3450-3427 Ma Panorama Formation of the Warrawoona Group (e.g. Smithies et al., 2007, 2009). Instead, it is proposed that the early felsic volcanic rocks represent products of tholeiitic differentiation of underlying basalts and komatiites (e.g. Smithies et al., 2007). The recognition of an early partial convective overturn event within the western Doolena Gap greenstone belt that initiated (i.e. extension and uplift) during deposition of the ca. 3470-3460 Ma bimodal Mount Ada Basalt - Duffer Formation succession (Chapter 2 and 3) highlights the importance to re-visit the petrogenetic affinity of the felsic and mafic volcanic rocks in the Doolena Gap greenstone belt. Considering that i) tectonic uplift due to Muccan Granitic Complex doming occurred during Duffer Formation deposition (Chapter 2), ii) granitic doming was associated with partial anatexis of older TTG components (Chapter 3), and iii) continuous TTG magmatism of the 3450-3420 Ma Tambina Supersuite affected the central parts of the dome (e.g. Van Kranendonk et al., 2006), it is likely that TTG and granitic magma genesis

98 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.1: a) Simplified geological map of the granite dome - greenstone keel East Pilbara Terrane (modified after Van Kranendonk et al., 2006); the rectangle indicates the position of the map in Figure 4.1d; b) location of the East Pilbara Terrane in Western Australia, Australia; c) generalized stratigraphy of the 3.53-3.23 Ga Pilbara Supergroup, including the Mount Ada Basalt and Duffer Formation of the ~3.47-3.46 Ga Coongan Supgroup (modified from Hickman, 2012); age ranges of the emplacement of granitic supersuites are from Van Kranendonk et al., 2006; d) geological map of the Doolena Gap - Marble Bar greenstone belt syncline (taken from Chapter 2); the studied Mount Ada Basalt (green) and Duffer Formation (yellow) are highlighted; Geochronological data from the Marble Bar greenstone belt is from: (1) Thorpe (1991), written comm. in Williams (1999), (2) McNaughton et al. (1993), (3) Nelson (1999), (4) Beintema (2003), (5) deVries et al. (2006); the rectangle indicates the position of the study area along 1 Mile Creek, shown in Figure 4.2. contributed to the Duffer Formation volcanism and influenced the mafic volcanics and intrusions in the Doolena Gap greenstone belt. The purpose of the present contribution is to discuss the petrogenesis and petrological- geochemical effects of intra-crustal processes, such as fractional crystallization and crustal contamination, and post-magmatic alteration, during upper crustal formation and evolution of the early Archaean Doolena Gap greenstone belt, East Pilbara Terrane. To this end, new geochemical data from the ca. 3470-3460 Ma Mount Ada Basalt - Duffer Formation bimodal greenstone succession are presented. The studied greenstone transect is characterized by i) absence of rocks that can be conclusively identified as komatiites, ii) field evidence for intra-crustal intrusions, iii) proposed eruption through precursor >3500 Ma continental crust, and iv), at least locally intensive

99 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______polyphase carbonate alteration (Chapter 2 and 3). In this way, it is aimed to answer [1] to what extent post-magmatic alteration affected the greenstones, [2] if the greenstone compositions were significantly affected by fractional crystallization and/or crustal contamination, and [3] if felsic volcanics were derived from underlying basalts, or TTG.

4.2 Brief Geology and Previous Geochemical Constraints on the Warrawoona Group Greenstone Volcanic Rocks

The greenstone belts of the early Archaean East Pilbara Terrane (Western Australia) form typical curvilinear synclinoria wrapping around polyphase granitic dome complexes (Fig. 4.1a,b), and comprise volcano-sedimentary successions of the 3530-3225 Ma Pilbara Supergroup (Fig. 4.1c). The 3530-3427 Ma Warrawoona Group includes the lowermost greenstone successions of the Pilbara Supergroup and is separated from the overlying 3350-3300 Ma Kelly Group by the Strelley Pool Formation, marking a regional unconformity (Fig. 4.1c; Hickman, 2012). The lower base of the Warrawoona Group is not exposed due to either tectonic or intrusive contacts with granitic rocks of the dome complexes. The Warrawoona Group consists of repeating (ultra-) mafic to felsic volcanic cycles that are stratigraphically represented, from bottom to top, by the Coonterunah, the Talga Talga, the Coongan and the Salgash Subgroups (Fig. 4.1c,d; Hickman, 2012). The most frequently cited geochemical data from the Warrawoona Group are based on regional sampling accomplished by the Geological Survey of Western Australia (e.g. Smithies et al., 2007). Particularly data from the Mount Ada Basalt and Duffer Formation (Coongan Subgroup; Fig. 4.1c), which are the focus of this study, come from sampling transects within the northern arm of the Marble Bar greenstone belt that represents the opposite synclinal limb to the studied Doolena Gap greenstone belt (Fig. 4.1d). Ultra-mafic komatiites (18-30 wt.% MgO) and komatiitic basalts (12-18 wt.% MgO) are mostly reported from the lower parts of the Warrawoona Group volcanic cycles (e.g. Green et al., 2000; Smithies et al., 2005; 2007). Most komatiitic lavas of the Warrawoona Group show flat normalized trace element patterns with values of about three to six times of that of chondrite, and belong to the Al-undepleted/Munro- (Nesbitt et al., 1979) type with Al2O3/TiO2 of ~20 to 24

(Smithies et al., 2007). However, Al-depleted komatiites (Al2O3/TiO2 ≈ 7), characterized by strongly depleted chondrite-normalized LREE have been reported (Gruau et al., 1987). Tholeiitic basalts form the most voluminous rock type in the Warrawoona Group volcanic cycles, and have been subdivided into high-Ti (>0.8 wt.% TiO2) and low-Ti groups, which occur interlayered and are indistinguishable in the field (Smithies et al., 2005; 2007). Both groups overlap in major-element compositions, however, the more common high-Ti basalts are characterized by higher iron content, higher concentrations in HFSE and overall REE, lower

Al2O3/TiO2 (~ 9 to 19), and display chondrite-normalized HREE depletion (Gd/Yb ~ 1.12-2.23). Smithies et al. (2005) interpreted the derivation of high-Ti basalts from Al-depleted (‘Barberton- type’) komatiites and komatiitic basalts formed during high-pressure melting in hot mantle plumes, with no substantial change in composition over time. In contrast, the volumetrically

100 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.2: Geological-structural map of the studied 1 Mile Creek transect with locations of samples used in the present study; map position in the western Doolena Gap greenstone belt is indicated in Figure 4.1d. subordinate low-Ti basalts show a systematic temporal chemical change with increasing in

Al2O3/TiO2 and Mg#, decreasing incompatible element concentrations, and decreasing La/Sm and Gd/Yb progressively up stratigraphy. Smithies et al. (2005) argued that the systematic variation of the low-Ti basalts reflect successive extraction from a progressively depleted mantle source that developed in isolation of the high-Ti plume source.

101 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

A continuous debate regarding the observed compositional variety of basaltic rocks from the Warrawoona Group is the possibility of crustal contamination. Green et al. (2000) proposed up to 25% contamination of even the lowermost Warrawoona basalts (Coonterunah Subgroup) by felsic crust. Arndt et al. (2001) and Bolhar et al. (2002) argued against significant contamination, and Smithies et al. (2007) noted that the range in LREE concentrations and La/Sm and La/Yb ratios of some of the basalts might indicate varying degrees of interaction with felsic crust. In a subsequent study Smithies et al. (2009) argued that incompatible element enrichment (Th, U, Nb, Zr, LREE) in some of the basalts from the lowermost Warrawoona Group could be inherited from an enriched mantle source, less depleted than modern N-MORB and enriched in recycled crustal components (Smithies et al., 2009). The major felsic volcanic formations at the tops of the Warrawoona Group volcanic cycles have been mostly ascribed to mafic-felsic differentiation series (Smithies et al., 2007). Based on the GSWA data (Smithies et al., 2007), only within the uppermost felsic succession of the Duffer Formation is a major distinctly trace element enriched group observed, but these volcanics are less L/HREE fractionated than average TTG. Despite this, it has been argued that volcanic rocks of the Duffer Formation represent surficial expressions of the TTG domes (e.g. Bickle et al., 1983; DiMarco and Lowe, 1989). TTG-like volcanics occur in the Panorama Formation, where they represent a minor component (Smithies et al., 2007). The Panorama Formation is the uppermost sub-group of the Warrawoona Group (Fig. 4.1c) and forms distinct volcanic centers each characterized by distinct compositions (e.g. DiMarco and Lowe, 1989; Smithies et al., 2007).

4.3 The Doolena Gap Greenstone Belt Transect: Field Relationships and Sample Selection

The presently studied rocks were sampled along a geological transect within the western Doolena Gap greenstone belt, previously ascribed to the Mount Ada Basalt and Duffer Formation (e.g. Van Kranendonk, 2010; Fig. 4.1c,d; Fig. 4.2). The stratigraphic affinity is not conclusive and based on regional mapping and correlation with geochronological constraints mainly from the southerly adjacent Marble Bar greenstone belt (Fig. 4.1d; Van Kranendonk, 2010). Within the study area, only the quartz-sandstone that unconformably overlies the proposed Duffer Formation and Mount Ada Basalt has been confirmed as representing the 3427-3350 Ma Strelley Pool Formation (e.g. Hickman, 2008; Chapter 3). Furthermore, Sm-Nd and Lu-Hf isochrones of basaltic rocks from the Low-Strain Belt slightly to the west of the here studied sampling transect yielded ages (within error) in agreement with the accepted Mount Ada Basalt depositional age (3470 Ma; Murphy et al., 2015, abstr.). In combination with the structural model proposed in Chapter 2, the above geochronological constraints support the stratigraphic correlation with the Mount Ada Basalt and Duffer Formation in the southern Low-Strain Belt of the study area. The stratigraphic affinity of rocks from the more highly deformed South Muccan Shear Zone and Central Fold Belt, however, remains unclear (Fig. 4.2; Chapter 2). Therefore, field relationships and analytical results from the

102 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.3: Representative field photographs; a) meta-basalt from the Mount Ada Basalt, Low-Strain Belt, showing well-preserved pillow structures; b) massive rock from the Duffer Formation, Low-Strain Belt, showing quartz-filled vesicles. different structural domains will be presented separately first, and possible stratigraphic and petrogenetic relationships considered in the Discussion of this Chapter (Section 4.8).

4.3.1 Low-Strain Belt

The Low-Strain Belt in the east of the mapped transect represents a relatively continuous and little deformed stratigraphic section dominated by pillow basalts of the Mount Ada Basalt (Fig. 4.3a) in the stratigraphically lower northern part, conformably overlain by silicic volcanic rocks (Fig. 4.3b) and volcaniclastic sediments of the Duffer Formation to the south. The pillow basalts of the Mount Ada Basalt often comprise abundant ocelli. Vesicular flows are rare, and mostly observed above chert horizons, and within the uppermost Mount Ada Basalt. No intrusions that crosscut the pillow flows are observed. However, some more massive, spinifex-textured mafic units occur that lack

103 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______the development of pillows. The latter rocks do not show any palisade-like alignment of the spinifex-texture, which would indicate the presence of flow tops. The spinifex-textured units are always overlain by pillow basalt flows, but direct contact relationships are obscured by alteration. Therefore it remains unclear if these rocks represent massive flows or sub-volcanic sills. One massive, medium- to coarse-grained anorthositic-gabbro forms a bedding-parallel intrusion or sill in the southwestern Low-Strain Belt. The studied silicic volcanic rocks of the Duffer Formation come from the southeastern part of the Low-Strain Belt. The lower contact to the Mount Ada Basalt is conformable, but disturbed where the NE-SW trending D4 faults sinistrally offset the stratigraphy (Chapter 2), and often marked by a few repeating cycles of interlayered vesicular pillow basalts and silicic volcanic layers. The emplacement style of the silicic volcanic layers is unclear, and they may represent either volcanic flows, or sub-volcanic sills that intruded the uppermost pillow basalts during the onset of the Duffer Formation volcanism. In fact, emplacement styles of the silicic volcanic rocks are ambiguous throughout most of the Duffer Formation. Only some aphanitic volcanic units that are characterized by abundant and large vesicles display a volcanic flow texture. The volcanic flows are interlayered with medium-grained volcanic rocks. Importantly, however, no crosscutting intrusive relationships are observed within the sample-transect.

4.3.2 Central Fold Belt and South Muccan Shear Zone

The Central Fold Belt to the north is affected by multiple fold generations (Fig. 4.2; Chapter 2), and mostly comprises greenschist-facies mafic schists that tectonically incorporate relatively undeformed lozenges within the regional EW S2-foliation. Only the relatively intact lozenges preserve primary magmatic textures. Overall, more massive, as well as distinctly intrusive mafic lozenges increase in abundance towards the Muccan Granitic Complex. Most of the contact relationships and the relative timing and style of emplacement are unclear due to the transposed nature of the lozenges (Chapter 2). The presence of lozenges of vesicular and pillow basalts in association with lozenges of former intrusive rocks (e.g. gabbro, pyroxenite), suggests that the latter intrusive rocks were emplaced into the pillow basalt flows. Importantly, all rocks within the

Central Fold Belt investigated here are affected by the D2 formation that occurred prior to 3427 Ma, during which the lozenges developed. Thus the Central Fold Belt lavas and intrusions are consistent with being part of the Warrawoona Group. Rocks of clearly intrusive origin include gabbro, clino-pyroxenite, anorthositic gabbro, and anorthosite. In the northern part of the Central Fold Belt one larger lozenge (100 m-scale) represents an intrusive body characterized by a compositional zoning of pyroxenite-anorthosite- anorthositic gabbro to gabbro. Other transposed lozenges of gabbro and anorthosite occur throughout the central to northern part of the Central Fold Belt. The emplacement style of abundant medium-grained, massive, non-pillowed mafics within the northwestern Central Fold Belt is unclear. The rocks may represent either the interiors of massive flows, or sub-volcanic sills.

104 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Table 4.1: Sample list with observed or inferred primary magmatic mineralogy and secondary metamorphic/metasomatic assemblage Sample Formation (proposed) magmatic mineralogy metamorphic assemblage text. preserv. type Low-Strain Belt 34-1 Mount Ada Basalt (Cpx-Pl-Ol?) Chl-clay-Ser±Qz(-Opq) poor ocelli pillow basalt 182-1 Mount Ada Basalt (Cpx-Pl-Ol?)-Opq Chl-Cc±Qz±Ab(?) poor spinifex mafic rock 307-2 Mount Ada Basalt Pl(-Cpx)-Opq Chl-Qz-Cc-clay(-Opq) poor ocelli pillow basalt 365-1 Mount Ada Basalt (Cpx-Pl-Ol?)-Opq Chl-clay±Qz±Ab(-Opq) good spinifex mafic rock 30-1 Mount Ada Basalt Pl-Cpx-Ilm Chl-Ab-clay-Opq(Rt)±Qz(±Ser?) good pillow basalt 38-1 Mount Ada Basalt Chl-clay-Opq(+Rt?)±Qz±Ab no pillow basalt 40-1 Mount Ada Basalt Pl(-Cpx?-Opq) Chl-clay-Ab-Qs-Opq(±Cc?) poor pillow basalt 48-1 Mount Ada Basalt Pl-Cpx-Opq Chl-Ab-clay±Qz fair pillow basalt 216-1 Mount Ada Basalt Pl-Cpx-Opq Chl-Ab-Qz poor pillow basalt 259-1 Mount Ada Basalt Pl-Cpx-Opq Chl-Qz(±Ab?) no/very poor pillow basalt 252-1 Mount Ada Basalt Pl+/-Cpx+/-Opq Ab±Chl±Ser fair anorthositic-gabbroic 186-2 Duffer Formation (Cpx?-Pl-Qz-Opq) Qz-Ser-clay-Chl fair unknown 175-1 Duffer Formation (Fsp-Qz-Opq) Qz-Ser-Kaol/clay-Rt fair unknown 356-1 Duffer Formation (Fsp-Qz-Opq) Qz-Ser-Kaol/clay-Rt fair unknown 357-2 Duffer Formation (Fsp-Qz-Opq) Qz-Ser-Kaol/clay-Rt fair unknown 109-1 Duffer Formation (Fsp-Qz-Opq) Qz-Ser-Kaol/clay-Chl? fair unknown 338-1 Duffer Formation (Fsp-Qz-Opq) Qz-Ser-Kaol/clay-Rt fair unknown 112-1 Duffer Formation (Fsp-Qz-Opq) Qz-Ser-Kaol/clay-Rt fair unknown Central Fold Belt 315-1 undefined* Cpx-Pl-Opq Chl-Qz-clay poor/fair massive mafic 346-1 undefined* Cpx-Pl-Opq Chl±Ep±Qz±Ab poor massive mafic 359-1 undefined* Ol pheno. (±Cpx-Pl?) Chl-clay±Qz good mafic dyke 359-2 undefined* (Cpx-Pl-Opq?) Chl±Qz±Ab very poor massive mafic 360-1 undefined* (Cpx-Pl-Opq?) Chl-clay±Qz±Ab no massive mafic 371-1 undefined* Cpx-Pl-Opq Chl-Cc±clay±Ab good massive mafic 234-1 undefined* Chl-Ep-Opq±Qz±clay no massive mafic 348-1 undefined* Cpx-Pl-Opq Chl-clay-Ser±Ab good massive mafic 358-1 undefined* Pl(-Cpx)-Opq Chl-clay±Ser±Ab fair massive mafic 364-1 undefined* Pl(-Cpx)-Opq Chl-Cc±Ab±Qz(-Opq) no massive mafic 318-1 undefined* Pl-Cpx-Opq Chl-Ep±clay±Qz(-Opq) fair massive mafic 347-1 undefined* Pl-Cpx-Opq Chl±Ep±Ab(-Opq) good massive mafic 99-5 undefined* Pl-Cpx±Opq Chl-Ab±Opq fair gabbro 97-2 undefined* fair gabbro 320-2/3 undefined* Pl±Opq(±Cpx) Ab-Cc-Ser/Ms-±Opq±Chl fair/good anorthosite 100-1 undefined* Cpx(±Pl)-Opq Chl fair clino-pyroxenite South Muccan Shear Zone 330-1/2 undefined* (chromite-Ol?) chromite?-serpentine no ultra-mafic schist 311-1 undefined* Hbl/Act-Chl-Ep-Qz-Opq no amphibolite schist 312-1 undefined* Hbl/Act-Chl-Qz-Opq±Ep no amphibolite schist Note: phases in brackets are inferred; Ab, albite; Act, actinolite; Cc, calcite; Cpx, clinopyroxenite; Ep, epidote; Fsp, feldspar; Hbl, hornblende; Kaol, kaolinite; Ol, olivine; Opq, opaque; Pl, plagioclase; Qz, quartz; Rt, rutile; Ser, sricite;

Within the South Muccan Shear Zone original field relationships are completely obscured by the tectonic imbrication of lozenges of various rock types (Chapter 2). The South Muccan Shear Zone comprises granitoid and amphibolite schists and mylonites from both the Muccan Granitic Complex and the greenstone belt. The first morphological ridge south of the Muccan Granitic Complex is formed by ultramafic schists/mylonites.

4.3.3 Sample selection

Sampling focused on the best-preserved and least carbonate-altered specimens. Particularly within the highly deformed Central Fold Belt sampling was concentrated on relatively continuous transects, not cut by major faults or shear zones (Fig. 4.2). One such transect has been identified in the northwestern Central Fold Belt (Fig. 4.2) that represents the limb of a map-scale F2 fold. Here, fault displacements and stratigraphic repetitions possibly had minor effects on the overall stratigraphic continuity. The selected samples are listed in Table 4.1, and sample locations are shown in the map of Figure 4.2. Although all studied rocks show at least partial sub-greenschist- to greenschist-facies overprinting, the prefix ‘meta-‘ is omitted for simplicity. For petrographic and geochemical

105 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.4: Selected petrographic images of representative textures of rocks from the Low-Strain Belt; refer to text for description. analyses, the following samples were chosen: i) seven silicic volcanic rocks come from the Duffer Formation and include both volcanic flows and samples from units of undefined emplacement style (186-2, 175-1, 356-1, 357-2, 109-1, 338-1, 112-1), ii) a total of eleven mafic rocks from the Mount Ada Basalt, including six pillow basalts (30-1, 38-1, 40-1, 48-1, 216-1, 259-1), two ocelli- bearing pillow basalts (34-1, 307-2), two spinifex-textured massive mafic rocks (182-1, 365-1), and one anorthositic-gabbroic mafic rock (252-1), iii) a total of 17 mafic rocks from the Central Fold Belt, including 10 massive mafic rocks (315-1, 346-1, 359-2, 360-1, 371-1, 348-1, 358-1, 364-1, 318-1, 347-1), one mafic dyke (359-1), one mafic schist (234-1), two gabbros (99-5, 97-2), two anorthosites (320-2, 320-3), and one clino-pyroxenite (100-1), and iv) two ultra-mafic schists (330-1, 330-2) and two amphibolite schists (311-1, 312-1) from the South Muccan Shear Zone.

106 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

4.4 Analytical Methods

Petrographic analyses were performed using a Leica DM750P polarizing microscope, fitted with a Leica ICC50HD camera system for image acquisition. For geochemical analyses, a total of 38 rock samples (29 mafic and 7 silicic) were sledgehammered to gravel-size in heavy-duty sample bags, and then crushed to sand-size in a tungsten-free rock-crusher. Powdering of the samples was performed with an agate mill (Queensland University of Technology). For major element X-ray fluorescence (XRF) analyses method and instrumentation refer to Chapter 3, Section 3.5. Because of possible F- and carbonate- complex formation using dilution- ICP-MS technique, trace element analyses were performed with a LA-ICP-MS (Agilent 8800 triple quadrupole ICP-MS, Central Analytical Research Facility, Queensland University of Technology) on fusion beads. The beads were manually fused from a 1:3 sample to flux (1:1.5 lithium metaborate-lithium bromide) mixture. Each fusion bead was ablated five times using a 193 nm excimer laser and a 100 µm-diameter spot size. The glass standards NIST610 and NIST612 (Jochum et al., 2011) were analyzed repeatedly, bracketing every sequence of 10 unknowns, throughout the run. LA-ICP-MS data reduction, including background correction and integration of the standards, was performed with the IgorPro based Iolite software (WaveMetrics). The LA- ICP-MS data were then corrected using the Si content derived from XRF major element analyses. The data used here (Table 4.3 and 4.5) represents average values of five ablation spots for each sample.

4.5 Petrography and Geochemistry of Rocks of the Duffer Formation and Mount Ada Basalt from the Low-Strain Belt

4.5.1 Petrography

4.5.1.1 Mount Ada Basalt

The pillow basalts of the Mount Ada Basalt often preserve magmatic textures (Table 4.1) characterized by pseudomorphic laths of former igneous minerals that are replaced by sub- greenschist to greenschist-facies minerals. The most well-preserved pillow basalts comprise a phaneritic texture of interlocking former medium-grained eu- to subhedral clinopyroxene and plagioclase laths, and opaque minerals (e.g. 30-1, 48-1; Fig. 4.4a). Clinopyroxene is often completely replaced by fine-grained or microcrystalline clay minerals, and plagioclase is mostly albitized, or replaced by fine-grained albite and chlorite. Intermediately preserved pillow basalts display a porphyritic to aphanitic texture characterized by fine-to medium grained thin laths of former clinopyroxene and plagioclase, now replaced by albite, chlorite, and clay minerals, set in a finely recrystallized matrix of mostly chlorite, quartz and clay minerals (e.g. 40-1). Basalts comprising a more poorly preserved magmatic texture, are either characterized by more abundant former aphanitic groundmass relative to the abundance of

107 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______phenocrystic mineral laths and are from pillow margins, or are found in the vicinity to carbonate alteration zones. In the latter case, the basalts occasionally comprise multiple generations of thin crosscutting quartz, chlorite and calcite vein networks. Secondary quartz often forms medium- grained aggregates overgrowing the matrix. Within the two ocelli-bearing samples, magmatic textures are rather poorly preserved, and only within certain domains, particularly at the rims of the ocelli, former magmatic laths can be recognized. The samples comprise an alteration assemblage of microcrystalline to fine-grained chlorite, quartz, albite, and clay minerals. One of the samples (34-1) is characterized by a smaller ocelli size (~5 mm) compared to the other (10 mm), and contains former phenocrysts. The primary mineralogy of the phenocrysts is unclear due to the complete pseudomorphic replacement by chlorite and clay minerals, but their crystal shapes resemble that of olivine. The phenocrysts are up to ~500 µm in size and show embayments along their margins. The other ocelli-bearing basalt is interlayered with the first silicic volcanics of the Duffer Formation and contains abundant vesicles. The vesicles are filled by secondary inward growing chlorite fibres and/or calcite and microcrystalline quartz (Fig. 4.4b). The ocelli are richer in clay minerals compared to the chlorite- dominated matrix (Fig. 4.4b). In both the spinifex-textured samples, former magmatic laths of up to 1.5 cm size are randomly oriented, and are inferred to represent former clinopyroxene, now replaced by chlorite and clay minerals. In the more well-preserved of the two samples (365-1), chlorite cores and clay mineral rims of the spinifex-type laths indicate pseudomorphs after an original compositional zoning (Fig. 4.4c). The finely recrystallized matrix of this rock comprises tiny laths of chlorite and albite, likely representing a former plagioclase and clinopyroxene rich aphanitic groundmass. Quartz and minor calcite form part of the alteration assemblage within the matrix. Within the more altered spinifex- textured basalt (182-1), which occurs in the vicinity of a carbonate alteration zone, the spinifex laths are poorly preserved in thin section and completely replaced by chlorite and minor microcrystalline calcite. The finely recrystallized chlorite-albite-quartz-clay mineral matrix comprises patches of fine- to medium grained calcite and chlorite. Some domains within the matrix are rich in tiny laths of parallel-aligned opaque minerals, representing the remnants of replacement textures, or secondary growth. The anorthositic-gabbroic mafic rock (252-1; Fig. 4.4d) comprises a seriate interlocking texture of eu-to subhedral coarse-grained blocky plagioclase (>500 µm) that is mostly albitized, and medium-grained laths of former plagioclase that occur in association with opaque minerals and former medium-grained clinopyroxene that is replaced by chlorite and clay minerals.

4.5.1.2 Duffer Formation

The silicic volcanic rocks from the Duffer Formation preserve aphanitic to porphyritic primary magmatic textures. However, the primary mineralogy is completely replaced by low-grade alteration products, mostly consisting of quartz-sericite-ilmenite/rutile-kaolinite assemblages. From the alteration assemblage it can be inferred that the primary magmatic mineralogy likely

108 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Table 4.2: XRF major element data for mafic and silicic rocks from the Low-Strain Belt Mount Ada Basalt Duffer Formation sample 365-1 34-1 307-2 182-1 216-1 30-1 48-1 40-1 259-1 38-1 252-1 186-2 175-1 356-1 357-2 109-1 338-1 112-1 (wt.%) spfx ocel. ocel. spfx pb pb pb pb pb pb an-gb SiO 2 48.72 44.86 50.91 37.96 50.37 48.17 57.19 47.87 45.50 45.99 52.13 79.54 77.00 74.15 80.61 79.29 81.11 78.82 TiO2 0.57 0.98 1.30 0.76 0.87 1.18 1.00 0.99 1.43 0.88 1.27 1.30 1.60 1.64 1.36 1.40 1.43 1.76 Al2O3 14.33 10.30 11.86 7.51 14.44 12.93 14.36 14.03 12.14 13.14 12.97 10.27 12.62 14.24 11.18 12.01 11.35 12.46 Fe2O3 11.90 17.69 17.69 12.86 11.33 13.39 12.20 13.78 15.75 10.67 13.06 1.08 1.13 0.70 1.04 0.73 0.58 0.80 MnO 0.13 0.33 0.08 0.31 0.20 0.18 0.09 0.25 0.23 0.13 0.17 0.02 0.01 0.01 0.01 0.01 0.01 0.01 MgO 11.60 12.05 10.01 10.84 7.81 5.85 6.81 6.38 5.00 7.89 6.34 0.55 0.76 0.47 0.40 0.49 0.20 0.27 CaO 3.43 8.44 1.26 16.24 5.61 7.65 0.80 10.43 8.39 7.65 7.25 0.18 0.19 0.02 0.01 0.11 0.02 0.03 Na 2O 1.25 0.44 0.07 0.06 4.09 0.08 0.06 1.86 2.59 0.07 3.11 0.11 0.09 0.12 0.09 0.09 0.11 0.11 K2O 0.53 1.79 0.58 0.02 0.94 1.52 2.48 0.06 0.02 2.47 1.76 5.77 4.57 7.00 3.08 4.06 3.60 3.96 P 2O5 0.05 0.10 0.11 0.08 0.09 0.08 0.09 0.08 0.11 0.09 0.15 0.14 0.17 0.04 0.04 0.11 0.08 0.04 LOI 7.50 3.04 6.14 13.36 4.26 8.96 4.93 4.27 8.85 11.02 1.79 1.05 1.86 1.61 2.18 1.71 1.53 1.76 Total 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 Mg# 79 73 69 77 73 63 69 65 56 75 66 67 73 73 61 73 57 57 Mg# calculated as Mg/(Mg+Fe) an-gb, anorthositic-gabbroic; ocel., ocelli pillow basalt; pb, pillow basalt; spfx, spinifex-textured included K-rich feldspar, quartz and ilmenite. Ilmenite forms either tiny needles, or skeletal remnants, and is often partly replaced by rutile. Where present, the former phenocrysts are medium-grained and form eu- to subhedral laths that often show marginal embayments. The rock matrix consists mostly of very fine-grained/microcrystalline secondary quartz, sericite and clay minerals (i.e. kaolinite). The original texture of the fine-grained recrystallized matrix is unclear, but likely represented an aphanitic groundmass due to the fact that phenocrystic textures are preserved. Overall, the preserved magmatic textures show a gradation from medium-grained porphyritic to fine-grained aphanitic (Fig. 4.4e,f). The more finer-grained rocks comprise abundant vesicles that are filled by microcrystalline quartz (Fig. 4.4e; Fig. 4.3b). The size of the vesicles increases with decreasing abundance and grain-size of former phenocrysts. The most fine-grained specimens comprise large vesicles that show oval shapes, surrounded by a microcrystalline, recrystallized matrix (quartz-sericite-clay) reminiscent of a magmatic trachytoid (i.e. flow) texture. The two samples from the lowermost Duffer Formation are interlayered with the uppermost vesicular pillow basalts of the Mount Ada Basalt and contain occasional magmatic-textured fine-grained and often radially aligned laths consisting of microcrystalline clay minerals. The laths resemble the pseudomorphs after clinopyroxene that are inferred from the underlying basalts.

4.5.2 Major and trace element geochemistry

XRF major element data is given in Table 4.2 and LA-ICP-MS trace element data is given in Table 4.3. Presented rare earth element (REE) diagrams and values (indicated by ‘N’) in the text are chondrite-normalized (McDonough and Sun, 1995), and presented multi-element spider- diagrams are primitive mantle-normalized (McDonough and Sun, 1995). As shown in the AFM diagram (Fig. 4.5a), the bulk of the mafic rocks from the Mount Ada Basalt plot along a tholeiitic trend. The anorthositic-gabbroic rock (252-1) straddles the transition towards calc-alkaline, and the more highly altered of the spinifex-textured rocks (182-1) plots well within the field for the calc-alkaline differentiation series. All of the silicic volcanic rocks from the

109 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Table 4.3: LA-ICP-MS trace element data for mafic and silicic rocks from the Low-Strain Belt

Mount Ada Basalt Duffer Formation

sample 365-1 34-1 307-2 182-1 216-1 30-1 48-1 40-1 259-1 38-1 252-1 186-2 175-1 356-1 357-2 109-1 338-1 112-1 (ppm) spfx ocel. ocel. spfx pb pb pb pb pb pb an-gb P 215.35 145.94 552.46 338.73 372.97 383.36 359.10 316.13 512.81 405.38 671.98 498.24 630.04 119.76 113.16 381.94 260.15 101.77 Sc 38.73 46.45 31.37 19.70 38.22 37.38 29.91 37.34 29.25 36.72 23.68 12.59 21.37 18.44 15.53 14.13 14.53 12.12 Cr 293.11 310.36 709.90 971.61 183.85 118.32 160.90 137.80 104.84 172.95 258.50 11.70 27.54 64.54 35.84 49.66 22.61 4.59 Ni 103.81 135.54 326.89 761.50 76.44 76.51 135.79 120.15 73.18 71.48 242.30 25.02 35.42 53.52 60.38 33.04 13.61 7.92 Cu 116.80 54.73 319.83 102.64 77.53 126.86 120.97 200.13 91.69 41.71 131.84 Rb 16.24 0.77 23.29 0.27 9.31 45.99 84.16 0.87 0.11 55.07 23.29 82.87 78.38 91.56 53.40 60.62 59.62 62.68 Sr 22.13 28.85 7.59 80.28 95.01 19.04 4.51 141.88 154.83 26.09 264.78 3.51 3.53 27.16 15.53 9.09 48.19 9.25 Y 14.77 12.24 21.72 13.57 22.73 22.05 12.60 20.16 24.89 20.60 21.57 17.72 34.62 29.48 15.13 17.42 24.21 54.54 Zr 35.40 26.88 92.56 59.47 69.49 61.34 53.38 53.06 89.26 71.24 130.32 122.70 158.17 165.32 129.40 136.13 135.25 332.99 Nb 1.37 1.42 6.74 4.22 3.28 2.75 2.71 2.31 5.49 3.38 6.90 7.80 10.14 10.79 8.44 9.06 9.03 21.03 Ba 115.87 46.13 86.71 12.15 199.84 165.72 51.84 11.22 6.44 103.40 508.23 195.70 122.75 281.56 166.93 128.27 219.38 108.73 La 2.27 2.61 8.45 7.62 5.16 3.16 2.51 2.64 7.14 4.81 18.78 11.01 18.74 23.65 17.37 15.32 19.90 32.77 Ce 6.03 6.52 23.99 17.33 12.69 8.90 7.76 8.00 18.38 12.58 40.51 23.13 39.15 50.13 35.94 32.29 38.54 69.69 Pr 0.74 0.96 3.19 2.28 1.76 1.36 1.10 1.23 2.64 1.77 4.90 3.31 5.33 6.84 4.40 4.30 5.22 9.54 Nd 3.41 4.91 14.73 10.02 8.50 7.22 5.43 6.48 13.18 8.10 20.71 16.35 26.25 33.31 21.24 21.06 25.36 47.57 Sm 1.06 1.51 3.41 2.14 2.55 2.41 1.62 2.01 3.46 2.32 4.25 3.84 6.27 7.46 4.33 5.07 5.65 11.17 Eu 0.46 0.55 1.07 0.73 0.72 0.87 0.50 0.80 1.26 0.78 1.40 1.25 1.99 2.20 1.13 1.59 1.70 3.28 Gd 1.85 1.93 3.74 2.54 3.25 3.18 2.08 3.22 4.16 2.92 4.80 4.02 6.72 7.24 3.92 5.05 5.88 11.83 Tb 0.31 0.33 0.58 0.41 0.55 0.55 0.31 0.51 0.67 0.52 0.71 0.58 1.07 1.07 0.55 0.73 0.84 1.79 Dy 2.33 2.16 3.96 2.58 3.74 3.66 2.18 3.45 4.60 3.48 4.25 3.65 7.13 6.30 3.40 4.29 5.16 11.30 Ho 0.49 0.48 0.84 0.53 0.78 0.82 0.44 0.75 0.91 0.74 0.79 0.74 1.47 1.25 0.66 0.78 1.05 2.32 Er 1.65 1.33 2.44 1.45 2.45 2.62 1.42 2.31 2.80 2.29 2.27 2.05 4.21 3.44 1.88 2.14 2.85 6.60 Tm 0.25 0.18 0.35 0.20 0.37 0.35 0.22 0.31 0.38 0.34 0.31 0.30 0.62 0.49 0.27 0.29 0.41 0.99 Yb 1.70 1.18 2.05 1.26 2.51 2.44 1.55 2.25 2.79 2.29 2.07 1.98 4.29 3.31 1.76 1.88 2.75 6.52 Lu 0.26 0.19 0.33 0.17 0.34 0.34 0.21 0.30 0.41 0.35 0.30 0.31 0.63 0.49 0.27 0.28 0.42 0.97 Hf 0.94 0.81 2.41 1.49 1.82 1.74 1.47 1.50 2.41 2.06 3.38 3.60 4.64 4.85 3.75 3.88 3.89 9.33 Ta 0.08 0.09 0.44 0.25 0.20 0.16 0.18 0.14 0.35 0.22 0.54 0.64 0.81 0.96 0.72 0.71 0.76 1.69 Th 0.26 0.30 1.40 1.04 0.67 0.30 0.31 0.25 0.68 0.68 3.42 1.76 2.76 3.11 2.33 2.41 2.41 4.63 U 0.06 0.07 0.33 0.23 0.16 0.06 0.08 0.07 0.17 0.16 0.74 0.43 0.52 1.05 0.41 0.68 0.53 0.40 an-gb, anorthositic-gabbroic; ocel., ocelli pillow basalt; pb, pillow basalt; spfx, spinifex-textured

Duffer Formation plot close to the Na2O+K2O apex, along a calc-alkaline trend (Fig. 4.5a). The two formations will be geochemically characterized separately, below.

4.5.2.1 Mount Ada Basalt

Based on discrimination diagrams for ultramafic to mafic rocks, the Mount Ada Basalt samples comprise both tholeiitic and komatiitic basalts (Fig. 4.5b,c). However, considering the strict sense of the definition for komatiitic basalts (Arndt et al., 1997; Arndt and Lesher, 2004), and the fact that the most-well preserved spinifex-textured rock (365-1) does not plot in the komatiitic basalt field (Fig. 4.5b,c) the proposed discrimination will not be used here. Instead, as seen in the major element oxide Harker variations (Fig. 4.6a-j), using MgO as an index of differentiation, it is preferred to generically group the samples into high-MgO (>10 wt.%) and low-MgO (<10 wt.%) basalts (Fig. 4.6). The high-MgO group (10.0-12.1 wt.% MgO) comprises the two spinifex- textured (365-1, 182-1), and the two ocelli-bearing pillow basalt samples (34-1, 307-1). The low- MgO basalts (5.0-7.8 wt.% MgO) are all non-ocelli-bearing pillow basalts. The anorthositic- gabbroic mafic rock (252-1; 6.3 wt.% MgO) falls within the low-MgO group. The SiO2 content of the two groups overlap, but displays an overall slight increase from the high-MgO (38.0-50.9 wt.%

SiO2) to the low-MgO (45.5-52.1 wt.% SiO2) basalts (Fig. 4.6a). Other major element oxides show either overall large scatter or display trends that may in fact represent differentiation systematics within each of the MgO groups (Fig. 4.6). Therefore, the geochemical characteristics of the two MgO groups will be described separately, below.

110 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.5: a) Ternary AFM diagram for the Low-Strain Belt samples; boundary between tholeiitic and

calc-alkaline from Irvine and Baragar (1971); b) ternary CaO-MgO-Al2O3 diagram separating tholeiitic basalts from komatiitic basalts (after Viljoen et al., 1982; modified after Jensen, 1976); c) ternary Al2O3-

Fe2O3+TiO2-MgO discrimination diagram for ultramafic and mafic rocks (after Jensen, 1976);

a) High-MgO group

Following a previous classification for basalts from the Warrawoona Group based on TiO2 content (0.8 wt.%; Smithies et al., 2007), the two ocelli-bearing pillow flows are high-Ti basalts (0.98-1.3 wt.%) and show corresponding higher Fe2O3 (Fig. 4.6e) and hence lower Mg# of 69-73, while the two spinifex-textured basalts are low-Ti basalts (0.57-0.76 wt.%) with slightly higher Mg# (77- 79). However, this subdivision does not correlate with any other major element variation, and, as shown in Figure 4.7a, does not correlate with the pattern and level of chondrite-normalized REE concentrations: samples 365-1 and 307-1 show higher HREE levels, lower MnO and CaO, and higher Al2O3, compared to samples 182-1 and 34-1 (Fig. 4.7a; Fig. 4.6a-j). On the other hand, one spinifex sample and one ocelli sample (365-1 and 34-1) show flat chondrite-normalized REE pattern (La/Yb(N) ~0.91-1.50), while the other two samples (182-1 and 307-1) display fractionated L/HREE pattern (La/Yb(N) ~2.81-4.10; Fig. 4.7a). The primitive mantle-normalized multi-element spider-grams reveal further discrepancies in the geochemical signatures that are unrelated to the Ti-groups (Fig. 4.8a). The spinifex-textured sample 182-1 with low HREE concentration and fractionated L/HREE shows the highest levels in Cr and Ni with Ni>Cr. The other samples have lower Ni and Cr concentrations, Cr>Ni fractionation, and display more negative Sr anomalies (Fig. 4.8a). Both of the LREE enriched samples (182-1, 307-2) show

111 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.6: (a-j) Harker variation diagrams for major element oxides for rocks from the Low-Strain Belt; using MgO as an index of differentiation. corresponding enrichment in Th and U, but only the ocelli-bearing 307-2 is characterized by enrichment in other LILE (Ba, Rb, Th, U), while Ba and Rb are generally depleted in the high- MgO basalts.

b) Low-MgO group

All of the low-MgO basalts are high-Ti basalts (0.87-1.43 wt.%) and display a systematic increase in Fe2O3 and TiO2 and a decrease in Al2O3 and Cr with decreasing MgO content and hence decreasing Mg# from 75 to 56 (Fig. 4.6), while other major elements show wide scatter. Similar to

112 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.7: (a-d) Chondrite-normalized (McDonough and Sun, 1995) rare earth element (REE) patterns for rocks from the Low-Strain Belt. the high-MgO basalts, where major element trends are present they do not correlate with REE patterns (Fig. 4.7b). Two samples (30-1 and 40-1) define a slightly LREE depleted (La/Yb(N) ~0.80-0.88) group, while the other samples display slightly enriched, L/HREE fractionated patterns (La/Yb(N) ~1.10-1.74; Fig. 4.7b). The degree of enrichment of the second group lies in- between the flat and the enriched LREE groups of the high-MgO basalts, while HREE concentration levels reach slightly higher values, i.e. ~16 times that of chondrite compared to ~13 times in the high-MgO group. Only the anorthositic-gabbroic rock displays significant L/HREE fractionation with La/Yb(N) ~6.17 and shows an almost parallel REE pattern to the enriched spinifex-textured rock (182-1; Fig. 4.7c). Similar to the high-MgO basalts, the primitive mantle- normalized Th and U concentration levels display a positive correlation with the observed LREE variation in the low-MgO basalts (Fig. 4.8b). Besides the most strongly L/HREE fractionated sample (259-1), which displays depleted Ba and Rb signatures compared to Th, U and REE, most of the low-MgO basalts are enriched in Ba and Rb (Fig. 4.8b). The fractionation of Cr versus Ni shows flat primitive mantle behavior within both the slightly LREE depleted and the most strongly L/HREE fractionated pillow basalts, while those samples with only slight LREE enrichment display Cr>Ni fractionation patterns (Fig. 4.8b).

4.5.2.2 Duffer Formation

The Duffer Formation silicic volcanics plot along a calc-alkaline trend (Fig. 4.5a) mostly due to their high K2O content (~3-7 wt.%), while Na2O is extremely depleted (<0.12 wt.%; Table 4.2;

Fig. 4.6b,d). The SiO2 contents are extremely high, ranging from 74-81 wt.%, the samples show

113 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.8: (a-c) Primitive mantle normalized (McDonough and Sun, 1995) multi-element spider-diagrams for rocks from the Low-Strain Belt.

high TiO2 (1.30-1.76 wt.%; Table 4.2), and Al2O3 contents are slightly lower than those of the high-Ti pillow basalts (Fig. 4.6i). All other major elements are extremely depleted (Fig. 4.6).

An overall increase in REE concentration levels (Fig. 4.7d) and decrease in Fe2O3, MgO and Mg# (~73 to ~57 Mg#) roughly correlates with the stratigraphic younging. The REE patterns of the silicic volcanics are relatively parallel and characterized by slightly depleted HREE (Gd/Yb(N)

114 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

~1.27-2.17), and enriched LREE (La/Sm(N) ~1.79-2.5) with overall high H/LREE fractionation (La/Yb(N) ~2.97-6.69) similar to the most enriched ocelli-bearing pillow basalt (Fig. 4.7). However, in the enriched pillow basalts an upward convex LREE pattern is observed (Fig. 4.7a,b), while the silicic volcanics display a downward convex LREE curvature (Fig. 4.7d). The silicic volcanics overall show very similar primitive mantle-normalized trace element behavior to the enriched pillow basalts, however, they tend towards much higher LILE enrichment, including strong positive K anomalies - compared to the pillow basalts that show K depletion - with increasing REE concentration, and show much lower Cr and Ni contents (Fig. 4.8c). The increasing differentiation within the silicic volcanics further correlates with an increasing negative Ti anomaly and increasing P depletion in the primitive mantle-normalized spider-gram (Fig. 4.8c), in agreement with the Harker variations (Fig. 4.6).

4.6 Petrography and Geochemistry of Mafic Rocks from the Central Fold Belt and South Muccan Shear Zone

4.6.1 Petrography

4.6.1.1 Central Fold Belt

Within the Central Fold Belt, some of the massive mafic rocks (e.g. 358-1; Fig. 4.9a) resemble the medium-grained phaneritic pillow basalts from the Low-Strain Belt (e.g. 30-1) in both texture and inferred primary mineralogy, but do not show evidence for pillows. One sample (364-1) displays a fine- to medium-grained assemblage comprising both preserved magmatic plagioclase with polysynthetic twinning and recrystallized secondary albite, quartz, chlorite and calcite. Within some massive mafic rocks, pseudomorphs after coarse-grained magmatic minerals display a remnant interlocking texture, with interstitial medium-grained opaque phases (359-2, 360-1). The coarse-grained former magmatic minerals are replaced by fine- to medium-grained either chlorite- quartz-albite-rich or chlorite-clay-rich assemblages. One sample (359-1) represents a ca. 20 cm wide mafic dyke that seems to be related to a topographic ridge formed by more massive mafic rock (359-2, 360-1). The dyke contains abundant euhedral phenocrysts (ca. 50-700 µm size) that resemble former olivine crystals (Fig. 4.9b). The phenocrysts often show resorption textures, characterized by embayments that are filled by a former aphanitic groundmass, now finely recrystallized to chlorite and clay minerals. The sampled mafic schist (234-1) from the northern Central Fold Belt comprises abundant epidote in association with chlorite, albite, quartz and opaque phases. The original magmatic texture of the rock is unclear. The gabbroic rocks are characterized by intermediate preserved coarse-grained magmatic textures, often partly preserving plagioclase and clinopyroxene. Greenschist-facies alteration is evident in chlorite and clay minerals replacing clinopyroxene and plagioclase along grain boundaries, rims or fractures. Opaque minerals occur in minor abundance. Anorthosites comprise

115 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______well-preserved plagioclase, displaying polysynthetic twinning, as well as rare simple concentric zoning. Some grains are affected by sericitization and albitization. One anorthosite body shows

Fig. 4.9: Selected petrographic images of representative textures of rocks from the Central Fold Belt and South Muccan Shear Zone; refer to text for description. intensive sericitization of plagioclase and contains abundant fine-to medium grained calcite grains growing between the coarse-grained feldspars (Fig. 4.9c). A relatively large ultra-mafic - mafic body within the northern Central Fold Belt comprises a (lower?) zone of clinopyroxenite (Fig. 4.9d), mantled by anorthosite, anorthositic gabbro and gabbro. The coarse-grained clinopyroxene is partly preserved, showing pleochroism and occasional twinning. The grains are partly replaced by a chlorite-epidote assemblage.

116 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Table 4.4:XRF major element data for mafic rocks from the Central Fold Belt and South Muccan Shear Zone

Central Fold Belt South Muccan Shear Zone

sample 371-1 346-1 315-1 359-2 360-1 359-1 234-1 364-1 97-2 347-1 99-5 358-1 348-1 318-1 320-3 320-2 100-1 312-1 311-1 330-1 330-2 (wt.%) kb kb kb kb kb kb ba ba gb gb gb do do do an an px ba ba hzbgt hzbgt

SiO 2 47.90 47.93 52.42 42.39 42.26 42.39 51.10 49.97 48.00 51.26 49.40 52.41 45.77 50.36 50.94 53.49 52.60 51.13 51.50 38.49 38.12

TiO2 0.48 0.68 0.59 0.86 0.92 0.86 0.91 0.95 1.46 0.99 0.64 1.58 1.04 0.96 0.76 0.76 0.66 1.41 1.02 0.19 0.18

Al2O3 4.39 6.15 6.36 5.94 6.57 7.29 9.67 13.48 12.34 13.78 15.99 13.07 9.09 13.90 24.44 26.08 4.54 14.15 16.20 1.26 0.98

Fe2O3 9.24 10.84 10.47 14.17 12.99 12.81 12.13 11.19 13.87 11.31 8.93 15.04 15.20 11.09 2.65 2.65 10.72 13.69 10.07 14.68 13.54 MnO 0.21 0.19 0.11 0.27 0.17 0.20 0.19 0.14 0.24 0.16 0.12 0.16 0.21 0.13 0.06 0.03 0.29 0.20 0.15 0.20 0.22 MgO 14.08 12.78 11.75 11.51 13.23 11.50 7.58 8.34 5.42 6.20 6.58 4.98 8.11 7.53 2.88 2.99 14.08 6.36 5.41 35.43 36.00 CaO 14.39 8.48 6.51 10.53 9.36 10.15 8.46 5.98 10.09 9.03 8.62 4.37 8.41 7.18 4.58 1.92 15.77 8.32 8.46 0.09 0.60

Na 2O 0.17 0.07 0.06 0.06 0.06 0.07 1.74 2.18 4.24 3.74 3.04 3.60 1.94 3.42 6.07 6.61 0.44 2.97 3.94 0.04 0.04

K2O 0.04 0.01 0.01 0.01 0.01 0.01 0.40 0.46 0.26 0.17 1.11 0.05 0.02 0.33 2.50 2.01 0.11 0.67 1.34 0.01 0.01

P 2O5 0.04 0.06 0.05 0.07 0.09 0.10 0.07 0.09 0.15 0.09 0.06 0.17 0.09 0.11 0.05 0.03 0.05 0.15 0.15 0.03 0.04 LOI 9.08 12.80 11.67 14.19 14.34 14.63 7.74 7.22 3.93 3.28 5.52 4.55 10.10 4.98 5.07 3.43 0.74 0.96 1.78 9.60 10.27 Total 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 100 Mg# 86 82 82 76 51 78 71 75 61 68 74 80 68 73 81 82 84 65 68 91 91 Mg# calculated as Mg/(Mg+Fe)

4.6.1.2 South Muccan Shear Zone

The two (mylonitic) amphibolite schists from the South Muccan Shear Zone comprise metamorphic hornblende porphyroclasts that are partly replaced by syn-kinematic actinolite, within high-strain shear bands of foliation-defining actinolite-chlorite-quartz-albite and elongated schlieren of opaque phases (Fig. 4.9e). The ultra-mafic (mylonitic) schist from the northernmost South Muccan Shear Zone is formed by ~500 µm-size sigma-shaped porphyroclasts that are almost entirely replaced by fine-grained serpentine (± chlorite) and surrounded by thin shear bands of opaque minerals (Cr-spinel?) and minor chlorite (Fig. 4.9f).

4.6.2 Major and trace element geochemistry

XRF major element data is given in Table 4.4 and LA-ICP-MS trace element data is given in Table 4.5. The AFM and ultramafic - mafic discrimination diagrams are shown in Figure 4.10a-c comparing the mafic rocks from the Central Fold Belt and South Muccan Shear Zone to the Low- Strain Belt.

4.6.2.1 Mafic rocks from the Central Fold Belt and South Muccan Shear Zone

With decreasing MgO content the mafic rocks from the Central Fold Belt and South Muccan Shear

Zone show an overall increase in Na2O, K2O, SiO2, Al2O3, TiO2, and a decrease in CaO, MnO (wt.%) and Cr (ppm) as shown in the Harker variation plots (Fig. 4.11a-j). Similar to the mafic rocks from the Low-Strain Belt, the mafic rocks from the Central Fold Belt and South Muccan Shear Zone can be grouped into high- and low-MgO compositions (Fig. 4.11). The high-MgO group includes five of the massive mafic rocks (11.50-14.08 wt.% MgO) and the mafic dyke (11.51 wt.% MgO). The rocks generally show high Mg# of 76-86 with one sample showing an exceptionally low Mg# (51; Table 4.4). The clino-pyroxenite shows similar high MgO and Mg# (14.08 wt.%, 84), and the two ultramafic schists from the South Muccan Shear Zone have

117 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Table 4.5: LA-ICP-MS trace element data for mafic rocks from the Central Fold Belt and South Muccan Shear Zone Central Fold Belt South Muccan Shear Zone

sample 371-1 346-1 315-1 360-1 359-2 359-1 234-1 364-1 347-1 358-1 348-1 318-1 99-5 97-2 320-3 320-2 100-1 312-1 311-1 330-1 330-2

(ppm) mas mas mas mas mas dyke mafs mas mas mas mas mas gb gb an an px amps amps umas umas P 146.87 207.02 206.61 390.49 285.49 368.52 327.92 436.47 331.61 780.10 398.48 434.67 457.72 628.75 116.20 -2.00 221.08 634.58 601.08 107.80 91.24 Sc 35.68 38.98 37.40 22.37 20.93 22.37 32.93 39.97 28.57 15.95 27.51 22.94 30.51 29.42 16.63 18.26 26.43 24.87 21.61 6.54 8.12 Cr 2073.07 885.42 1180.46 1086.45 1220.42 1575.23 335.79 182.04 192.14 29.75 258.81 299.66 973.63 105.20 3.16 1.28 20.32 149.90 267.22 3842.73 3662.90 Ni 199.34 163.50 151.00 409.88 543.24 1034.09 98.93 70.98 96.53 50.90 139.69 147.84 315.86 71.83 28.55 22.06 63.41 94.16 80.50 1741.89 1571.71 Cu 26.89 44.07 37.83 133.38 35.16 101.95 100.55 66.54 205.29 135.71 241.31 242.01 152.39 34.25 56.37 16.56 24.96 32.27 9.13 27.16 17.45 Rb 0.53 0.23 0.17 0.15 0.10 0.11 23.39 12.51 3.71 0.85 0.58 6.05 35.13 2.13 51.45 41.40 37.47 14.14 32.80 0.61 0.38 Sr 145.66 107.02 99.84 87.71 125.80 143.16 236.35 21.80 143.93 102.31 172.26 187.03 72.52 187.56 84.31 128.63 172.38 173.26 295.13 2.17 2.46 Y 9.75 12.56 9.26 17.96 13.98 12.28 19.27 22.31 17.50 28.92 20.30 19.62 16.41 25.52 22.06 21.72 12.97 25.12 24.99 2.59 2.93 Zr 25.15 40.72 36.62 71.98 51.39 74.35 64.30 76.98 62.30 133.97 85.91 76.48 72.98 104.77 85.91 121.39 37.12 113.30 119.21 20.59 15.22 Nb 1.25 1.99 1.78 4.22 3.54 3.71 3.19 3.61 3.64 8.38 4.92 4.37 5.59 5.58 4.14 3.88 2.10 5.82 5.85 0.75 0.88 Ba 36.74 118.04 14.99 34.85 230.64 6.28 867.76 73.34 99.90 48.22 14.55 1511.74 190.42 69.96 212.78 219.41 398.30 437.21 438.63 9.24 18.99 La 2.30 3.10 2.44 6.30 4.61 7.96 4.72 4.47 5.16 8.99 7.98 6.34 6.44 6.22 3.49 2.89 3.25 10.14 9.59 0.99 1.37 Ce 5.98 8.20 7.38 16.14 13.31 18.35 12.45 11.83 12.57 23.83 20.20 17.14 18.33 18.80 8.46 7.38 8.55 23.96 23.76 3.04 3.46 Pr 0.82 1.18 1.00 2.26 1.87 2.35 1.78 1.66 1.82 3.35 2.84 2.28 2.66 2.20 1.06 0.90 1.25 3.10 3.13 0.35 0.42 Nd 3.96 5.63 4.61 10.99 8.58 10.44 8.73 8.26 9.05 16.00 13.07 10.82 12.71 10.36 5.27 4.77 5.97 14.08 14.20 1.58 1.81 Sm 1.22 1.67 1.21 2.76 2.24 2.45 2.56 2.77 2.69 4.34 3.39 2.90 3.20 3.02 1.45 1.51 1.66 3.66 3.53 0.46 0.42 Eu 0.58 0.56 0.41 0.98 0.92 1.04 0.90 0.89 0.97 1.43 1.12 1.03 1.19 0.96 1.00 0.63 0.76 1.14 1.16 0.12 0.15 Gd 1.62 2.17 1.55 3.45 2.69 2.65 3.58 3.51 3.20 5.00 3.81 3.73 3.19 3.69 2.67 2.27 2.23 4.38 4.17 0.44 0.54 Tb 0.28 0.36 0.27 0.53 0.41 0.39 0.56 0.62 0.53 0.81 0.61 0.59 0.50 0.63 0.47 0.34 0.35 0.71 0.69 0.07 0.10 Dy 1.80 2.56 1.85 3.36 2.76 2.40 3.54 4.03 3.31 5.36 3.90 3.66 3.19 4.66 3.07 3.06 2.31 4.61 4.53 0.47 0.56 Ho 0.38 0.46 0.33 0.65 0.49 0.51 0.68 0.84 0.65 1.08 0.77 0.68 0.61 0.95 0.65 0.71 0.45 0.91 0.87 0.09 0.11 Er 1.12 1.39 1.08 1.89 1.59 1.36 2.03 2.43 1.92 3.31 2.24 1.93 1.69 2.95 1.85 1.76 1.37 2.80 2.88 0.26 0.37 Tm 0.15 0.20 0.17 0.27 0.22 0.19 0.32 0.34 0.26 0.47 0.32 0.29 0.23 0.44 0.24 0.24 0.19 0.39 0.40 0.04 0.05 Yb 1.00 1.36 1.06 1.66 1.41 1.10 1.83 2.54 1.94 3.10 2.09 1.85 1.52 2.97 1.60 2.06 1.26 2.73 2.78 0.21 0.28 Lu 0.13 0.18 0.15 0.27 0.19 0.18 0.27 0.38 0.24 0.43 0.29 0.24 0.22 0.41 0.29 0.27 0.17 0.40 0.40 0.04 0.04 Hf 0.73 1.20 0.96 1.92 1.43 2.05 1.78 2.12 1.63 3.34 2.34 2.10 1.87 2.80 2.14 2.50 1.01 2.93 3.07 0.48 0.35 Ta 0.07 0.11 0.11 0.28 0.20 0.28 0.20 0.24 0.18 0.51 0.28 0.27 0.35 0.34 0.33 0.20 0.12 0.38 0.37 0.05 0.05 Th 0.26 0.40 0.35 0.79 0.54 1.65 0.61 0.71 0.74 1.51 1.61 0.91 1.32 1.50 1.53 1.13 0.36 2.06 2.19 0.16 0.17 U 0.06 0.09 0.09 0.20 0.14 0.35 0.14 0.18 0.21 0.35 0.34 0.23 0.31 0.22 0.25 0.26 0.12 0.39 0.36 0.03 0.03 Mg# calculated as Mg/(Mg+Fe; amps, amphibolite schist; an, anorthosite; dyke, mafic dyke; gb, gabbro; mafs, mafic schist; mas, massive mafic rock; px, clino-pyroxenite; umas, ultramafic schist extremely high MgO of 35.43 and 36.00 wt.%, respectively, at very high Mg# (91). The other massive mafic rocks, the two gabbros, and the amphibolite schists belong to the low-MgO group (4.98-8.11 wt.%) and have lower Mg# (61-80). Both the anorthosite samples display the lowest MgO content (2.88 and 2.99 wt.%), but have high Mg# (~81). The low-MgO mafic rocks belong to the high-Ti group (0.91-1.58 wt.% TiO2). Within the high-MgO mafic rocks three massive mafic rocks belong to the low-Ti group (0.48-0.68 wt.%) and two massive mafic rocks (0.86-0.92 wt.%) and the mafic dyke (0.86 wt.%) belong to the high-Ti group. Similar to the high-MgO group from the Low-Strain Belt, the high-MgO mafic rocks from the Central Fold Belt can be subdivided into a group with relatively flat chondrite-normalized REE patterns (La/Yb(N) ~1.55-1.57) comprising the three low-Ti group mafic rocks, and a group with higher L/HREE fractionation patterns (La/Yb(N) ~2.22-2.58; Fig. 4.12a). The more LREE enriched group also displays overall higher REE concentration levels (Fig. 4.12a). The high-MgO olivine phenocrystic mafic dyke displays the highest L/HREE fractionation (La/Yb(N) ~4.9), slightly higher than the adjacent LREE enriched massive mafic rock (359-2). The dyke shows remarkably similar REE concentration levels and patterns to the spinifex-textured rock 182-1 from within the Low-Strain Belt. The ultramafic schists from the South Muccan Shear Zone show L/HREE fractionation (La/Yb(N) ~3.15-3.35) in-between the more LREE enriched massive rocks and the mafic dyke, but at much lower HREE concentration levels (~1.5-2 times chondrite; Fig. 4.12a). The low-MgO mafic rocks, including the amphibolite schists from the South Muccan Shear Zone, are all enriched in incompatible LREE with L/HREE ratio variations (La/Yb(N) ~1.2- 2.6; Fig. 4.12b) that overlap with those from the relatively flat and the enriched high-MgO REE groups.

118 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.10: a) Ternary AFM diagram for the Central Fold Belt and South Muccan Shear Zone samples; boundary between tholeiitic and calc-alkaline from Irvine and Baragar (1971); b) ternary CaO-MgO-

Al2O3 diagram separating tholeiitic basalts from komatiitic basalts (after Viljoen et al., 1982; modified after Jensen, 1976); c) ternary Al2O3- Fe2O3+TiO2-MgO discrimination diagram for ultramafic and mafic rocks (after Jensen, 1976).

The REE patterns of the intrusive rocks are shown in Figure 4.12c and display relatively flat REE patterns for the anorthosites (La/Yb(N) ~0.95-1.48), the clino-pyroxenite (La/Yb(N) ~1.74) and one of the gabbros (La/Yb(N) ~1.42). The gabbro, however, plots at overall elevated REE concentration levels of ~17 times that of chondrite compared to the other intrusions (7-11 times chondrite). One gabbro shows more highly fractionated L/HREE with La/Yb(N) ~2.88 (Fig. 4.12c), which is in the range of the enriched high-MgO massive mafic rocks. Overall, the different REE patterns of both the high- and low-MgO mafic rocks from the Central Fold Belt resemble the different REE signatures observed in the Low-Strain Belt. Despite the varying L/HREE fractionation, the primitive mantle-normalized patterns show relatively consistent signatures of the high-MgO mafic rocks from the Central Fold Belt (Fig. 4.13a). Compared to the high-MgO group from the Low-Strain Belt, Sr anomalies are absent, or only very slightly positive or negative, depending on the enrichment of the surrounding LREE (Fig. 4.13a). All samples show depletion in Ba, K, and Rb. The major differences within the high- MgO group are the Ni-Cr fractionation systematics (Fig. 4.13a). The high-MgO, low-Ti samples with the relatively flat REE pattern, show the highest Cr>Ni fractionation, while the two high-Ti group massive mafic rocks show little Cr>Ni fractionation. In contrast, the mafic dyke (high-Ti group) shows a Ni>Cr primitive mantle-normalized fractionation pattern (Fig. 4.13a), similar to the spinifex-textured mafic rock 182-1 from the Low-Strain Belt.

119 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.11: (a-j) Harker variation diagrams for major element oxides of the Central Fold Belt and South Muccan Shear Zone mafic components in the Low-Strain Belt (closed circles) and DF rhyolites (open circles) using MgO as an index of differentiation.

The low-MgO mafic rocks from the Central Fold Belt display either LILE enriched or depleted (Ba, Rb, K) signatures (Fig. 4.13b) The Cr-Ni fractionation patterns range from Cr>Ni to Ni>Cr. Where Ni is more abundant than Cr, the overall Ni and Cr concentrations are low (Fig. 4.13b,c). The highest Ni and Cr concentrations with Cr>Ni fractionation are found in the very high-MgO ultramafic schists from the South Muccan Shear Zone, while the lowest Ni-Cr concentrations are observed in the anorthosite with Ni>Cr (Fig. 4.13b,c).

120 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.12: (a-c) Chondrite-normalized (McDonough and Sun, 1995) rare earth element (REE) patterns for rocks from the Central Fold belt and South Muccan Shear Zone.

4.7 Discussion

The presented field relationships and petrographic and geochemical data demonstrate that the Doolena Gap greenstone belt transect comprises mafic and silicic rocks characterized by a variety of geochemical signatures and emplacement styles, including volcanic flows, intrusions and massive rocks of undefined either volcanic or intrusive origin. Below, I will use the presented data to discuss petrogenetic processes and relationships between the rock types to shed light on the volcanologic-magmatic growth of the early Archaean upper crust in the study area. One of the main aims is to elucidate whether the silicic rocks of the Duffer Formation represent differentiation products of the underlying mafic rocks or if they represent surficial expressions of mid crustal TTG magmatism. Because the stratigraphic affinity between the Low-Strain Belt and the vertically juxtaposed Central Fold Belt and South Muccan Shear Zone (Chapter 2) remains unclear, I will first discuss the petrogenetic affinities within the continuous stratigraphic mafic to

121 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.13: (a-c) Primitive mantle normalized (McDonough and Sun, 1995) multi-element spider-diagrams for rocks from the Central Fold Belt and South Muccan Shear Zone. felsic succession of the Low-Strain Belt. Then, I will critically analyze the stratigraphic position of the Central Fold Belt and discuss its possible petrogenetic relationship with the Low-Strain Belt. Finally, I will test two end-member models, i.e. different mantle sources versus intra-crustal differentiation processes, for the derivation of the observed compositional mafic rock types. In order to test which of the presented geochemical signatures are reliable and represent primary compositions form which petrogenetic processes can be inferred, I will first integrate the data in regards to possible alteration and metasomatic modification.

122 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

4.7.1 Effects of alteration and metamorphism/metasomatism

Early Archaean greenstone rocks are affected by at least sub-greenschist- to greenschist-facies metamorphism, and more or less pervasive seafloor hydrothermal alteration (e.g. Appel et al., 2001; Terabayashi et al., 2003; Van Kranendonk and Pirajno, 2004; Thébaud and Rey, 2013). Moreover, as outlined in Chapter 2, the western Doolena Gap greenstone belt experienced five deformational events (D1-D5), each associated with localized fluid-induced carbonate alteration. The effects of seafloor alteration and polyphase metasomatism need to be accounted for in constraining petrogenetic processes and relationships based on primary geochemical signatures. As outlined in Chapter 2, the pillow flows of the Mount Ada Basalt were erupted in a deep marine environment. Therefore it is likely that even the well-preserved mafic flows of the Low- Strain Belt experienced some degree of interaction with Archaean ocean water, and possibly locally with element-enriched hydrothermal vents, in association with chert formation, as reported from other greenstone successions (e.g. de Wit and Furnes, 2016; Van Kranendonk and Pirajno, 2004). Based on the composition of various generations of fluid inclusions it has been proposed that ancient seawater contributed to an upper crustal-scale hydrothermal plumbing system, spanning most of the Pilbara Supergroup depositional history (Thébaud et al., 2006; Thébaud and Rey, 2012). Furthermore, it is suggested that hydrothermal alteration episodes associated with chert formation (hydrothermal vents) affected basalts of the Warrawoona Group repeatedly (Van Kranendonk and Pirajno, 2004). In addition to possible alteration through hydrothermal fluids and interaction with seawater, the dominantly greenschist-facies Central Fold Belt, is intensively affected by localized fluid-induced carbonate alteration zones and associated chlorite, calcite and quartz vein networks, which were active during up to five deformational episodes (Chapter 2). In the vicinity of these alteration-deformation zones, both the addition of species by element-enriched fluids, as well as elemental mobilization and leaching out of the host rock might be expected. Fluids released from crystallizing granitic magma in the dome or from metamorphic de-hydration reactions that took place in the more deeply buried and/or sagducting pillow basalt pile likely infiltrated the uppermost crust from below. The presence of secondary quartz, chlorite, and calcite veins, quartz and epidote aggregates and euhedral pyrite overgrowing highly altered chlorite-rich rock matrix, as well as syn-deformational Ti-phases (e.g. rutile; Chapter 2) suggest open-system hydrothermal growth of these minerals (i.e. chloritization, epidotization, silicification). None of the studied rocks is fully equilibrated, and partial alteration seems to be controlled by the local effect of metasomatic fluid influx, often preserving magmatic textures and mineralogy in other domains. Within the South Muccan Shear Zone some of the mafic schists show evidence for early amphibolite-facies metamorphism. The spatially restricted amphibolite-facies assemblages are interpreted as the result of contact-style metamorphism as a response to magmatism within the dome (Chapter 2). These rocks therefore likely experienced burial and seafloor hydrothermal greenschist-facies alteration, followed by prograde amphibolite-facies contact metamorphism. Subsequent retrogression is evident in the partial replacement of the peak metamorphic assemblage by a greenschist-facies assemblage of chlorite-actinolite-quartz-calcite. The Duffer

123 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Formation in the Low-Strain Belt shows replacement of primary ilmenite by rutile and silicification (Fig. 4.3b). In summary, field and petrographic evidence suggests that polyphase alteration by different mechanisms affected the Doolena Gap greenstone belt. This requires further evaluation of elemental mobility and redistribution. Overall, however, it should be noted that in the present study only freshest samples from within relatively fresh and un-deformed lozenges were used for geochemical investigation. Therefore, the following assessment of alteration effects is restricted to determine primary signatures of least altered samples and does not present an overall assessment for the entire transect, which shows at least locally higher alteration intensity zones (Chapter 2; Burke-Shyne et al., 2015, abstr.).

4.7.1.1 Assessment of least altered geochemical signatures

As shown in the Harker diagrams (Fig. 4.6; Fig. 4.11) the major element oxides display systematic variations with decreasing MgO content. However, most trends are characterized by some degree of scatter (e.g. CaO, Na2O and K2O), possibly related to alteration related modification. A commonly used assessment of the intensity of alteration is based on the loss on ignition (LOI). There is no systematic correlation between LOI (Table 4.2 and 4.4) and major oxide variations in the studied samples. This is demonstrated in comparing the high-MgO and the low-MgO group rocks from within the Central Fold Belt. The high-MgO rocks display higher LOI (9.08-14.34) than the low-MgO group rocks (LOI ~3.28-10.10), and show lower Na2O and K2O content.

However, lower Na2O and K2O are expected features of more primitive mafic rocks, and the Na2O and K2O trends display some systematic behavior with other major element oxide variations (Fig. 4.11). More importantly, within each of the MgO-groups the LOI does not show systematic correlation to any oxide. The LOI, therefore, is not useful to assess elemental mobility in the studied rocks. This is supported by the fact that there is no correlation between the LOI and the textural preservation of the rocks. Considering the commonly accepted immobile behaviour of high-field strength elements (HFSE), zirconium content is used as a proxy to test the reliability of primary geochemical signatures (Polat et al., 2002). Figure 4.14 shows relatively systematic correlations of selected elements with Zr contents of both high- and low-MgO mafic rocks from the Low-Strain Belt,

Central Fold Belt and South Muccan Shear Zone. Only CaO and K2O, and Al2O3 of a few samples, display a poor correlation. Critical major element oxides (e.g. MgO, TiO2) and trace elements (HFSE, REE) do not seem to be disturbed in correlation to Zr, and hence are not significantly affected by mobility. However, the possibility of HFSE mobility needs to be considered as it has been shown that HFSE can be mobilized under specific pressure-temperature-fluid composition (P-T-X) conditions (e.g. Woodhead et al., 2001; Jiang et al., 2005; Torres-Alvorado et al., 2007). The danger in relying on correlations between HFSE and other elements to infer immobile behavior lies in the fact that the mobility of HFSE has shown to be coupled with the mobility of other elements, such

124 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.14: Zr (ppm) versus selected trace elements, trace element ratios, and major element oxides; refer to text for discussion. as REE. Alteration fluids that comprise mobilized HFSE often retain the elemental ratios of the leached host rock (review in Jiang et al., 2005). Such a scenario may have some meaningful geological background in the present case. As shown in Figure 4.14 HFSE and REE show an increase with increasing Zr, but retain the same elemental ratios. In a sagduction environment envisaged by the sinking of vertically stacked pillow basalts (Chapter 2), similar original compositions may characterize both surface rocks and the ones that are being buried (i.e. sagducted) and de-hydrated to release fluids. In fact, secondary Ti-phases (titanite, rutile) are observed replacing primary Ti-phases (magnetite, ilmenite) or overgrowing chlorite-rich matrix.

125 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

These observations may be interpreted as indicative of at least local HFSE mobilization (Jiang et al., 2005). Some of the secondary textures of the studied basalts indicate open-system non- equilibrium growth. The Ti-phases may have been precipitated from Ti-rich metamorphic fluids, released by de-hydration reactions. Known controls on HFSE mobility include specific P-T conditions, pH of solution, and fluid chemistry (review in Jiang et al., 2005; Linnen et al., 2014). The mobility of Zr is enhanced in highly saline fluids during metamorphic conditions. Zirconium forms stable F-OH complexes (e.g. ZrF(OH)3, ZrF2(OH)2) in solutions up to 400 °C, 0.7 kbar. Other evidence indicates HFSE and REE mobilization as phosphate complexes (Jiang et al., 2005; Linnen et al., 2014). In the studied rocks, increasing LREE show a good correlation with increasing Zr and with P (Fig. 4.14). Metabasalts from the Dresser Formation and Apex Basalt of the Warrawoona Group that occur in association with (hydrothermal) chert display both LREE depletion and/or enrichment as the result of LREE leaching and mobilization (Van Kranendonk and Pirajno, 2004). LREE enrichment has been observed in the basalts overlying the chert, while LREE depletion has been observed in the basalts below the chert, interpreted as due to redistribution in acidic alteration solutions (e.g. Michard, 1989; Van Kranendonk and Pirajno, 2004). It should be noted that the latter study (Van Kranendonk and Pirajno, 2004) does not consider different primary REE types for the basalts. In fact, the presence of chert indicates a hiatus in volcanism, therefore making it likely that new eruptions may show a different composition with different REE signatures. Nevertheless, numerous other studies have demonstrated the redistribution of LREE in pillow basalts. Rims of zoned pillow basalts have been shown to being depleted in LREE (e.g. Polat et al., 2007) or enriched in LREE (e.g. Frey et al., 1974; Ludden and Thompson, 1978) relative to the interiors of the pillows. In the present study, the slightly depleted LREE signatures of the two texturally altered basalts from the Low-Strain Belt are therefore not considered reliable, as these samples may indeed represent outer zones of pillows, affected by more intensive alteration. However, the strong enrichment in LREE of other basalts characterized by relatively well- preserved textures and overall high REE concentration levels is considered too large to be the result of LREE redistribution, as REE and Zr mobility is not easily achieved in significant amounts. Overall, I consider the mobilization of LREE as a possible mechanism for depletion and/or enrichment of some of the texturally altered basalts to some degree, especially in the rocks with overall low REE concentration levels. However, the extent of LREE enrichment as observed in the more highly enriched basalts is considered too high to be caused by alteration. This is supported by the fact that the observed different REE patterns are relatively parallel even from very different structural domains (e.g. South Muccan Shear Zone versus Low-Strain Belt). Furthermore, the petrographic observations demonstrate that the REE enriched basalts do not show higher textural alteration or deformation than the flat REE types that occur in close spatial association. Therefore, it is proposed that the observed REE groups represent primary magmatic types. Finally, if HFSE mobility indeed occurred, critical elemental ratios, as shown in Fig. 4.14, remain the same and may be used as representing primary signatures.

126 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.15: Discrimination diagram based on immobile elements; after Winchester and Floyd (1977).

4.7.2 Petrogenetic relationship between the Duffer Formation and the Mount Ada Basalt in the Low-Strain Belt

The presence of bimodal mafic-felsic volcanic suites is a typical feature of Archaean greenstone belts worldwide (review in Kamber, 2015). Two types of felsic volcanic rocks occur: i) a crustal- derived type that form the surficial expressions of TTG magmatism, or reworking of older TTG (e.g. Bickle et al., 1983; DiMarco and Lowe, 1989; Kröner et al., 2013; Kamber, 2015), and ii) a type that is petrogenetically related to underlying mantle-derived basalts, as part of ultramafic- mafic-felsic differentiation series (e.g. Smithies et al., 2007). As shown in Figure 4.8c the trace element signature of the Duffer Formation silicic volcanic flows do not resemble those of TTGs from the study area (Chapter 3), or other typical Archaean TTG (not shown; Moyen and Martin, 2012). The major difference is the much lower L/HREE fractionation compared to TTG (Fig. 4.8c). Furthermore, Mg#’s in the Duffer Formation rocks are significantly higher (57-73) compared to Archaean TTGs and granites (8-19; Chapter 3; average

Archaean TTG ~43; Moyen and Martin, 2012). Similarly, TiO2 contents are much higher in the

Duffer Formation rocks with TiO2 ~1.3-1.76 wt.%, compared to typical Archaean TTG (Moyen and Martin, 2012). The stratigraphically lowermost Duffer Formation rocks comprise pseudomorphs after inferred former clino-pyroxene, further arguing against a petrogenetic relationship to the TTG rocks, in which clino-pyroxene is not observed. Instead, the REE patterns and trace element signatures, TiO2 content and Mg# resemble the LREE enriched pillow flows and massive mafic rocks of the uppermost Mount Ada Basalt some of which (e.g. 307-2) are interlayered with the lowermost Duffer Formation rocks (Fig. 4.8c). The stratigraphic younging within the Duffer Formation samples roughly correlates with increasing overall REE concentration levels. This supports the idea that the Duffer Formation silicic volcanics represent the final products of differentiation from the underlying tholeiitic rocks, with successive outpourings increasingly enriched in overall REE concentrations (Fig. 4.7d).

127 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

The intensive silicification of the Duffer Formation flows likely resulted in post-magmatic SiO2 enrichment, preventing rock classification based on major element oxides. Using the classification based on immobile elements (Winchester and Floyd, 1977) demonstrates that the Duffer Formation rocks were originally andesites (Fig. 4.15). This is supported by the inferred presence of clinopyroxene laths. This would imply a compositional transition zone from mafic rocks of the uppermost Mount Ada Basalt towards increasingly evolved intermediate compositions within the Duffer Formation. Such a transition zone has been suggested for a mafic to felsic suite in the lowermost Warrawoona Group (Coonterunah Subgroup), based on field and geochemical evidence (Smithies et al., 2007). I conclude that the Duffer Formation silicic volcanic flows were not derived from TTG. The geochemical trace element signatures and field relationships suggest an affinity to the underlying mafic rocks of the Mount Ada Basalt.

4.7.3 Petrogenetic processes and correlation between mafic rocks from the Low-Strain Belt and mafic rocks from the Central Fold Belt and South Muccan Shear Zone

A critical question in reconstructing the upper crustal growth in the Doolena Gap greenstone belt is whether the mafic rocks from the Central Fold Belt and South Muccan Shear Zone are part of the Mount Ada Basalt. The rocks from the Central Fold Belt are all affected by D2 deformation, i.e. they are either directly affected by S2 foliation, or form elongated lozenges that are aligned in the S2 foliation. Therefore the Central Fold Belt must have been emplaced prior to 3427 Ma (D2; Strelley Pool Formation; Chapter 2). Considering the abundance of intrusive rocks within the northern Central Fold Belt, compared to the absence of crosscutting intrusive rocks within the Mount Ada Basalt and Duffer Formation (despite very rare doleritic dykes in the Duffer Formation, Chapter 2), it is unlikely that the intrusive rocks of the Central Fold Belt belong to the younger Apex Basalt (Fig. 4.1c). The relative timing suggests that the entire Central Fold Belt in fact represents the lower Mount Ada Basalt, or is older. Alternatively, the pillow flows in the Central Fold Belt could be part of an older basaltic succession (e.g. part of the Talga Talga or Coonterunah Subgroup; Fig. 4.1c), while only the intrusive rocks belong to the Mount Ada Basalt. Nevertheless, the intrusive rocks must have been emplaced into the shallow upper crust during or subsequent to massive pillow basalt eruptions. Therefore, the possibility of a petrogenetic relationship between the intrusive rocks from the Central Fold Belt and the pillow basalts from the Low-Strain Belt is considered for the purpose of discussion. Most of the studied samples come from an area in the NW Central Fold Belt that can be considered as a relatively undisturbed succession. Although cm- to m-scale F2 tight to isoclinal folds are observed, the area overall forms a limb of a map-scale F2 fold, that internally is not affected by major map-scale folding, or carbonate shear zone related displacements (Fig. 4.2). Furthermore, at the southern limit of this block, an aphanitic olivine phenocrystic dyke (ca. 20 cm width) is spatially associated with a medium-grained, massive lozenge, and can be traced for a few metres southward, until it is cut by the EW-trending S2 foliation. Considering the dyke moved upwards, this indicates an overall southward stratigraphic younging of this block.

128 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.16: Pseudo-stratigraphic map section through the greenstone belt showing selected geochemical proxies for proposed volcanic cycles; refer to text.

Figure 4.16 shows a pseudo-stratigraphic diagram for selected geochemical proxies through the Doolena Gap greenstone belt transect. The pseudo-stratigraphy does not account for major

129 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______stratigraphic repetition due to F2 folding, however, as discussed above a relatively undisturbed block within the Central Fold Belt is considered as representing an overall stratigraphic upward younging sequence. Figure 4.16 indicates distinct geochemical trends that may suggest a systematic magma evolution showing similar trends for two (fault-separated) cycles within the Low-Strain Belt, including the silicic volcanic flows of the Duffer Formation, and a cycle that is recognized in the relatively undisturbed block within the Central Fold Belt. Particularly in focusing on the high-MgO mafic rocks, the three cycles all display a stratigraphic upward increase in L/HREE fractionation and TiO2 content, while Cr/Ni, Sc/Zr and Mg# decrease (Fig. 4.16). This indicates a cyclic nature of geochemical signatures. Despite the unknown stratigraphic affinity of the Central Fold Belt this implies a cyclic and systematic infra-crustal magma evolution (a) or a systematic emplacement of magma from different sources (b) during bulk upper crustal magmatic construction of the Doolena Gap greenstones. To investigate this further, two end-member models for the petrogenesis of the compositional types that form the bottom and top of the proposed cycles are discussed below: Model I – compositional end-members were generated from different mantle sources, but emplaced in a systematic fashion Model II – the intra-formational cycles represent the result of infra-crustal differentiation processes.

4.7.3.1 Model I: distinct mantle sources

In previous studies of mafic rocks from the Warrawoona Group the presence of interlayered high- and low-Ti group basalts has been interpreted as the contribution of distinct mantle sources from different depths (e.g. Smithies et al., 2007) based on the fact that increasing TiO2 content correlates with increasing depth of the mantle source (e.g. MacGregor, 1969). Such interpretation is often applied to Phanerozoic large igneous provinces (e.g. Xiao et al., 2004). However, not only the depth of the mantle source, but also the degree of partial melting controls the TiO2 content of the melts produced (e.g. Hirschmann et al., 1999). Furthermore, contamination with TiO2-enriched material as argued for some modern-day ocean island basalts can lead to elevated TiO2 contents (e.g. Prytulak and Elliott, 2007). Finally, a combination of the above is possible, e.g. it has been shown that the geographical distribution of low-Ti compared to high-Ti basalts can dependent on the distance in regards to a proposed plume core/head position, i.e. proximal low-Ti melts form through high degree partial melting directly above the plume, while high-Ti melts form through lower degrees of partial melting distal to the plume core and are likely affected by higher degrees of wall-rock contamination (e.g. Maury et al., 2003). In the studied local transect it is impossible to make such distinction. However, using the previously proposed TiO2-based grouping (Smithies et al., 2007) the high- and low-Ti groups show some distinct characteristics. Figure 4.17a shows that the low- and high-Ti groups (shown for high-MgO rocks) plot within distinct fields, with the high-Ti group at low Sc/Zr, but higher L/HREE, and the low-Ti group at higher Sc/Zr and lower L/HREE with the latter having values close to chondritic. Combining this distinction with the

130 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Fig. 4.17: a) La/Yb(N) versus Sc/Zr; DF, Duffer Formation b) Sc/Zr versus Al2O3/TiO2; c) TiO2 versus Al2O3; d) Gd/Yb(N) versus Al2O3/TiO2 with fields for Barberton-type and Munro-type komatiites after Arndt (2004); e) Nb/Th(PM) versus Th/Yb showing the trend for crustal assimilation after Lesher et al. (2001) and values for TTG from Chapter 3; f) Sm/Nd versus Nb/U.

Sc/Zr versus Al2O3/TiO2 plot of Figure 4.17b one may propose that the high-Ti group indeed reflects a distinct mantle source. This is indicated by the garnet removal trend (i.e. retention of Sc; e.g. Nesbitt et al., 1979; Fig. 4.17b), which, in agreement with the high TiO2, would imply partial melting of a deep mantle source (>8 GPa) with majoritic garnet in the solid residue (Agee, 2002; Arndt and Lesher, 2004). This is further supported by Figures 4.17c and d, in which low

131 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Al2O3/TiO2 and high Gd/Yb indicate that garnet retained in the solid mantle residue (e.g. Sun and Nesbitt, 1978; Nesbitt et al., 1982; Arndt and Lesher, 2004). The low-Ti group, on the other hand shows somewhat lower HREE depletion (Gd/Yb(N); Fig. 4.17d), rather flat REE and no significant Sc retention (Fig. 4.17a). The low-Ti group melts could have derived from shallower depth, where garnet melts near the peridotite solidus and is removed from the residue (e.g. Arndt, 2003; Fig. 4.17c,d). In combination with the observed cyclic emplacement, the above signatures may indicate that melt was derived from two different source depths in successive pulses. The ascent of a deep sourced mantle plume would cause partial melting within shallower upper mantle regions above the rising plume head to derive the low-Ti melts, followed by eruption of the high- Ti melts from the deep-seated plume sources. However, the above discrimination diagrams and elemental ratios are generally only applied to the most primitive mantle melts, i.e. komatiites (>18 wt.% MgO; e.g. Arndt and Leher, 2004) that can more directly indicate source signature imprints. Supported by the field relationships, even the high-MgO mafic rocks in the Doolena Gap greenstone Belt are unlikely to represent primitive mantle melts, but rather differentiation products of more primitive compositions. Therefore, the geochemical effect of fractional crystallization will be discussed in Model II.

4.7.3.2 Model II: significance of infra-crustal processes – fractional crystallization and crustal assimilation

The various textural and mineralogical types of mafic rocks within the Doolena Gap greenstone belt in association with intrusive cumulate rocks, such as anorthosite and layered clino-pyroxenite- anorthosite-gabbro bodies suggests that significant fractional crystallization occurred. As mentioned above, even the high-MgO mafic rocks likely already form fractionation products of more primitive, possibly komatiitic melts. Crystal fractionation affects the elemental ratios used above to argue for possibly distinct mantle sources. It has been shown that upon plagioclase fractionation Al2O3/TiO2 ratios drastically drop (e.g. Arndt and Nesbitt, 1982) and may result in a “Al-depleted/Barberton-type” ratio, although the parental melt had much higher ratios (Fig. 4.17c).

Chondritic parental melts could therefore evolve into low-Al2O3/TiO2 melts if plagioclase fractionates. In the Doolena Gap greenstone belt significant fractional crystallization of plagioclase is evident in numerous anorthosite bodies. Furthermore, Clinopyroxene fractionation can be inferred. Sc is compatible in clinopyroxene (e.g. Cattell and Taylor, 1990), resulting in lower Sc/Zr with fractionation of clinopyroxene. Clino-pyroxenite has been observed in the study area. The combined fractionation of plagioclase (anorthosite) and clinopyroxene (pyroxenite) therefore results in the same trend that is expected for garnet removal in a mantle source rock (Fig.

4.17b). Finally, increasing Al2O3, TiO2, P2O5 and Cr-Ni fractionation with decreasing MgO suggest olivine fractionation. Rollinson (1999) showed that LREE enrichment in high-Ti basalts could be produced through olivine and clinopyroxene fractionation of flat REE high-Ti komatiitic basalts. Based on their high Cr and Ni contents and the secondary assemblage of almost pure serpentine (and chromite), the protolith of the ultramafic schists from the South Muccan Shear

132 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

Zone (330-1/2) are here interpreted as an olivine-chromite rock, most likely representing a cumulate. In contrast to the mantle source end-member model (Model I), I propose that the derivation of different mafic rock compositions was controlled by the fractionation of plagioclase, olivine and clinopyroxene. Integrating the geochemical trends of the proposed eruption/emplacement cycles (Fig. 4.16), this model becomes even more viable: more primitive flat REE low-Ti melts erupted (were emplaced) first, followed by fractional crystallization products with lower Al2O3/TiO2 (plagioclase!), lower Sc/Zr and LREE enrichment

(clinopyroxene!), and lower Cr/Ni, lower Mg# and higher TiO2 (olivine!). As shown in Figures 4.17e and f, the Doolena Gap greenstones follow a trend towards the TTG (Chapter 2), in agreement with both the crustal contamination trend of Th enrichment (Lesher et al., 2001; Fig. 4.17e) and U enrichment relative to Nb (Shimizu et al., 2005; Fig. 4.17f). As demonstrated in Chapter 3, a thick crust comprising TTG and granite (ca. 3500-3590 Ma Doolena Suite) must have formed a basement to the Doolena Gap greenstones. It is likely that the mantle- derived melts were contaminated with TTG components during their crustal ascent. Varying degrees of crustal contamination could have caused additional enrichment in incompatible elements (LREE, HFSE). In Chapter 2, I proposed that the Doolena Gap pillow basalts erupted through extensional faults and fissures, likely tapped from intra-crustal magma chambers. In view of the presented data I here propose following volcanological model for the volcanic cycles: shallow crustal tholeiitic chambers were fed from mantle-derived primitive (komatiitic?) magma that differentiated through fractional crystallization and varying degrees of crustal assimilation. The proposed extensional regime (Chapter 2) and the cooling of the pillow basalt pile in which the mantle-magma intruded led to brittle fractures and normal fault systems. Surface water was allowed to penetrate into the developing fractures and caused phreatomagmatic sub-marine volcanism, followed by eruptions of the more primitive, high-MgO, low-Ti, near-chondritic magma, in turn followed by massive outpourings of fractionated enriched and possibly crustal-contaminated surface pillow flows. The more evolved late-stage outpourings (including the final Duffer Formation volcanic rocks) are expected to have higher viscosities and lower temperatures making them to erupt last, supporting the changing eruptive environment observed in the Duffer Formation. The final more evolved and more highly viscose volcanic rocks of the Duffer Formation possibly erupted by forming domes, and hence providing topography for the observed basin that was filled by authigenic volcaniclastics and affected by turbiditic influx from the more felsic volcanic topography (Chapter 2). The final-stage of fractional crystallization differentiation (Duffer Formation) may have been the result of the temporal cessation of further mafic magma infiltration from the mantle. The paucity of komatiitic melts remains unclear. Komatiites may have been erupted, but are not readily recognizable due to intensive alteration and deformation, or are not preserved due to tectonic removal during sagduction (Chapter 2). Nevertheless, significant amounts of komatiitic melts are suggested to have differentiated through fractional crystallization during crustal ascent into the flat REE low-Ti high-MgO melts, which in turn further evolved through fractional crystallization.

133 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

4.8 Conclusions

In this Chapter I presented field observations, and new petrological and geochemical data of greenstone rocks from a sampling transect through the studied western Doolena Gap greenstone belt. The main aim of this Chapter was to test if the silicic volcanic rocks of the Duffer Formation in the relatively undeformed Low-Strain Belt represent surficial expressions of contemporaneous TTG magmatism (or reworking of TTG rocks), or if they form part of a tholeiitic differentiation series and hence show a petrogenetic relationship with mafic rocks from the underlying Mount Ada Basalt. Furthermore, I addressed the petrogenesis of mafic rocks and the petrogenetic relationship between various mafic components within the greenstone belt, including pillow basalt flows, massive mafic rocks of undefined emplacement style, and mafic – ultra-mafic intrusive rocks, both within the well-preserved Low-Strain Belt and in correlation to the highly deformed Central Fold Belt and South Muccan Shear Zone. In order to identify primary geochemical signatures, I assessed the extent and intensity of alteration effects. I put forth following conclusions:

1) Primary geochemical signatures of critical major (MgO, TiO2) and trace elements (HFSE, REE) are demonstrated based on the recognition of geochemical trends and systematic variations in relation to immobile HFSE (i.e. Zr), and primary magmatic signatures are preserved in the relationship between select elements. This is supported by the preservation of magmatic textures and part of the primary mineralogy. Intensive and multiphase alteration is observed in the study area, however, for the purpose of this study only freshest samples from most intact lozenges were sampled. Some major elements that

are considered mobile, do show significant scatter (e.g. Na2O, K2O, CaO), and were not used to imply petrogenetic processes. 2) The silicic volcanic rocks of the Duffer Formation show no compositional affinities to TTG, indicated by much lower L/HREE fractionation and significantly higher Mg# and

TiO2 contents compared to TTG, and the presence of pseudomorphs after clinopyroxene. Instead, the trace element signatures of the silicic volcanic rocks resemble those of LREE enriched pillow basalt flows and massive mafic rocks from the underlying upper Mount

Ada Basalt. Based on immobile element concentrations (Zr/TiO2 – Nb/Y) the rocks are

classified as andesites, indicating that SiO2 enrichment was the result of intensive post- depositional silicification. Stratigraphic younging within the Duffer Formation correlates with increasing REE concentration levels. This supports a model in which the Duffer Formation silicic (andesitic) volcanic rocks represent successive outpourings as the final differentiation products of underlying Mount Ada Basalt tholeiites. 3) The stratigraphic affinity of mafic rocks in the Central Fold Belt remains unclear.

However, all rocks within the Central Fold Belt are affected by D2 deformation (Chapter 2), and therefore must be older than 3427 Ma (Chapter 2, 3). Because of the abundance of intrusive rocks in the Central Fold Belt, but rare occurrence of intrusive rocks in the Low- Strain Belt, I argue that the Central Fold Belt is part of the Mount Ada Basalt, or older.

134 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

The pillow basalt flows and massive mafic rocks in the Central Fold Belt do not show any geochemical difference to their equivalents in the Low-Strain Belt. Therefore, a petrogenetic relationship, or similar petrogenetic processes are considered. 4) Pillow basalt flows and massive mafic rocks from the Central Fold Belt and the Low- Strain Belt cannot be distinguished based on their geochemistry, and include high (10-15 wt. %)- and low (5-10 wt. %)- MgO types. Both high- and low-MgO types include flat REE and LREE enriched compositions. Based on the recognition of repeated, upward stratigraphic, systematic geochemical variation cycles of the most primitive high-MgO components, I argue for the presence of distinct volcanic cycles. From base to top, these

cycles display increasing TiO2 and LREE enrichment, while Mg#, Cr/Ni and Sc/Zr decrease. I explain this systematic magma evolution as the result of infra-crustal fractional differentiation processes (Model II), as opposed to a systematic emplacement of different mantle sources (Model I). Supported by petrographic and field observations of mafic – ultra-mafic cumulate rocks, such as anorthosites and anorthosite – clino- pyroxenite – gabbro layered bodies, and possibly olivine-chromite cumulate (i.e. serpentine-chromite schist), fractional crystallization processes are implied. Fractional crystallization within infra-crustal chambers resulted in the observed systematic magma

evolution and emplacement/eruption, controlled by plagioclase (decreasing Al2O3/TiO2), clinopyroxene (decreasing Sc/Zr, enrichment in LREE) and olivine (decreasing Mg# and

Cr/Ni and increasing TiO2) fractionation. Additional enrichment in incompatible elements (LREE, HFSE) may have resulted from some degree of crustal contamination (i.e. TTG basement, Chapter 3) during magma ascent.

135 Chapter 4 Geochemistry of the Mount Ada Basalt and Duffer Formation ______

136

Chapter 5 ______

SYNTHESIS AND DISCUSSION

In the introduction (Chapter 1) I outline major knowledge gaps and controversies regarding the formation and evolution of the early Archaean granite dome – greenstone keel focusing on the East Pilbara Terranes. These controversies are: I) The geometry, kinematics, and tectonic significance of an intense deformation phase predating the regional 3310 Ma dome-and-keel event in the East Pilbara Terrane; II) The timing, nature and formation of pre-dome-and-keel felsic crust; III) Lateral variability in the architecture, thickness and composition of the greenstones of the East Pilbara Terrane. I addressed these knowledge gaps in a detailed lithostratigraphic, structural, geochronological and geochemical study of the Muccan Granitic Complex dome – Doolena Gap greenstone keel transect in the three main chapters of the thesis. After a short summary of my main findings, the most significant outcomes from answering these knowledge gaps are synthesized in the tectonic evolution of the Doolena Gap greenstone belt. A discussion of larger-scale implications for the formation and evolution of the early Archaean crust follows. Finally, I point out remaining questions and how to address these in future research to further improve our understanding of early Archaean tectonics and crustal growth.

5.1 Summary of Key Findings of the Thesis

My lithostratigraphic and deformation analysis of the Doolena Gap greenstone belt inspired a comprehensive tectonic model of the research area. It entails structural and stratigraphic evidence for pre-3427 Ma dome-and-keel formation in the northern East Pilbara Terrane. For the first time, a detailed kinematic model for the development of the ubiquitous transposed tight to isoclinal folds is presented (D2 in the Central Fold Belt, Chapter 2). These folds likely correlate with early tight folds reported from other adjacent greenstone belts (e.g. Buick et al., 1995). This high-strain

D2 deformation accommodating doming in the middle crust (corresponding to the Central Fold Belt and South Muccan Shear Zone) is correlated with protracted upper-crustal response (preserved in the Low-Strain Belt): first EW-extension during deposition of the Mount Ada Basalt; Chapter 5 Synthesis and discussion ______then uplift during the eruption of the ~3460 Ma Duffer Formation; and finally 45˚ tilting of the entire greenstone stack ascribed to doming and keel descent prior to the formation of the <3427 Ma Strelley-Pool unconformity (Chapter 2 and 3). These observations present time constraints for partial convective overturn within ≤ 30 Ma, in agreement with numerical models (e.g., Robin and Bailey, 2009). This local tectonic model also explains the exceptional thickness of the Strelley Pool Formation in the area as a developing sub-basin, related to the formation of a greenstone drip at the nearby dome triple junction (Chapter 2). Regional variations in composition and thickness of greenstone formations can therefore be explained by differential spatial and temporal extension and uplift during progressive dome-and-keel development. An interesting observation not explored in detail here is the continuous formation of localized carbonate-alteration zones associated with all recognized deformation events. These metasomatic alteration zones may provide a possible mechanism for significant mechanical weakening of the greenstone keel. This interpretation may inspire future testing with numerical forward models. My petrological, petrogenetic, geochemical and geochronological studies reveal new insights into the formation of the igneous protoliths of the dome and keel. For the first time, a potentially co-magmatic suite (Doolena Suite) of 3499-3591 Ma granitoid gneisses is identified, and their petrogenesis is discussed. A model is favored in which parental TTG melts were formed in a zone of partial melting at the base of over-thickening hydrated mafic crust (>40 km depth). The parental melts infiltrated the mid- to shallow lower-crust where they crystallized as TTG and/or fractionated to derive granodioritic-dioritic cumulates and evolved potassic granite components. Detrital zircon provenance ages of the Strelley Pool formation reflect the evidence for continental material of up to 3500-3600 Ma age: they show an age peak that correlates with the emplacement ages of the Doolena Suite components. This observation lends further support to the notion of pre- 3427 Ma dome-and-keel formation, during which this ancient felsic basement was partly un- roofed and eroded (Chapter 3). Geochemical and petrological data from mafic greenstone rocks provides evidence for tholeiitic magma chambers, in which (ultra-)mafic mantle magma differentiated to produce various magma-types that erupted on and intruded the upper crust. A petrogenetic relationship between felsic volcanic rocks and underlying mafic rocks is demonstrated. The absence of TTG-like felsic volcanic rocks indicates that felsic dome magmatism does not contribute to near coeval greenstone volcanism (Chapter 4).

138 Chapter 5 Synthesis and discussion ______

5.2 Discussion

5.2.1 Tectonic evolution of the western Doolena Gap greenstone belt

The findings of the thesis in combination with the previously established regional model for the East Pilbara Terrane formation, suggest the following geological-tectonic evolution for the western Doolena Gap greenstone belt:

3591-3499 Ma: Doolena Suite stage - formation of the Doolena Gap greenstone belt basement

The crystallization ages of the 3499-3591 Ma Doolena Suite granitoid gneisses provide evidence for the existence of early granitic mid-crust, previously inferred from detrital zircon, isotopic data, and rare xenoliths (e.g. Hickman and Van Kranendonk, 2012; Van Kranendonk et al., 2015). Petrological and geochemical constraints for the Doolena Suite suggest that these early rocks were derived from fractional crystallization of a TTG parental melt that was generated through partial melting of hydrous metabasalt at the base of over-thickening crust at depth of ca. 45-50 km. The ancient melts infiltrated the mid- to shallow lower levels of a thick vertically growing crust, and crystallized to form the basement to the supracrustal rocks of the Doolena Gap greenstone belt.

3490-3470 Ma: Callina Supersuite/Talga Talga Subgroup stage - not recognized in the research area: unidentified, non-deposition, or tectonic removal?

No evidence for rocks belonging to the regional granitic 3490-3460 Ma Callina Supersuite (e.g. Van Kranendonk et al., 2006) or to the supracrustal ~3490-3480 Ma Talga Talga Subgroup of the lower Warrawoona Group (e.g. Hickman, 2012) have been identified. However, central parts of the Muccan Granitic Complex to the north of the study area have been mapped as belonging to the Callina Supersuite (Van Kranendonk et al., 2006; Van Kranendonk, 2010). Metabasalts from the westerly adjacent Warralong greenstone belt and from the southerly adjacent Marble Bar greenstone belt have been ascribed to the 3490 Ma North Star Basalt (Talga Talga Subgroup; e.g. Beintema, 2003; Van Kranendonk, 2010). It remains unclear if part of the studied South Muccan Shear zone or Central Fold Belt belong to formations of this age range, or if the lowermost Warrawoona Group greenstones were lost during later dome-and-keel tectonics in the Doolena Gap greenstone Belt. In fact, it is possible that older greenstone successions were deposited in the Doolena Gap area prior to the Mount Ada Basalt and Duffer Formation. Tectonic removal may be expected particularly if the lower successions were formed by more primitive (e.g. komatiitic) material. Numerical models of Rayleigh-Taylor instability delamination drips of lower crustal material into underlying mantle have demonstrated that the most negatively buoyant material will sink, while positively buoyant material remains at the crustal base (e.g. Johnson et al., 2014). In

139 Chapter 5 Synthesis and discussion ______

combination with the higher petrophysical susceptibility of primitive (ultramafic) rocks to weaken during low-temperature alteration (e.g. Hooper and Hatcher, 1989), it may be expected that such rocks were lost through sagduction during partial convective overturn, while mechanically stronger layers were preserved. Within the Central Fold belt of the Doolena Gap greenstone belt, wide carbonate-altered deformation zones penetrate the greenstone succession parallel to the inferred original bedding-planes (Chapter 2). The carbonate zones accommodated dome-up vertical displacements during non-coaxial shear deformation. It is likely that significant amount of stratigraphy was lost during deformation along these alteration zones. It is proposed that these zones accommodated strain through syn- deformational alteration weakening of mafic volcanic rock (Chapter 2), possibly ultramafic.

3470-3460 Ma: Coongan Subgroup stage - formation of the upper crust, initial doming, and the emergence of early life in the Doolena Gap greenstone belt

The Mount Ada Basalt and Duffer Formation supracrustal deposits (Coongan Subgroup) represent the well preserved and stratigraphically constrained (Low-Strain Belt) main phase of volcanism and upper crustal growth in the studied Doolena Gap greenstone belt. The Mount Ada Basalt was formed through the vertical stacking of pillow basalt flows and sub-volcanic intrusions and sills. Volcanological constraints imply that the pillow flows erupted in a deep- marine extensional environment. Geochemical and petrological data suggests that compositionally distinct types of pillow basalts and interspersed massive mafic rocks and shallow intrusions were derived from fractional crystallization of more primitive mantle- derived melts (Chapter 4). Parental primitive (komatiitic?) melts, however, are not observed. It is proposed that parental melts may not have been able to penetrate the ductile mid-crust, which underlay the study area at that time and was continuously infiltrated by granitic magma of the Callina Supersuite. Alternatively, as discussed above, primitive volcanic rocks may have formed part of the lowermost Mount Ada Basalt, but were tectonically removed during subsequent partial convective overturn (see next stage, below) due to their high-density and low-viscosity. Compositional types of mafic rocks were derived from intra-crustal fractionation chambers. The surficial volcanism likely occurred in cycles, as indicated by intercalated chert horizons, starting with explosive eruptions, followed by outpouring pillow flows with a change from more primitive to fractionated melt compositions. The melts erupted through an extensive upper crustal fracture system. East-west syn-depositional extension in the upper crust of the Low-Strain Belt is interpreted as a response to the initial expansion and

doming of the granitic mid-crust below. Similarly, the only locally preserved earliest S1 foliation of the Central Fold Belt is ascribed to vertical flattening and propagation-related lateral flow across the interface of the nascent dome and overlying greenstones within lower upper- and mid-crustal levels. East-west crustal extension of the upper crust continued past the cessation of mafic volcanism in the Low-Strain Belt and gave rise to outpourings of the final differentiation (i.e. fractionation) products that form the silicic volcanic portion of the

140 Chapter 5 Synthesis and discussion ______

Duffer Formation in the eastern study area. Protracted doming finally led to the uplift of the marine extensional volcano-sedimentary basin, reflected in the regressive sedimentary sequence in the western basin-forming part of the Duffer Formation. Ultimately, the upper crust of the Low-Strain Belt emerged in a shallow-marine environment, where the earliest proposed life forms (i.e. stromatolites; e.g. Schopf et al., 2007) were able to flourish in the Doolena Gap greenstone belt.

3460-3427 Ma D2: early partial convective overturn and associated reworking of the ancient basement

The 45˚ angular unconformity underneath the Strelley Pool Formation indicates that the >3460 Ma greenstone belt experienced a relatively rapid tilting to the south until ~3427 Ma. The tilting is consistent with kinematics of the ascent of the Muccan Granitic Complex dome to the north and coeval sagduction of the greenstones. The relative upward movement of the dome resulted in a wide ductile shear zone where both dome- and greenstone-belt components were tectonically imbricated and mylonitized within a dominantly non-coaxial strain regime (South Muccan Shear Zone). The S- and SL-fabrics near the dome-keel interface turn into L- tectonites indicating constrictional flow towards the south-southwest. This is interpreted as a consequence of a possible greenstone drip near the triple junction between the Muccan, Mount Edgar and North Pole Granitic Complex domes. The geometrical relationship between greenstone drips and dome triple junctions has also been observed within the lower crust of the Western Dharwar Craton and in the Barberton granite-greenstone belt (e.g. Anhaeusser, 1984; Bouhallier et al., 1995; Anhaeusser, 2001), and is possibly recognized in the foliation triple junction west of the Warrawoona Syncline in the East Pilbara Terrane (e.g. Collins et al., 1998). The cause of the final rapid overturn of the granitic mid-crust remains unclear, but was associated with low-temperature magmatic reworking of the most ancient components. At the dome margin, partial melting was likely initiated by volatile influx that derived from continuous magmatic activity in the dome center. It is proposed that continuous granitic melt influx from the base of the crust, in combination with continuous deposition of supracrustal volcanic rocks, ultimately led to a gravitational crustal instability. The presented detrital zircon provenance ages suggest that the mid-crustal basement represented by the Doolena Suite was being exposed during this event. Based on commonly accepted pressure estimates for the emplacement of epidote-bearing dome components (e.g. Zen, 1984), this indicates an exhumation of at least 25-30 km. It remains unclear if the Salgash Subgroup that represent the uppermost Warrawoona Group in the East Pilbara Terrane was deposited at all and eroded since the onset of the tilting, or if the crustal uplift or some other mechanism prevented the deposition of the uppermost Warrawoona Group in the Doolena Gap greenstone belt. Possible indications for sub-aerial weathering and hence erosion are found below the <3427 Ma angular unconformity.

141 Chapter 5 Synthesis and discussion ______

3427-3350 Ma Strelley Pool Formation stage

From ca. 3427 Ma onwards, massive amounts of clastic sediments started to fill the developing Strelley Pool Formation (sub-)basin to the south. The most ancient TTG components and those that were emplaced during the partial convective overturn event were eroded and their detritus shed into this sub-basin. Within the Doolena Gap greenstone belt, the Strelley Pool Formation presented a shallow-water fluviatile to marginal environment, as evident in interbedded conglomerate layers and cross-beddings. The sediment was likely transported in rivers from proximal sources of the surrounding topographically elevated domes (particularly the Muccan Granitic Complex) and shed into the basin where the sandstone matured in shallow water (Chapter 2 and 3). The local accommodation of up to ~1 km of sandstone may have been possible due to ongoing subsidence over the sinking greenstone drip as a late-stage response to partial convective overturn. Emplacement ages of granitic magmatism within the domes are virtually absent between 3420 and 3330 Ma in the East Pilbara Terrane (e.g. Van Kranendonk et al., 2006), and the ca. 75 Ma in which the Strelley Pool Formation was deposited marks a hiatus in supracrustal volcanism (e.g. Hickman, 2012). It remains unclear why there is no record of magmatic activity during the Strelley Pool Formation stage. On first hand, it may be speculated that depletion of a fertile mantle source at least temporally inhibited the continuation of supracrustal volcanism and hence vertical crustal thickening, which in turn prevented the generation of granitic melts at the lower crustal base. If this were true, one would expect a possible coupling mechanism between the cessation in mantle and crustal magmatism and the coinciding partial convective overturn. It has been proposed that surficial outpourings of TTG affinity did not occur until the uppermost Warrawoona Group (Panorama Formation; e.g. Smithies et al., 2007, 2009). This TTG volcanism coincided with final TTG magmatism of the Tambina Supersuite (e.g. Van Kranendonk et al., 2006). It is therefore possible that voluminous generation of magma from the mantle already ceased just before the final Warrawoona Group volcanism took place (during ca. lower Salgash Subgroup). Subsequently, the final pre-3420 Ma TTGs were generated from the lower crustal base. The vanishing contribution from mantle magma into and onto the crust may have allowed the supracrustal pile to cool and support the ascent and eruption of the TTG magma through fissures in the brittle upper crust. This coincided with the final felsic volcanism that derived from tholeiitic differentiation of the last mantle-derived melts that had slowly penetrated the thick crust, producing the compositionally distinct volcanic centers of the Panorama Formation (e.g. Smithies et al., 2007). As evident in the Doolena Gap greenstone belt, mid-crustal doming initiated during the Coongan stage (Chapter 2), marking the presence of a crustal instability at that time. However, rapid (>30 Ma) partial convective overturn did not occur until after the deposition of the Duffer Formation in the Doolena Gap greenstone belt. I therefore speculate that temporal termination of mantle magmatism since the late Warrawoona Group may have played a significant role in

142 Chapter 5 Synthesis and discussion ______

contributing to the rapid nature of the partial convective overturn. If mantle-derived volcanic stacking on top of the crust lasted, longer time-scales for partial convective overturn would be expected. In fact, numerical models that demonstrate rapid partial convective overturn are most often based on the parameterization of a crustal rheology comprising cold dense greenstones on top of felsic substrate (e.g. Sandiford et al., 2004; Thébaud and Rey, 2013). It would be interesting to explore partial convective overturn in future numerical models that incorporate more variability in the addition of both granitic mid-crustal magma and mafic greenstone volcanism.

3310 Ma D3: regional partial convective overturn?

To the south of the study area, conformable greenstone formations are ascribed to the younger Kelly Group, indicating renewed volcanic activity (e.g. Euro Basalt; Van Kranendonk, 2010).

The D3 deformation event in the studied part of the Doolena Gap greenstone belt can be geometrically related to further dome-and-keel development (Chapter 2). It is likely that this event coincided with the regional partial convective overturn event in the East Pilbara Terrane at ca. 3310 Ma (e.g. Sandiford et al., 2004; François et al., 2014; Van Kranendonk et al., 2015). The event occurred during the final stages of Kelly Group greenstone deposition that spawned further and/or renewed gravitational instability within the crust. During this event the Doolena Gap greenstone belt experienced further westward rotation towards the dome- triple junction between the Muccan, Mount Edgar, and North Pole Granitic Complexes. This indicates that the renewed sagduction of greenstones was likely most intensive where the prior (>3427 Ma) greenstone drip occurred. Given the above interpretation that the Strelley Pool Formation in the Doolena Gap greenstone belt developed during final subsidence of the early partial convective overturn, it is likely that massive Kelly Group greenstone deposition further filled the developed basin and hence supplied renewed supracrustal loading on top of the old drip. In contrast to the reworking mechanism of granitic material through volatile and heat supply from continuous infiltration of juvenile granitic magma into the mid-crust, as proposed in Chapter 3 for the early partial convective overturn, post-3330 Ma TTG and granitic rocks have been interpreted to represent reworked older TTG crust without the presence of juvenile magmatic input (e.g. Collins, 1993; Smithies et al., 2003). In the latter case, softening of the older (>3420 Ma) granitic mid-crust and associated generation of partial melts that buoyantly rose to support doming, has been interpreted as the result of a combination of i) heating from the radioactive element-enriched older TTG crust through blanketing, ii) heating from mantle plume activity, and iii) burial of shallow granites into warmer mid-crust (e.g. review in Van Kranendonk et al., 2015). This is in agreement with the fact that sagduction occurred at the sights of prior partial convective overturn related drips and further doming affected the already dome-structured granitic crustal layer. As evident in later angular unconformities (see below; Van Kranendonk, 2010), the original site of most intensive Doolena Gap greenstone

143 Chapter 5 Synthesis and discussion ______

sagduction within the domal triple junction to the SW also formed the zone of greenstone sinking during later partial convective overturn events.

Meso- to Neoarchaean evolution

Late-stage structures in the Doolena Gap greenstone belt include the D4 sinistral NE-SW faulting event that can be traced on satellite images to affect rocks of the <3050 Ma De Grey

Supergroup, and the D5 dextral EW-structural reactivation that must have occurred after the

intrusion of the 2772 Ma Black Range Dyke swarm (Wingate, 1999). The D5 event marks the final juxtaposition of the Central Fold Belt and the Low-Strain Belt. The time constraints and

the inferred compressional regime during both the D4 and the D5 event suggest structural imprints and reactivation after the initiation of plate tectonics (~3200 Ma) when the East Pilbara Terrane formed part of the Pilbara Craton (e.g. Hickman and Van Kranendonk, 2012).

5.3 Remaining Questions and Future Research Directions

1) East Pilbara Terrane-wide and possible global correlation of the early partial convective overturn event This thesis conclusively demonstrates that early partial convective overturn affected the studied western Doolena Gap greenstone belt between ca. 3460 and 3420 Ma, prior to the regional main event around 3310 Ma. Previous studies from other localities within the East Pilbara Terrane have presented evidence of deformational structures below the 3427 Ma Strelley Pool Formation (SW Carlindi Granitic Complex; Buick et al., 1995), and metamorphic monazite growth ages of ~3443 Ma (Warrawoona Syncline/Mount Edgar Granitic Complex; François et al., 2014). This suggests that the early partial convective overturn may have affected the early East Pilbara Terrane crust on a regional scale. In order to test this I propose to undertake further detailed field and geochronological studies throughout the East Pilbara Terrane. Such studies should focus on the oldest deformed greenstone packages unaffected from late-stage overprinting within synclinal keel centers. Correlation between the East Pilbara Terrane and the Barberton granite-greenstone belt suggest that deformation affected both terrains at ca. 3220, 3310 and 3430 Ma (Van Kranendonk et al., 2015). The ca. 3430 Ma deformation in the Barberton area has not been ascribed to partial convective overturn so far. Future studies should focus on testing if this deformation event was the result of partial convective overturn in the Barberton area, following the example put forth in this thesis for part of the East Pilbara Terrane. In case the events can be correlated, this has large-scale implications, further underlining a proposed East Pilbara – Barberton palaeo-geographic affiliation, or even a global coincidence of tectonic events.

2) Incomplete geochronological record and the time-frames and interaction of tectono-magmatic processes

144 Chapter 5 Synthesis and discussion ______

As demonstrated in this thesis based on the relative timing of overprinting deformation in relationship to the stratigraphic position of affected rocks, early partial convective overturn occurred within 30 Ma or less, in agreement with previous numerical models. Most numerical models, however, do not account for the formation of the required crustal configuration to initiate partial convective overturn. The gravitational instability that led to the first (recognized) crustal re- organization occurred towards the end of an early phase of >100 Ma-duration of almost continuous juvenile crust formation in the East Pilbara (i.e. Warrawoona Group, Callina and Tambina supersuites). Future models need to account for both progressive granitic magmatic input and contemporaneous supracrustal loading in order to better constrain the time frames required to develop observed crustal instabilities. While current models explain the later 3220 and 3310 Ma partial overturn events, in which pre-existing mid-crustal granitoid basement was buried and blanketed through supracrustal mafic volcanism (e.g. Sandiford et al., 2004), the early 3460-3420 Ma instability occurred during juvenile magmatic input. In combination with future mapping and extension of geochronological databases, numerical models could be used to estimate the volumes of required granitic melt production during continuous thickening of mafic crust until instabilities develop. The volume and rate of continental crustal growth represent some of the most debated controversies in early Earth research (e.g. Hawkesworth et al., 2016). As demonstrated in this thesis, the geochronological record of the East Pilbara Terrane is still incomplete. A more comprehensive geochronological record would allow us to better constrain the duration of both granitic magmatism and supracrustal volcanism, and their respective volumes. These data need to be integrated into advanced numerical models, such as performed by Johnson et al. (2014), who explored the time frames of development of lower crustal instabilities and associated delamination drips. What impact does delamination of melt depleted lower crustal drips have on the continuation of mantle volcanism and granitic melt extraction? Are delamination processes reflected in age gaps of granitic emplacements?

3) Provenance of Archaean detrital zircon I argue for a simple pattern of sediment dispersal sourced from the Muccan Granitic Complex dome into proximal dome-circumferential basins represented by the sinking greenstone keels during the early partial convective overturn event. The exceptional preservation of the East Pilbara Terrane would allow testing of this statement. Future studies are proposed to investigate palaeo- currents based on sedimentary structures in the Strelley Pool Formation. If palaeo-current directions correlate geometrically to the dome, can this be shown elsewhere around the Muccan Granitic Complex or elsewhere in the East Pilbara? Such studies could help to solve a long debate on the provenance of Archaean zircon and the source of host bimodal greenstone sediments. In the East Pilbara and many other Archaean terrains, the detrital zircon record exceeds the oldest emplacement ages within granitic domes. I propose future studies linking zircon chemistry with host rock composition and petrogenesis, to better constrain provenances of detrital grains.

145 Chapter 5 Synthesis and discussion ______

Furthermore, more data on the spatial distribution and kinematics of syn-depositional extensional and/or normal faults is deemed valuable. As previously demonstrated by Nijman and de Vries (2001), such data can be used to reconstruct basin geometries and extents.

4) Stratigraphic correlations Particularly in Chapter 4 of this thesis I demonstrate the problematic in interpreting petrogenetic relationships and processes of mafic greenstone rocks, which are not conclusively dated. Future advances in dating of other accessory minerals, such as titanite and rutile, may help to constrain the ages of mafic greenstone formations that do not comprise zircon-bearing lithologies.

5) Repeated partial convective overturn events and the transition to plate tectonics In the East Pilbara Terrane, plate tectonics initiated with Wilson Cycle rifting ca. 3200 Ma ago (Van Kranendonk et al., 2010). Including the early partial convective overturn event evident in this thesis, a total of three partial convective overturn events are recognized in the East Pilbara prior to the onset of plate tectonics. As outlined in Chapter 1, dome-and-keel tectonics coincided with predicted peak mantle temperatures. Mantle potential temperature estimates between the peak at ca. 3500 Ma and the onset of plate tectonics at 3200 Ma, however, are not too drastically different (~20 °C; Korenaga, 2013). Could such little drop in mantle temperatures have caused the lithosphere geodynamics to change so significantly? Numerical simulations have demonstrated that the tectonic state of a planet depends on its evolutionary pathway (Weller and Lenardic, 2012; O’Neill et al., 2016). I propose to explore the possibility that the evolution of the lithosphere through repeated partial convective overturn played a significant role in developing thermal- mechanical properties in favor of Wilson-style rifting, in future numerical simulations.

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Appendix 1

Lithostratigraphy and structure of the early Archaean Doolena Gap Greenstone Belt, East Pilbara Terrane (EPT), Western Australia Wiemer, D., Schrank, C. E., and Murphy, D. T. American Geophysical Union, Fall Meeting, San Francisco, 2014, poster abstract #V43C-4901 Bibliographic Code: 2014AGUFM.V43C4901W

Abstract We present a detailed lithostratigraphic and structural analysis of the Archean Doolena Gap greenstone belt to shed light on the tectonic evolution of the EPT. The study area is divided into four structural domains: i) marginal orthogneisses of the MGC (Muccan Granitoid Complex), ii) a dominantly mafic mylonitic shear zone (South Muccan Shear Zone, SMSZ) enveloping the MGC, iii) a Central Fold Belt of dominantly mafic greenschists (CFB), and iv) a lower greenschist- to sub-greenschist southern domain. Toward the dome margin, abrupt increases in deformation intensity occur across domain boundaries. Domain boundaries and intra-domain shear zones are marked by significant carbonate +/- quartz alteration and high-strain non- coaxial deformation with dome-up kinematics. The southern domain comprises pillow basalts of the Mount Ada Formation (MAF), conformably overlain by clastic sediments and minor pillow basalts of the Duffer Formation (DF). The MAF and DF are overlain by an up to 1km thick package of quartzite (Strelley Pool Formation) across an angular unconformity. Isoclinal folds (F2) within the CFB to the North deform an early foliation (S1) within dominantly mafic schists and associated carbonate veins. F2 folds are preserved within lozenges that are parallel to the axial planes of F2 folds in a regional E-W trending foliation (S2) and to the SMSZ. Lozenges are often bound by zones of significant carbonate alteration. The lozenges are folded recumbently (F3), with sub-vertical fold axes pointing towards the dome. The F3 axes are parallel to mineral stretching lineations on S2 indicating dome-up movement. The entire belt is cut by late NE-SW-striking faults that exhibit dominantly brittle deformation in the southern domain but ductile drag folding (F4) in the CFB. Therefore, the southern domain must have overlain the CFB during this D4 event. We propose a protracted structural history of the greenstone belt where successive deformation events relate to the episodic emplacement of the MGC. We demonstrate that the greenstone keel is mainly characterised by an anastomosing shear zone network, induced by hydro-chemical weakening of mafic schists. This implies that previous estimates of stratigraphic thickness are significantly overestimated.

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Appendix 2

Pillow basalts from the Mount Ada Basalt, Warrawoona Group, Pilbara Craton: Implications for the initiation of granite-greenstone terrains Murphy, D. T., Trofimovs, J., Hepple, R. A., Wiemer, D., Kemp, A. I. S., and Hickman, A. H. Goldschmidt, Prague, 2015, talk abstract Presenter: D. Wiemer

Abstract The Pilbara Craton represents the archetypal Archean granite-greenstone terrain in which mafic volcanic dominated supercrustals are intruded by granitic domes. This crustal morphology reflects distinct tectonic settings that formed in a hotter early Earth. The ambient temperature in the Paleoarchean mantle is estimated to be 1600 °C [1] and corresponds with the liquidus temperature of Barberton komatiites [2]. In the Paleoarchean mantle a pyrolite composition at depths of less than 100 km is expected to melt and generate ultramafic magmas. Here we present volcanology, petrology and geochemistry data of well-preserved basaltic lavas ascribed to the Mount Ada Basalt, Warrawoona Group from the Doolena Gap Greenstone Belt. The Mount Ada Basalt was coeval with the Callina plutonic event that marks the initiation of dome formation in the Pilbara Craton [3]. The Doolena Gap sequence is exclusively pillow basalts with MgO < 10%. Isotopically the basalts are indistinguishable from contemporary non-chondritic Bulk Earth (εNd, 1.0 ± 0.2 and εHf, 2.3 ± 0.2). Here we address the implications of Paleoarchean basalts with MgO% < 10 derived from melting of a source indistinguishable from non-chondritic Bulk Silicate Earth to the initiation and subsequent evolution of the Pilbara Craton.

[1] Korenaga, J. (2008) Reviews of Geophysics, 46 (2008) [2] Puchtel, I. et al. (2013) Geochimica et Cosmochimica Acta 108, 63-90 [3] Hickman, A. H. (2012) Island Arc 21, 1-31

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Appendix 3

Structural development of the early Archaean Doolena Gap greenstone belt, East Pilbara Terrane (Western Australia) Wiemer, D., Schrank, C. E., Murphy, D. T., and Hickman, A. H. Specialist Group in Tectonics and Structural Geology, The Geological Society of Australia, Caloundra, 2015, poster abstract

Abstract Regional crustal-scale dome-and-keel structures in the Archaean East Pilbara Terrane (EPT, Western Australia) are explained through gravitational overturn of buoyant felsic substrate overlain by dense mafic greenstones [1]. In the southeastern EPT, dome-and-keel development is thought to postdate deposition of the ~3.33 Ga Euro Basalt. Thermal softening of buried radiogenic 3.47-3.43 Ga granitic mid-crust triggered vertical re-organization through sagduction of greenstones and buoyant rise and expansion of domes [2]. Older (>3.43 Ga) deformational structures have been reported locally, but their development remains largely unexplained [3]. Here, we present a detailed lithostratigraphic and structural analysis of a granite-greenstone traverse in the western Doolena Gap greenstone belt (DGGB) of the northern EPT. The traverse reveals four structural domains: i) orthogneisses of the polyphase Muccan Granitoid Complex (MGC), ii) the mylonitic South Muccan Shear Zone (SMSZ), iii) a Central Fold Belt (CFB) of dominantly mafic greenschists, and iv) a well- preserved southern Low Strain Belt (LSB) featuring pillow basalts and intercalated metasediments. From the keel to the dome margin, metamorphic facies and deformational intensity increase across domain boundaries. Domain boundaries and anastomosing intra-domain shear zone networks are marked by significant carbonate ± quartz alteration and high-strain non-coaxial shear deformation with dome-up kinematics. The studied greenstone successions have been ascribed to the 3.47 Ga Mount Ada Basalt (MAB) and the Duffer Formation (DF), overlain by the <3.43 Ga Strelley Pool Formation (SPF) across an angular unconformity [4] that suggests ~45° tilting of the MAB and DF prior to ~3.43 Ga, provided that the stratigraphic correlation with the SPF is correct. We interpret this tilting as a consequence of doming of the MGC to the North.

Geometry and shape of syn-anatectic tight folds within the MGC resemble those of isoclinal folds (F2) within the CFB. We therefore propose that partial melting within the MGC triggered its buoyant rise and expansion, resulting in D2 passive shear folding and heterogeneous shear deformation accompanied by influx of carbonaceous fluids along discrete deformation zones within the CFB, and development of the mylonitic

SMSZ at the dome-keel interface. Within the CFB, prolate fabrics indicated by D2 L-tectonites become increasingly common towards the South. In conjunction with mineral lineations (L2) this may indicate constrictional flow (sagduction) of the keel rocks to the SW, likely coupled with the early doming. F2 folds deform an earlier localized shear foliation (S1) and associated D1 carbonate veins, which possibly record the initiation of the main D2 event along the flanks of the rising MGC. Structures post-dating ~3.43 Ga SPF deposition include asymmetric, open upright folds (F3) superimposed on earlier structures within the CFB, and shear zone networks penetrating the LSB including the SPF. D3 indicates tightening of the D2 greenstone keel associated with protracted dome-up dextral flow towards a dome triple junction to the SW. The entire belt is cut by late NE-SW-striking faults (D4) that exhibit dominantly brittle deformation in the LSB but ductile drag folding (F4) in the CFB. Therefore, the LSB may have overlain the CFB during this D4 event. The two domains were juxtaposed during late-stage oblique strike slip (D5) along local reactivated S2 foliation and carbonate zones. In summary, we propose a structural history of the DGGB in which successive deformation events can be related to the episodic emplacement of the MGC and associated sagduction of the keel rocks. Our study suggests that dome-and-keel initiation commenced at least locally prior to ~3.43 Ga. Polyphase localized thermo-hydro-chemical weakening of mafic greenstones accommodated strain during the deformational episodes within the keel.

[1] Collins et al., 1998. Journal of Structural Geology, 20, 9/10, pp. 1405-1424 [2] Sandiford et al., 2004. Tectonics, 23, TC1009, doi:10.1029/2002TC001452 [3] Van Kranendonk et al., 2002. Economic Geology, 97, pp. 695-732 [4] Van Kranendonk, 2010. Geology of the Coongan, Geological Survey of Western Australia Report

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Appendix 4

Carbonate alteration is the dominant weakening mechanism in the Doolena Gap Greenstone Belt preserving the keel rock in a dome-and-keel terrane Burke-Shyne, D., Wiemer, D., Schrank, C. E., and Murphy, D. T. Specialist Group in Tectonics and Structural Geology, The Geological Society of Australia, Caloundra, 2015, poster abstract

Abstract We hypothesise that strain localisation facilitates large vertical displacements in dome-and-keel structures by a coupled thermo-hydro-mechanical process controlled by carbonate alteration of the mafic keel rocks. The partial crustal overturn of these dome-and-keel structures is often ascribed to a Rayleigh Taylor (RT) instability [1]. Numerical models suggest that RT instabilities can only progress in dome-and-keel terranes if the effective viscosity of the mafic rock decreases such that localised shear zones form, allowing for protracted periods of deformation accommodating large strain [2, 3]. We examined anastomosing shear zones and preserved lozenges of dense keel rock exposed in the Doolena Gap Greenstone Belt (DGGB), East Pilbara Terrane (EPT, Western Australia), to test our hypothesis. The DGGB provides an excellent study area because it represents the synclinal limb of a greenstone keel, containing dominantly mafic schists and minor felsic volcanics and meta-sediments, in which a brittle to ductile gradient of deformation is observed towards the domal Muccan Granitoid Complex. The increasing deformational intensity within the mafic keel rocks and associated occurence of localised anastomosing carbonate shear zones record large vertical displacement between dome-and-keel while preserving the early structures associated with dome-and-keel formation due to not being completely overprinted by subsequent deformation events. We present microstructural and geochemical analyses of three detailed transects across anastomosing shear zones within the most deformed part of DGGB greenstones, the Central Fold Belt, which abuts the mylonitic shear zone separating the Muccan dome in the North from the keel rocks [4]. Carbonate alteration is spatially associated with shear zones at all scales and with all deformation phases, as indicated by [4] overprinting relationships. Here, we focus on the deformation event (D2 after ), which accommodated the largest amount of greenstones displacement relative to the Muccan dome. D2 displays highly transposed, rootless isoclinal folds in an anastomosing shear zone network, which locally preserves ultramafic to mafic intrusives and extrusives in metric to decametric lozenges. The studied transects feature three 10-m wide shear zones with carbonate alteration bounded by relatively less deformed lenses of mafic rock exhibiting remnant igneous textures. We investigate the spatial correlation of relative strain intensity, microfabrics, deformation mechanisms, and trace- and major-element geochemistry to unravel the interplay of carbonate alteration and strain localisation. The mafic low-strain lenses comprise mafic schists that have been altered by carbonate-rich fluids under greenschist-facies conditions. Carbonate rocks within the shear zones represent alteration products of the mafic protolith having experienced almost complete replacement by carbonate minerals gradationally transitioning to mylonitic carbonates at the centre. The shear zones exhibit a gradual increase of foliation intensity towards their centres, which is interpreted as an increase in finite strain. The mafic schists in the low strain domain are composed of dominantly chlorite and calcite ± dolomite with minor quartz and opaques. Within the shear zones the composition is almost entirely calcite ± dolomite with minor quartz, opaques, goethite and accessory micas. The mineralogical transition displays a sharp disappearance of chlorite at the mafic schist to carbonate contact replaced by calcite, with gradually decreasing quartz towards the centre of the shear zones. Some relict magmatic textures are observed within the mafic schists. These remnant igneous textures gradually disappear in the carbonate towards the centre of the shear zones. The S2 foliation is defined by cleavage domains of chlorite and carbonate minerals, anastomosing around microlithons of microcrystalline quartz. Deformation mechanisms on the micro-scale include the shape preferred orientation of chlorite and carbonate minerals and significant grain size reduction in quartz by bulging and subgrain rotation and similar mechanisms in carbonate minerals. From major element geochemical analysis there is a clear trend of decreasing SiO2 and associated relative increase in CaO towards the centre of the shear zones. Trace element analysis shows that the carbonate material in the shear zones is derived from a mafic protolith. The observed spatial assocation of increasing foliation density with an increase in dynamic recrystallisation and the proportion of carbonate minerals suggests that carbonate replacement played a major role during strain localisation. Ongoing work focuses on identifying the source and composition of fluids associated with the multiple carbonate replacement events.

[1] O'Neill, C. and Debaille, V. (2014) Earth and Planetary Science Letters 406, 49-58. [2] Thébaud, N. and Rey, P. F. (2013) Precambrian Research 229, 93-104. [3] de Bremond d'Ars, J. et al. (1999) Tectonophysics 304, 29-39. [4] Wiemer, D. et al. (2014) AGU Fall Meeting 2014

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Appendix 5

Improving LA-ICP-MS dating techniques: experiments on zircon from a 3.51 Ga dioritic gneiss, East Pilbara Terrane, Western Australia Allen, C. M., Wiemer, D., and Murphy, D. T. The Geological Society of America, Annual Meeting, Denver, 2016, talk abstract

Abstract The “gold standard” U-Pb zircon dating method is ID-TIMS because highly precise and accurate ages obtain when Pb and U are purified, concentrated and calibrated using isotope dilution. Mattinson (2005) suggested that more concordant results ensue following annealing and chemical abrasion. “Spot” or in-situ dating methods are less accurate and precise because of “matrix effects”, a catchall for the subtle differences in composition among standards and unknowns. For LA-ICP-MS dating, Marillo-Sialer et al. (2014) and Crowley et al. (2014) have shown that a variety of factors result in age inaccuracies, but that the dominant effects stem from the ablation processes. We have studied the effect of annealing and/or chemical abrasion as compared to pristine ~3.5 Ga zircons in order to understand the impact of treatments on ancient zircons and the standards used to calibrate the isotope ratios (ages). We suggest that published reports of reverse discordance among Archean zircon commonly are caused by calibration artefacts and not common Pb. We propose improvements in the LA-ICP-MS dating workflow partially based on experiments conducted on large euhedral zircons from a dioritic gneiss of the Muccan Granitic Complex, East Pilbara Terrane. The weighted mean 207Pb/206Pb age (with annealed Temora2 as the reference standard) yielded 3.51 ± 0.03 Ga (n=37), making it the oldest documented coherent intrusion in the region. Zircon trace element concentrations help constrain petrogenesis. Ti-in-zircon thermometry yielded 720 ± 30 oC; Ca content shows positive correlation with domains of high alpha-dose and associated radiation-induced microstructural damage, and is useful for detecting discordant (Ca>200 ppm) grains.

Crowley et al. (2014) Minerals, 4, 503-518 Marillo-Sialer et al. (2014) Journal of Analytical Atomic Spectrometry 29, 981-989 Mattinson (2005) Chemical Geology 220, 47-66

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Appendix 6

Effects of thermal annealing and chemical abrasion on the U-Pb isotopic systematics, and the microstructure of complex, metamict ~3.5 billion year old zircon: insights for U-Pb LA-ICP- MS dating Wiemer, D., Allen, C. M., Murphy, D. T., and Kinaev, I. Australian Microbeam Analysis Society (AMAS), 14th Biennial Symposium, Brisbane, 2017, abstract submitted

The power of zircon in U-Pb geochronological application incorporates its own caveat: self-irradiation by α- decay of U and Th damages the microstructure, hence reducing the physico-chemical durability of zircon over time. Vulnerable to chemical modification, especially at near surface crustal conditions, affected zircon can suffer isotopic disturbance, complicating U-Pb age determination. Pre-analytical thermal annealing (TA) and chemical abrasion (CA) to recover crystallinity and remove high-damage portions has been successfully applied to the “gold standard” ID-TIMS dating method. Optimum workflows for the more cost- and time- efficient LA-ICP-MS spot analytical technique are debated [1]. We present a microstructural and U-Pb systematics study comparing pristine, TA and TA+CA treated ~3.5 Ga zircon from a quartz dioritic gneiss (East Pilbara Terrane, WA), with the aim to improve pre-analytical workflows for more accurate and precise LA-ICP-MS U-Pb dating of particularly ancient zircon. Based on CL imaging and trace element contents, four zircon domains are identified: i) low- to medium-U concentric- oscillatory zoned cores, ii-iii) two porous high-U outer domains (alt-1, alt-2), and iv) low-U narrow inward growing outermost rims (Fig.1a). Raman spectroscopy on the pristine zircon reveals a positive correlation between increasing structural damage and U content. The porous high-U outer domains show a drastic increase in non-formula Ca above estimated amorphous fractions of ~0.8 (Fig.1b). We ascribe this to hydrothermal alteration of high-damage zircon that is characterized by percolating networks of amorphous areas. Upon treatment (TA, CA), crystallinity is improved in all domains, as evident in de-colorization, improved translucence, and recovery of CL intensity. Hyper-spectral CL and Raman suggest structural recovery of point defects, but not full repair of high-damage amorphous areas. Calculated Raman radiation damage ages suggest that natural annealing affected all domains at 504 ± 19 Ma. This is consistent with our observation that laser ablation matrix effects are negligible between pristine cores and TA cores, which show little difference in U-Pb systematics. Both pristine and TA cores are reversely discordant and yield negative lower intercept ages (Fig.2a). In contrast, high-U alteration domains are normal discordant and display dispersion along Discordia with upper intercept ages of 3413 ± 62 (alt-1) and 3409 ± 110 Ma (alt-2), respectively. Lower intercepts are roughly in agreement with the 500 Ma-event. Upon CA, U-Pb discordance and scatter are reduced in the cores, yielding intercepts of 3503 ± 14 Ma and 560 ± 550 Ma (MSWD=1.0; Fig.2b). The lower intercept matches the timing of Pb-loss recorded in the alteration domains and the Raman radiation damage age. We argue that short distance Pb redistribution within the cores was controlled by U-Pb systematics recording both the crystallization age (upper intercept) and the timing of Pb redistribution (lower intercept). Some excess Pb was redistributed from the partially re-set (~3410 Ma) alteration domains into the cores, leading to mixed U-Pb systematics, observed in negative lower intercepts (Fig.2a). Only the latter excess Pb was leached upon CA. We argue that Pb from the alteration domains accumulated within distinct damage sites, accessible for CA leaching, and representing enhanced diffusion pathways that promoted microstructure-related open-system behavior during alteration. Such diffusion pathways may be recognized in radial fractures (Fig.1a), possibly formed as a result of volume swelling during radiation-induced point defect accumulation. We conclude that standardization of pre-analytical treatment is not recommended; we propose to investigate the structural state of unknown specimen by means of Raman spectroscopy, in combination with commonly applied CL imaging, in order to evaluate case-specific treatment requirements. During LA- ICP-MS analyses, co-measurement of non-formula and selected trace element contents is deemed highly valuable in detecting ablations for use in reliable U-Pb age determination.

[1] U. Schaltegger, et al., Chemical Geology, 402 (2015), p. 89-110. [2] The authors thank L. Rintoul (Raman), H. Cathey (EMPA), K. H. Moromizato (LA-ICP-MS), and other CARF staff for assistance in sample preparation and analytics. D. Wiemer acknowledges fruitful input during the U-Th-Pb LA-ICP-MS network workshop, Prague, 2015.

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Appendix 6 continued

Figure 1. a) CL image of part of an annealed (TA) zircon, showing the four identified domains; note dark CL holes (a) indicating porous alteration texture, and radial fractures in rim and core, terminating in the alteration domains (b); b) plot showing increase in non-formula Ca above 2nd percolation point;

Figure 2. Concordia diagrams of core analyses; a) pristine; (note: pristine Temora2 reductions shown as points, annealed Temora2 reductions shown as error ellipses); b) chemically abraded (TA+CA);

165

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Appendix 7

Appendix Table 3.A-I QUT CARF Axios major element detection limits (DL) Oxide avgerage DL (wt%) PbO 0.00297

HfO2 0.00301 ZnO 0.00066 CuO 0.00090 NiO 0.00098

Fe2O3 0.00169

Mn3O4 0.00165

Cr2O3 0.00172

V2O5 0.00191 BaO 0.00389

TiO2 0.00199 CaO 0.00213

K2O 0.00153

SO3 0.00194

ZrO2 0.00360

P2O5 0.00307 SrO 0.00510

SiO2 0.00559

Al2O3 0.00397 MgO 0.00433

Na2O 0.00340 Note: this is an average DL from 30+ major element

analyses (mainly SiO2-rich materials). The DL is dependent on bulk sample composition.

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Appendix 8

Appendix Table 3.A-II ICP-MS trace element detection limits, Central Analytical Research Facility, QUT Element Detection limit (ppb) 7 Li [ No Gas ] 0.02514416 45 Sc [ He ] 0.00279251 47 Ti [ He ] 0.13380244 51 V [ He ] 0.00616899 52 Cr [ He ] 0.08288774 59 Co [ He ] 0.00613051 60 Ni [ He ] 0.02213658 63 Cu [ He ] 0.00811341 66 Zn [ He ] 0.09736629 72 Ge [ He ] 0.00165372 85 Rb [ No Gas ] 0.00257178 88 Sr [ He ] 0.00570788 89 Y [ No Gas ] 0.00113012 90 Zr [ No Gas ] 0.00239880 93 Nb [ No Gas ] 0.00035689 95 Mo [ No Gas ] 0.00367485 107 Ag [ No Gas ] 0.00027563 111 Cd [ He ] 0.00071211 111 Cd [ No Gas ] 0.00078595 118 Sn [ He ] 0.00594457 118 Sn [ No Gas ] 0.00769095 121 Sb [ No Gas ] 0.00238309 133 Cs [ No Gas ] 0.00061880 137 Ba [ No Gas ] 0.02593383 138 Ba [ No Gas ] 0.01822295 139 La [ No Gas ] 0.00078232 140 Ce [ No Gas ] 0.00068934 141 Pr [ No Gas ] 0.00048180 146 Nd [ No Gas ] 0.00042084 147 Sm [ No Gas ] 0.00035641 153 Eu [ He ] 0.00032285 157 Gd [ He ] 0.00052717 159 Tb [ He ] 0.00026958 163 Dy [ He ] 0.00044778 165 Ho [ He ] 0.00028960 166 Er [ He ] 0.00031758 169 Tm [ He ] 0.00034509 172 Yb [ He ] 0.00029493 175 Lu [ He ] 0.00029250 178 Hf [ He ] 0.00005175 181 Ta [ He ] 0.00009243 182 W [ He ] 0.00023498 182 W [ No Gas ] 0.00024067 204 Pb [ No Gas ] 0.01552539 205 Tl [ No Gas ] 0.00093331 206 Pb [ He ] 0.00810196 206 Pb [ No Gas ] 0.00792231 207 Pb [ He ] 0.00803611 207 Pb [ No Gas ] 0.00758525 208 Pb [ He ] 0.00658801 208 Pb [ No Gas ] 0.00725809 232 Th [ No Gas ] 0.00059225 238 U [ No Gas ] 0.00027186

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Appendix 9

Appendix Table 3.A-III: Zirco n chemical data (LA-ICP -MS) Spo t Zr Hf Y P Fe Ca Ti La Ce Pr Sm Nd Eu Dy Lu # (ppm) Quartz-s ands to ne 335-1 1 428000 7840 2140 686 na 7 4.0 40.7 34.2 8.0 234.0 95.4 2 1120 113 95 800 0 3730000 6.8 11.1 7.2 3.2 28.0 1.0 3 450000 9520 704 480 na 5 3.7 50.4 25.6 6.0 69.8 45.9 4 630000 16290 28900 34400 6500 648 255.0 1874.0 2084.0 661.0 5140.0 684.0 5 453000 10620 6750 12010 1400 145 59.9 453.0 491.0 144.0 1077.0 196.4 6 434000 9710 4230 7340 1000 57 47.6 385.0 348.0 89.9 679.0 116.3 7 452000 9140 4940 7710 910 81 57.1 353.0 406.0 112.5 833.0 127.8 8 536000 14500 27300 40800 5530 404 353.0 1814.0 1896.0 551.0 4610.0 543.0 9 444000 10400 5470 8790 1510 110 58.0 418.0 409.0 112.2 897.0 151.1 10 462000 10930 2690 3960 650 40 23.9 224.0 188.0 51.7 427.0 95.2 11 497000 10190 5920 8750 1930 1850 56.4 471.0 466.0 128.2 1004.0 132.4 12 523000 13640 20290 31900 4180 275 159.0 1290.0 1380.0 410.0 3470.0 467.0 13 435000 10350 5980 9860 1490 114 75.7 591.0 599.0 163.3 1027.0 152.9 14 449000 10050 6790 11540 1590 100 80.1 584.0 550.0 141.9 1157.0 188.0 15 436000 9000 6180 10890 1820 111 63.5 509.0 557.0 145.5 976.0 146.4 16 449000 8110 6680 8820 1490 92 51.1 394.0 430.0 123.2 1014.0 175.3 17 455000 11040 7270 10890 2710 191 79.9 572.0 620.0 183.0 1220.0 185.5 18 429000 9560 3000 5080 990 35 26.2 243.0 228.0 64.7 465.0 99.4 19 430000 10080 3180 5040 700 67 48.2 263.0 279.0 73.7 510.0 90.2 20 409000 9110 4110 6750 1270 52 48.7 382.0 331.0 92.8 634.0 108.4 21 453000 12330 3540 7700 670 62 33.0 192.0 241.0 69.0 600.0 96.0 22 447000 7930 1160 255 60 5 0.0 22.3 2.5 2.4 102.4 74.9 24 417000 9940 4190 6640 930 110 91.8 369.0 460.0 114.4 660.0 114.8 25 437000 8010 1440 2830 260 20 18.5 148.0 137.0 34.8 228.0 35.0 26 445000 9520 709 1030 120 8 6.2 54.0 41.4 10.4 92.4 27.4 27 464000 10980 10110 15920 2010 173 168.4 888.0 914.0 243.0 1688.0 225.4 28 433000 11690 8820 12700 1810 186 87.2 475.0 595.0 166.0 1475.0 202.2 29 421000 10230 2950 5000 640 45 40.7 259.0 264.0 67.7 496.0 87.7 30 300000 6050 908 4830 1210 3450 1020.0 1480.0 665.0 36.5 132.0 31.1 31 449000 7860 276 104 na 6 0.0 10.4 1.0 0.9 22.0 12.7 32 444000 7810 716 837 30 14 4.6 38.7 34.9 8.7 92.8 25.3 33 443000 8260 342 312 60 8 0.8 19.1 8.6 2.9 35.4 14.6 34 435000 8690 288 510 na 7 119.0 183.0 92.0 4.4 27.6 13.8 35 428000 8460 873 1530 260 13 9.1 80.6 77.8 19.6 117.1 24.1 36 423000 7880 1255 1500 20 23 11.9 88.0 88.7 25.9 192.0 35.4 37 357000 9900 9600 15700 na 300 129.0 680.0 840.0 231.0 1690.0 213.0 38 437000 8310 4150 5920 810 81 36.3 321.0 310.0 92.8 708.0 91.3 39 440000 8620 343 132 130 5 0.4 12.2 3.8 1.4 32.6 16.2 40 418000 8590 5320 5800 1520 85 48.3 352.0 423.0 125.0 860.0 116.0 41 428000 8100 5820 9040 980 107 79.7 449.0 538.0 143.2 993.0 131.8 43 418000 9130 857 980 na 7 4.4 74.0 41.8 13.4 100.0 53.6 44 429000 7660 1770 1770 310 22 9.0 124.0 76.0 24.4 225.0 75.3 45 421000 7040 1422 502 0 12 2.6 32.3 22.5 8.3 149.0 68.8 46 421000 8810 1374 1660 180 16 12.0 107.0 98.3 26.6 191.0 59.1 47 438000 8130 627 586 310 11 1.5 25.9 18.8 6.4 72.8 26.0 48 440000 10020 5220 8070 1240 80 81.0 515.0 543.0 137.5 852.0 110.3 49 409000 7910 2740 2990 180 32 26.5 177.2 181.0 60.5 478.0 62.9 50 421000 9730 1230 1410 720 19 10.2 126.0 89.0 22.7 158.0 55.3 51 436000 7470 1450 1820 390 26 12.7 104.0 105.0 29.9 228.0 52.2 52 430000 7600 2220 3480 890 38 21.2 393.0 290.0 63.0 303.0 44.6 53 460000 9830 9150 15270 3680 156 113.3 846.0 696.0 210.3 1484.0 243.5 54 450000 9370 4290 7840 1300 53 55.5 484.0 408.0 108.7 732.0 100.7 55 433000 9800 3470 6360 1220 47 43.0 323.0 297.0 82.1 570.0 103.7 56 426000 9920 2980 3990 190 53 39.7 292.0 282.0 68.0 456.0 94.5 57 420000 7920 744 760 70 13 9.2 52.2 51.3 12.2 104.0 29.3 Quartz-jasper conglomerate 332-1 59 424000 7190 1215 1560 140 19 2.4 30.4 34.1 19.7 165.0 43.4 60 418000 8020 374 2980 420 990 3.9 20.3 10.9 4.8 40.3 15.3 61 405000 7060 2580 2550 380 24 3.1 46.4 77.0 47.4 380.0 73.3 62 474000 8340 17800 8500 1200 49 415.0 555.0 270.0 338.0 3160.0 191.0 63 455000 10960 6290 11080 1920 209 36.7 215.0 307.0 131.1 951.0 186.4 64 419000 8950 310 10 na 4 0.0 11.8 1.8 0.9 27.6 18.0 65 431000 7230 396 200 0 12 0.3 12.3 4.1 2.1 39.8 17.5 66 409000 6930 1360 2010 130 29 3.0 36.3 44.6 26.3 191.0 41.5 67 422000 7730 1310 1640 290 19 3.2 33.9 45.0 18.9 172.0 41.0 68 452000 8890 3230 5240 870 28 10.0 79.9 114.4 58.4 459.0 86.4 69 412000 9770 4160 6160 1440 76 8.9 107.9 145.8 78.5 608.0 110.8 70 426000 8180 2180 3050 550 29 4.7 49.8 83.0 42.1 329.0 54.2 71 416000 7720 2370 5170 510 46 16.7 87.7 95.0 43.0 345.0 67.8 72 436000 7740 571 190 50 8 0.1 14.7 3.3 1.4 47.7 35.0 To nalitic-tro ndhjem itic gneis s 310-4 1 446000 9840 1410 460 236 4 13.3 44.6 4.5 13.7 4.3 130.0 63.8 2 415000 11510 1260 327 2560 5 22.2 183.0 17.6 33.5 11.5 120.3 66.9

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Appendix Table 3.A-III: Co ntinued Spo t Zr Hf Y P Fe Ca Ti La Ce Pr Sm Nd Eu Dy Lu # (ppm) 3 426000 10320 1320 201 1120 6 6.7 47.1 4.2 16.7 4.6 126.0 64.3 4 403000 9030 1662 1080 1478 10 22.2 112.3 11.5 37.7 9.1 169.4 68.6 5 417000 10170 1620 231 1270 6 14.8 88.0 9.1 39.0 11.3 165.0 67.7 6 423000 9050 1206 201 1165 6 8.7 80.7 8.6 32.5 8.4 126.9 55.4 7 389000 11640 861 82 1062 5 14.4 135.7 12.9 46.3 10.9 110.0 41.9 8 399000 9110 4370 11100 3280 21 78.7 700.0 71.5 306.0 83.9 715.0 86.9 9 378000 9770 1634 2550 6100 8880 36.4 187.0 19.6 56.9 13.8 181.5 61.7 10 404000 8850 2140 1700 3320 10 45.1 394.0 21.2 57.8 17.2 251.0 70.9 11 400000 10610 1231 89 1653 17 94.5 344.0 43.4 91.1 27.3 182.4 41.3 12 439000 9620 960 124 293 3 2.7 26.7 1.3 7.8 2.7 86.0 47.3 13 427000 9360 1117 145 736 5 4.4 36.5 2.9 8.4 3.1 101.4 55.7 14 430000 11260 1486 131 3040 18 26.6 165.0 18.5 72.9 18.2 209.2 55.4 15 398000 8930 1737 430 1241 6 26.5 115.8 14.8 48.5 12.8 190.5 68.8 73 416000 13800 1219 230 6290 4 11.7 29.4 18.0 12.2 97.3 80.3 74 408000 8220 1630 810 2950 4 9.6 97.2 54.3 8.3 172.0 69.4 75 389000 8380 1574 420 2940 253 12.4 59.3 42.5 7.7 166.3 63.1 76 398000 8710 1286 3230 10800 8 6.1 57.7 34.7 6.4 120.4 57.5 77 391000 9870 1173 282 1820 7 16.6 79.6 59.1 9.6 124.9 51.4 78 388000 8760 1800 243 1840 14 32.3 181.0 133.0 20.5 214.0 61.4 79 417000 9060 1236 75 440 8 30.8 128.0 110.0 16.2 170.0 40.5 80 410000 8930 1554 1140 3490 3 3.5 31.9 13.8 3.2 141.3 72.4 81 391000 8830 1308 1980 7600 4 1.1 22.8 4.0 2.1 116.9 56.1 82 396000 11850 1076 52 4110 1 64.0 113.8 76.4 8.3 114.2 50.7 83 400000 8380 1555 203 4380 7 11.1 91.1 40.8 8.7 179.9 54.9 84 398000 8090 1217 240 1080 3 0.1 18.3 2.0 1.8 104.0 62.0 85 429000 10820 667 460 2350 11 3.4 16.3 8.8 2.1 58.2 37.9 86 439000 10590 1183 116 1970 6 7.8 45.8 26.2 6.0 107.4 60.9 87 409000 8390 1977 2330 7370 17 18.3 144.3 89.6 16.0 253.0 65.6 88 404000 7940 1166 7600 25900 5 3.5 29.4 14.4 3.1 118.4 49.7 89 392000 10020 948 3480 12100 740 5.0 33.7 15.8 2.8 89.1 46.2 90 403000 7910 1554 183 1640 8 18.1 137.2 85.2 13.5 173.8 62.0 91 384000 8480 1655 2210 8960 12 47.7 259.0 151.7 20.9 235.0 49.1 92 405000 9980 1923 332 3590 3 4.2 36.0 15.0 4.2 188.7 88.5 93 396000 7850 1271 546 2170 4 2.0 21.9 5.3 2.3 113.3 59.0 94 382000 8020 1292 75 1450 4 3.3 28.1 3.2 2.2 121.4 57.6 95 421000 8900 1380 176 1460 4 2.1 27.6 8.7 2.5 119.4 70.8 96 391000 8460 1375 160 4600 5600 88.7 365.0 197.0 19.3 178.5 55.8 97 405000 8940 1431 1070 7630 6 73.0 263.0 99.0 11.0 164.0 52.8 98 380000 7550 1411 1550 5800 11 24.4 170.2 116.2 19.0 186.0 43.7 99 444000 10410 1334 450 2640 20 4.9 40.5 16.9 4.0 120.9 65.6 100 398000 10630 1186 62 2240 2 30.0 122.0 104.0 17.4 161.0 37.6 101 388000 9460 2126 230 4760 112 16.6 121.0 77.0 10.6 231.0 82.0 102 354000 9570 1407 22200 79600 19500 21.6 72.9 57.6 11.3 156.0 52.3 Granitic gneiss 308-1 103 386000 22300 803 383 1570 1 12.3 16.7 28.8 3.4 63.5 104.6 104 366000 22680 536 197 1170 0 1.6 7.7 4.0 0.4 35.7 74.8 105 384000 25700 1542 423 5020 2 0.1 2.9 2.3 0.3 125.1 138.2 106 411000 23000 2000 650 6400 0 0.1 3.3 0.8 0.1 162.3 173.3 107 428000 8480 3210 840 870 41 22.0 176.0 145.0 17.6 385.0 128.0 108 479000 12660 1514 227 80 13 0.7 18.7 4.1 0.2 147.7 74.2 109 367000 22100 2070 716 1720 5 60.0 134.0 233.0 26.9 227.0 135.9 110 372000 13010 1559 602 4330 2 1.1 6.8 2.5 0.4 111.2 108.6 Dio ritic gneis s 370-1 1 191 13 523 95 6730 214500 268.0 896.0 141.2 179.1 61.1 107.6 5.0 2 447000 11640 807 114 8 3 0.0 12.9 0.0 2.2 1.0 57.5 58.3 3 454000 9760 891 168 10 3 0.4 21.5 0.6 8.0 2.8 84.7 47.0 4 440000 8060 1688 203 11 8 0.2 12.7 0.6 13.7 6.1 178.4 66.7 5 433000 8130 871 208 42 8 4.5 65.9 4.1 17.2 5.8 80.3 42.2 6 437000 9530 958 146 23 5 0.9 22.2 1.3 10.4 3.7 89.3 46.6 7 440000 8860 939 136 10 7 0.1 9.7 0.4 7.4 2.7 91.2 44.6 9 428800 8530 386 174 4 5 0.0 3.7 0.0 1.6 0.8 33.7 23.9 10 402000 10470 972 126 295 3 1.9 24.9 1.3 10.7 4.3 86.9 51.5 11 424000 9050 1172 187 219 9 1.5 23.6 1.1 7.6 3.0 90.7 68.5 12 428000 8580 1411 168 21 6 0.0 12.0 0.4 10.5 4.5 140.3 64.4 13 431000 9060 668 200 43 6 7.0 87.1 7.7 30.0 12.5 62.7 35.4 14 424000 9430 591 131 11 3 0.0 8.8 0.0 1.8 1.1 45.4 42.4 15 435000 8910 855 132 8 4 0.0 8.5 0.3 6.0 2.5 84.8 39.8 16 405000 11120 975 117 29 2 0.2 13.2 0.2 6.8 2.5 87.3 51.7 17 412000 8280 733 219 27 6 15.3 51.8 11.7 33.0 10.4 66.9 37.2 18 406000 8390 553 205 12 6 0.8 18.3 0.4 4.4 1.6 49.2 32.2 19 400000 11260 470 130 36 7 13.5 105.0 9.9 29.1 9.5 39.3 23.7 20 412000 8920 764 109 18 6 5.0 96.6 3.8 21.1 7.4 79.5 33.0

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Appendix Table 3.A-III: Co ntinued Spo t Zr Hf Y P Fe Ca Ti La Ce Pr Sm Nd Eu Dy Lu # (ppm) Grano dio ritic gneis s 309-2 1 437000 10670 397 86 125 24 19.6 188.0 11.1 21.7 6.1 28.3 38.3 2 451000 8110 5080 301 44 9 0.7 15.6 0.7 19.5 6.0 503.0 228.4 3 454000 8280 4340 253 24 7 1.5 29.2 1.4 20.3 5.8 425.0 194.4 4 426000 14210 2500 837 1230 25 8.9 66.7 7.4 29.3 7.2 146.7 186.0 5 154 9 683 87 4720 228800 218.0 887.0 152.2 243.8 129.8 146.1 7.3 6 310 27 1137 155 6380 222800 592.0 1966.0 296.0 403.0 108.9 230.9 15.0 7 460000 8340 1320 222 71 15 12.1 87.0 8.1 27.3 15.1 121.0 60.2 8 443000 7950 742 164 12 5 1.1 18.7 1.5 7.3 3.2 70.2 38.3 9 440000 8760 1175 182 680 21 5.2 49.7 2.8 13.7 5.3 114.1 55.0 10 448000 8700 1030 186 16 6 2.4 35.2 2.1 12.4 4.3 96.0 48.0 11 432000 8750 983 133 76 56 19.3 106.0 11.5 25.9 22.0 85.7 40.5 12 440000 7720 963 197 46 45 26.4 271.0 18.9 62.0 13.7 107.1 42.0 13 444000 8280 1305 2510 42 13 11.4 57.4 4.8 12.2 3.4 122.4 55.0 14 444000 9310 3730 208 14 3 0.0 6.6 0.2 11.2 2.8 349.5 170.4 15 418100 8130 816 199 38 8 2.0 20.7 1.6 10.7 3.9 82.8 38.3

170