Neoarchaean metamorphic evolution of the Yilgarn Craton: a record of subduction, accretion, extension and lithospheric delamination

Journal Pre-proofs Neoarchaean metamorphic evolution of the Yilgarn Craton: a record of subduction, accretion, extension and lithospheric delamination Ben Goscombe, David A. Foster, Richard Blewett, Karol Czarnota, Ben Wade, Bruce Groenewald, David Gray PII: DOI: Reference:

S0301-9268(18)30558-8 https://doi.org/10.1016/j.precamres.2019.105441 PRECAM 105441

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Precambrian Research

Received Date: Revised Date: Accepted Date:

29 October 2018 24 February 2019 30 August 2019

Please cite this article as: B. Goscombe, D.A. Foster, R. Blewett, K. Czarnota, B. Wade, B. Groenewald, D. Gray, Neoarchaean metamorphic evolution of the Yilgarn Craton: a record of subduction, accretion, extension and lithospheric delamination, Precambrian Research (2019), doi: https://doi.org/10.1016/j.precamres.2019.105441

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Neoarchaean metamorphic evolution of the Yilgarn Craton: a record of subduction, accretion, extension and lithospheric delamination. Ben Goscombe1,2*, David A. Foster2, Richard Blewett3, Karol Czarnota3, Ben Wade4, Bruce Groenewald5, David Gray6. 1Integrated

Terrane Analysis Research (ITAR), 18 Cambridge Rd, Aldgate, 5154, SA, Australia.

2Department

of Geological Sciences, University of Florida, Gainesville, Florida, 32611, USA.

3Geoscience

Australia, GPO Box 378, Canberra, 2601, ACT, Australia.

4Adelaide 566

Microscopy, University of Adelaide, Adelaide, 5005, SA, Australia.

Chessell Drive, Duncraig, WA, Australia.

6School

of Earth Sciences, University of Tasmania, Hobart, Tasmania. Australia.

*Corresponding author. Email: [email protected]. http://www.terraneanalysis.com.au Supporting datasets (Appendices 1 to 8) are contained in electronic appendices. ABSTRACT

The greater part of the metamorphic record of the Yilgarn Craton formed during a protracted middle Neoarchaean orogenic cycle, spanning from about 2750 to 2620 Ma. At least six distinct metamorphic events are defined by parageneses in different tectono-stratigraphic settings. This review characterizes the middle Neoarchaean metamorphic response of the entire craton using large databases. Rock descriptions from ~60,000 sites in government databases and ~1,040 PT determinations from all sources, are used to construct metamorphic map patterns for each of these metamorphic events. P-T paths and timing of each metamorphic event are characterized using petrology and the well-constrained PT determinations from this study (n=287), combined with metamorphic age determinations from all sources (n=114). The spatial patterns and metamorphic conditions of each event, provides critical constraints on tectonic settings and changes in tectonics over 130 million years. The record of metamorphism in conjunction with stratigraphic, kinematic and magmatic constraints shows that the Neoarchaean craton evolved from an accretionary crustal growth phase (Ma and M1), to thermal reworking of the crust during massive influx of granitic intrusions (M2), and lithospheric extension after the termination of subduction (M3). These events were followed by lithospheric delamination, resulting in a craton-wide, diffusion-delayed thermal pulse that coincided with renewed contraction (M4). M1 metamorphism at 2748±19, 2727±8 and 2706±10 Ma experienced high-P/moderate-T hairpin, clockwise P-T paths during burial of magmatic arc margins during terrane accretion events. M2 regional-contact type metamorphism at 2671±6 Ma tracked low-P/moderate-T clockwise P-T paths in back-arc settings during voluminous felsic magmatism related to shallow subduction. M3 metamorphism at 2656±5 Ma experienced lowP/moderate-T anticlockwise P-T paths in post-volcanic rift basins formed during lithospheric extension. M4 metamorphic peaks at 2644±4 and 2629±7 Ma experienced low-P/high-T clockwise P-T paths during regional-scale thermal pulses resulting from lower-crust and mantle lithosphere delamination at 2665-2668 Ma.

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

TABLE OF CONTENTS

Abstract ...................................................................................................................................................................1 Table of Contents .................................................................................................................................................2 Key Words ..............................................................................................................................................................3 Abbreviations........................................................................................................................................................3 1. Introduction ......................................................................................................................................................4 2. Crustal Architecture .......................................................................................................................................7 3. Stratigraphy and Deformation History ....................................................................................................8 3.1. Early Extension and Volcanism .............................................................................................................................8 3.2. Contraction and Termination of Volcanism.....................................................................................................9 3.3. Late Rift Basins .............................................................................................................................................................9 3.4. Reactivation, Au-Mineralization and Cratonization ..................................................................................11 4. Granitoid Magmatism..................................................................................................................................12 5. Neoarchaean Metamorphic History........................................................................................................13 5.1. Ma Magmatic Arc Metamorphism .....................................................................................................................14 5.2. M1 High-Pressure Metamorphism ....................................................................................................................14 5.3. M2 Regional-Contact Type Metamorphism...................................................................................................15 5.4. M3 Extensional Metamorphism..........................................................................................................................15 5.5. M4 High-Grade Metamorphism ..........................................................................................................................16 6. Metamorphic Petrology ..............................................................................................................................17 6.1. Ma Magmatic Arc Mafic Granulite .....................................................................................................................17 6.2. M1 High-Pressure Amphibolite ..........................................................................................................................17 6.3. M2 Metapelite.............................................................................................................................................................18 6.4. M3 Metapelite.............................................................................................................................................................19 6.4.1. Belches Formation, Kurnalpi Terrane....................................................................................................19 6.4.2. Ockerburry Shear Zone, Kalgoorlie Terrane .......................................................................................20 6.4.3. Youanmi Terrane.............................................................................................................................................20 6.5. M4 Metapelite.............................................................................................................................................................21 6.5.1. Youanmi Terrane.............................................................................................................................................21 6.5.2. Southwest Terrane .........................................................................................................................................22 6.5.3. Narryer Terrane...............................................................................................................................................23 7. Metamorphic Methodologies ....................................................................................................................24 7.1. Mineral Chemistry ....................................................................................................................................................24 7.2. PT Determinations by THERMOCALC..............................................................................................................24 7.3. PT Determinations by Conventional Geothermobarometry..................................................................25 7.4. Estimation of Prograde Conditions...................................................................................................................26 7.5. Semi-Quantitative P-T Paths................................................................................................................................26 8. Integrated Metamorphic Results.............................................................................................................27 8.1. Ma Metamorphism (~2715-2750 Ma) ............................................................................................................27 8.2. M1 Metamorphism (2700-2750 Ma)................................................................................................................28 8.3. M2 Metamorphism (2665-2685 Ma)................................................................................................................29 8.4. M3 Metamorphism (2650-2665 Ma)................................................................................................................32 8.4.1. Kalgoorlie and Kurnalpi Terranes............................................................................................................32

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8.4.2. Youanmi Terrane.............................................................................................................................................34 8.5. M4 High-Grade Metamorphism (2615-2650 Ma).......................................................................................35 8.5.1. Youanmi Terrane.............................................................................................................................................36 8.5.2. Southwest Terrane .........................................................................................................................................38 8.5.3. Narryer Terrane...............................................................................................................................................39 8.6. M4 Alteration and Mineralization (2610-2655 Ma) ..................................................................................40 9. Geochronology of Metamorphic Events ................................................................................................41 9.1. Analytical Procedures .............................................................................................................................................42 9.1.1. Lu-Hf and Sm-Nd Garnet Chronology.....................................................................................................42 9.1.2. U-Pb Monazite Chronology .........................................................................................................................42 9.1.3. U-Pb Titanite Chronology ............................................................................................................................43 9.1.4. U-Pb Zircon Chronology ...............................................................................................................................43 9.2. Relict (pre-M1) Metamorphism..........................................................................................................................43 9.3. M1 Metamorphism...................................................................................................................................................44 9.4. M2 Metamorphism...................................................................................................................................................45 9.5. M3 Metamorphism...................................................................................................................................................45 9.6. M4 Metamorphism...................................................................................................................................................46 9.7. Alteration and Mineralization .............................................................................................................................47 9.8. Isotopic Resetting and Cooling Ages.................................................................................................................47 10. Discussion and Conclusions: Neoarchaean Accretionary Orogenic Cycle...............................47 10.1. Magmatic Arcs [Ma] ..............................................................................................................................................48 10.2. Arc Accretion [M1] ................................................................................................................................................49 10.3. Subduction and Granite Bloom [M2] .............................................................................................................49 10.4. Lithospheric Extension [M3].............................................................................................................................50 10.5. Delamination [M4].................................................................................................................................................53 10.6. Comparison with Phanerozoic Accretionary Orogens...........................................................................55 Acknowledgements ..........................................................................................................................................57

KEY WORDS

P-T calculations; metamorphic petrology; metamorphic evolutions; lithospheric extension; lithospheric delamination; accretionary orogens ABBREVIATIONS

EYC

East Yilgarn craton.

WYC

West Yilgarn craton.

P

Pressure in kilobars (kbar).

T

Temperature in degrees Celsius (ºC).

G

Average thermal gradient or temperature/depth ratio (ºC/km).

PT

Pressure and temperature determination or locus.

P-T

Pressure-temperature space, as in pressure-temperature evolution path.

Z

Depth in crust (km), calculated from pressure (Z ≈ 3.5xP) assuming average rock density of 2.8 g/cm3.

Minerals

Mineral abbreviations after Kretz (1983). Additional minerals listed in Table (3).

M…

Metamorphic event.

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Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

D…

Deformation event.

Ma

Millions of years before present. All age ranges are in numerical order, not chronologic order.

GA

Geoscience Australia.

GSWA

Geological Survey of Western Australia.

ITAR

Integrated Terrane Analysis Research.

1. INTRODUCTION

The Yilgarn craton is a large (950 by 1100 km), contiguous lithospheric element (Figure 1) that contains a near continuous record of Neoarchaean crustal growth, recycling and formation of worldclass gold and nickel fields. Not unsurprisingly, the craton has attracted significant attention for over more than a century (e.g. Cassidy et al., 1998; Bierlein et al., 2006; Czarnota et al., 2010b). These studies have resulted in large, rich and comprehensive geological, stratigraphic, magmatic, structural, geophysical and isotopic datasets (as summarized in Blewett et al., 2008), and a mature understanding of the tectonic history (e.g. Champion et al., 2006; Czarnota et al., 2008, 2010a; Blewett et al., 2010; Korsch et al., 2013; Zibra et al., 2017). The Yilgarn craton experienced metamorphic conditions ranging from sub-greenschist to granulite facies (e.g. Binns et al., 1976) and contains metamorphic parageneses spanning almost the whole Neoarchaean (e.g. Nemchin et al., 1994; Rennie, 1998; Nemchin and Pidgeon, 1999; Fletcher and McNaughton, 2001; Mueller et al., 2004; Sircombe et al., 2007; Van Kranendonk, 2007). On this foundation, the Yilgarn craton is ideal for generating a record of Neoarchaean metamorphism at craton-scale, and test models for late Archaean metamorphism and tectonics in general. The Yilgarn craton spans the crucial Earth history transition from vertical or plume tectonics with symmetric advection, to horizontal or plate tectonics with asymmetric subduction. Metamorphic expressions of these two tectonic paradigms are fundamentally different, and therefore, the history of metamorphism is among the best ways to recognise and document this transition (Brown 2006; Condie, 2008). Metamorphic processes such as heat production, heat diffusion and isostatic response are continuous variables that vary smoothly across large scales and time frames; in response to more episodic parameters like stress switches, kinematic partitioning and erosion for example (e.g. England and Thompson, 1984). As a result, metamorphic rocks potentially preserve a near continuous record of the thermal and barometric evolution of orogenic systems that can be back-engineered from samples representing the different parts of the system (Figure 2). Furthermore, metamorphic rocks are the only reliable data source on crustal depth and particle paths giving insights into lithospheric thickening and thinning events. Spatial variation in metamorphic response documented by metamorphic maps and field gradients, and temporal variation documented by P-T-t paths, constrain relative vertical movement between different tectono-metamorphic domains and events. Consequently, metamorphic datasets are crucial to reconstructing the evolution of orogenic systems, and defining probable tectonic settings, inturn offering insights into fluid flow and mineralization (e.g. Inger and Harris, 1992; Spear, 1993; Vannay and Hodges, 1996; Brown and Solar, 1999; Hodges et al., 2001; Jamieson et al., 2002; Goscombe et al., 2017a, 2018; Blewett et al., 2010). The rationale behind this study is the integration of craton-wide spatial-temporal metamorphic datasets with pre-existing geological datasets for deformation, stratigraphy and magmatism (Goscombe et al., 2007, 2009; Czarnota et al., 2010a). The Yilgarn craton is characterized by [1] heterogeneous strain distribution, which is typical of well-endowed gold terranes (e.g. Bierlein et al., 2002, 2006), and [2] steep metamorphic gradients that are known to be important drivers of fluid formation, fluid flow and gold mineralization (e.g. Hall, 1997; Sheldon et al., 2007, 2008). Consequently, there is a prima facie need to link metamorphic and structural patterns, and by correlating the timing of metamorphic crystallization events with deformation and magmatism, robust models for crustal evolution can be developed (e.g. Czarnota et al., 2010a; Korsch et al.,

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2013). A sustained research effort over the last 20 years, including large government and industry sponsored programs (i.e. Y1, Y3 and Y4), have established robust frameworks for the structural evolution (e.g. Blewett et al., 2004a, 2008, 2010; Blewett and Czarnota, 2005, 2007b,c; Korsch et al., 2013), extrusive stratigraphy (e.g. Groenewald et al., 2006; Kositicin et al., 2008), clastic stratigraphy (e.g. Krapez et al., 2000; Squire, 2006, 2007), magmatic and geochemical evolution of the crust (e.g. Cassidy et al., 2002, 2006; Champion and Sheraton, 1997) and mineralization systems (e.g. Dugdale and Hagemann, 2001; Salier et al., 2005; Miller and Rasmussen, 2006: Walshe et al., 2008a,b). Similarly, the need for large-scale synthesis of geological data layers with spatial patterns of metamorphic response has been recognised as crucial in other Neoarchaean gold provinces (Valli et al., 2004; Thompson, 2005; Bleeker and Hall, 2007; Beakhouse, 2007; Berman, 2010). For the Yilgarn craton there are relatively few detailed metamorphic studies, most of which are widely scattered and undertaken in the 1980’s and 1990’s (e.g. Purvis, 1978; Gole and Klein, 1981; Archibald et al., 1978, 1981; Bickle and Archibald, 1984; Spray, 1985; Williams and Currie, 1993; Bloem et al., 1994; Dalstra, 1995; Ridley et al., 1997; Mikucki, 1997; Witt, 1998; Dalstra et al., 1998, 1999; Vielreicher et al., 2002). This legacy metamorphic data was generated by a large number of methods that are not directly comparable, and pressures were mostly assumed to be ~4.0 kbar (e.g. Binns et al., 1976; Ahmat, 1986; Wilkins, 1997), and thus insufficiently robust to form a relational database (Goscombe et al., 2007, 2009). Similarly, until recently the timing of Neoarchaean metamorphic events were estimated by correlation with magmatism (e.g. Kinny et al., 1988; McMillan, 1996; Dalstra et al., 1998; Bagas, 1999; Wilde, 2001; Blewett et al., 2004b; Van Kranendonk and Ivanic, 2009), and direct dating of metamorphic minerals was largely restricted to the Southwest and Narryer terranes (e.g. Pidgeon et al., 1990; Nemchin et al., 1994; Rennie, 1998; Campbell et al., 1998; Yumin et al., 1998; Nemchin and Pidgeon, 1999; Muhling et al., 2008; Iizuka et al., 2008; Rasmussen et al., 2010). The number of distinct Neoarchaean metamorphic events began to emerge with direct dating of metamorphic minerals in the lower-grade Youanmi, Kalgoorlie and Kurnalpi terranes (e.g. Wang et al., 1998; Schiotte and Campbell, 1996; Mueller and McNaughton, 2000; Fletcher and McNaughton, 2001; Mueller et al., 2004; Sircombe et al., 2007; Thebaud and Miller, 2009; Goscombe et al., 2007, 2009). Until relatively recently, tectonic models for the Yilgarn suffered from an absence of craton-wide metamorphic datasets and a general under-appreciation that thermal gradients, metamorphic conditions and P-T paths evolved through multiple metamorphic events in the Neoarchaean (Figure 2 and 3). This is in contrast to the more detailed analysis of other Neoarchaean terranes in South Africa (e.g. Diener et al., 2005; Dziggel et al., 2006) and Canada (e.g. Valli et al., 2004; Thompson, 2005; Berman, 2010). In the absence of a comprehensive analysis of the metamorphic history, the synthesis of other geological datasets listed above lead to a subduction-accretion tectonic model grounded in the plate-tectonic paradigm (e.g. Belt and Standing, 2007; Mary Gee and Swager, 2007; Blewett et al., 2010; Czarnota et al., 2008, 2010a; Korsch et al., 2013), with similarities to Phanerozoic accretionary orogens such as the Tasmanides being proposed (e.g. Foster and Gray, 2000; Foster et al., 2005; Gray and Foster, 2004; Gray et al., 2002). The subduction-accretion model has largely replaced two previous models for Yilgarn evolution: (1) diapirism involving partial convective overturn (e.g. Bloem et al., 1997; Dalstra et al., 1998; Rey et al., 2003; Van Kranendonk, 2007; Van Kranendonk and Ivanic, 2010), and (2) collisional tectonics involving crustal overthickening (e.g. Groves and Phillips, 1987; Myers, 1992, 1995; Qiu and Groves, 1999). The overturn and thickening models presuppose simple metamorphic histories with one cycle of prograde metamorphism (Binns et al., 1976; Swager et al., 1990; Wilkins, 1997; Dalstra, et al., 1999), and were thus unable to address a long-lived (~130 m.y.) and complex metamorphic history (Figure 3). At the same time that these models were being developed, a number of workers recognized that metamorphic response was not this simple and involved multiple events with contrasting structural modes (Bickle and Archibald, 1984; Ridley, 1992, 1993; Ridley et al., 1997; Williams and Whitaker, 1993; Williams and Currie, 1993; Witt, 1998; Mikucki and Roberts, 2003).

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The diapirism and the collisional models cannot account for a protracted and complex metamorphic cycle with multiple metamorphic peaks, different thermal regimes and contrasting P-T evolutions (Figure 3). For example, diapiric models are incompatible with the low variation in pressures (P ~0.8-1.6 kbar) recorded between margins and cores of granite-gneiss domes (Dalstra, et al., 1999). Crustal-scale extensional shear zones, core complexes, extensional rift basins, anticlockwise P-T paths and widespread extensional kinematics are consistent with lithospheric extension and incompatible with both diapiric and collisional models (Blewett and Czarnota, 2007c; Goscombe et al., 2007, 2009; Czarnota et al., 2010a; Blewett et al., 2010; Korsch et al., 2013). Deeply buried greenstone parageneses are rare, not associated with the margins of granite domes, pre-date the highCa granite bloom by ~16-37 m.y. and development of granite-gneiss domes by ~41-62 m.y., and thus implausible candidates for sinking keels accreted to the margins of diapirs (e.g. Cassidy et al., 2002, 2006; Champion and Sheraton, 1997; Goscombe et al., 2007, 2009; Blewett et al., 2010). Collisional tectonic models are incompatible with the highly partitioned strain pattern, multiple shortening and vergence directions, far-field stress switches, extensional kinematics and rift basins, and absence of large-scale thrust sheets and nappes (e.g. Blewett et al., 2004a, 2008, 2010; Blewett and Czarnota, 2005, 2007b,c; Czarnota et al., 2010a). Deep seismic imaging shows a craton-wide pattern of asymmetric structural grain and low-angle crustal-scale shear zones with extensional offsets, incompatible with symmetric diapiric models and collisional tectonics (Goleby et al., 1993, 2002a,b; Blewett et al., 2002a,b; Drummond et al., 2000; Henson and Blewett, 2006; Blewett et al., 2010; Korsch et al., 2013). The widespread low-P/high-T regional-contact type metamorphism, clockwise and anticlockwise P-T paths showing only minor burial, and absence of widespread deep burial and Barrovian-series zonation are all incompatible with collisional tectonics and crustal overthickening (Goscombe et al., 2007, 2009). The aim of this review is to characterize the metamorphic evolution throughout the Yilgarn craton during the middle Neoarchaean orogenic cycle. The approach taken has been to develop craton-wide relational datasets with unified systematic methods as described in Goscombe et al. (2017a, 2018). This allows testing for internal consistency and increased accuracy of pooled results, as well as documenting spatial and temporal patterns of metamorphic response. The strength of this approach is the use of sample populations to characterize metamorphism in different tectono-metamorphic zones and during different events. In order to achieve this, quantitative methods have been kept simple to generate a sufficiently large metamorphic dataset. An internally consistent database of robust peak metamorphic conditions (n~287) was generated by average-PT calculations in THERMOCALC v3.25 (Powell et al., 1998), using typical metapelite and meta-tholeiite assemblages. This data forms the basis of this project and was culled from a larger database generated using a range of geothermobarometry methods of variable confidence, sourced from this study (n~617) and literature (n~422). Similarly, a large dataset of semi-quantitative P-T paths covering all metamorphic events, have been generated using prograde, peak and post-peak metamorphic PT determinations in combination with evolving mineral parageneses within generic petrogenetic grids. Qualitative data such as metamorphic facies (n~60,000), was generated by reinterpretation of field and petrology descriptions in government databases and literature, combined with PT determinations and metamorphic mapping from this study and literature, to generate craton-wide metamorphic maps for each metamorphic event (e.g. Figures 5, 7, 8 and 9). New metamorphic mineral dates (n~52) were generated using a range of different rock types, minerals and isotopic systems: Pb-Pb titanite, U-Pb monazite and zircon, and Sm-Nd and Lu-Hf garnet. These results are combined with metamorphic dates from literature (n~62), to constrain each of the metamorphic events defined on the basis of petrology and tectono-stratigraphic relationships. These new metamorphic datasets are understood in the context of the structural framework, kinematic history and stratigraphy established in the literature and summarized by Blewett and Czarnota (2007b), Czarnota et al. (2010a), Blewett et al. (2008, 2010) and Korsch et al. (2013). The result is an integrated model for the middle Neoarchaean tectonic setting, and the metamorphic response as the orogenic system evolved within the plate

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tectonic paradigm from subduction-accretion to extension, delamination and cratonization (e.g. Czarnota et al., 2010a; Korsch et al., 2013). 2. CRUSTAL ARCHITECTURE

The Yilgarn Craton is divided into distinct tectono-metamorphic terranes that partitioned different deformation intensity, and metamorphic, stratigraphic and magmatic histories. From west to east, these are the Narryer, Southwest, Youanmi, Kalgoorlie, Kurnalpi and Burtville terranes (Figure 4). At largest scale, the long-lived, east-dipping, listric Ida Fault zone is a fundamental boundary separating hinterland West Yilgarn (WYC) crust, with Palaeoarchaean-Mesoarchaean precursors, from relatively juvenile, mostly Neoarchaean East Yilgarn (EYC) crust (e.g. Cassidy et al., 2006). The Ida Fault zone has an early extensional history during rifting at the margin of the Youanmi Terrane, and subsequent formation of back-arcs, magmatic arcs and rifted crustal slivers in the EYC (Figure 1). Ida Fault zone may have originated as the site of an early Neoarchaean subduction system on the eastern margin of the Youanmi Terrane (Korsch et al., 2013). Sm-Nd model age map patterns indicate that rifting and juvenile crust formation was focused in the Kalgoorlie and Kurnalpi terranes, with minor rifting within the Youanmi Terrane focused in the Murchison and Southern Cross domains (Figure 5; Cassidy and Champion, 2004; Champion et al., 2006; Champion and Cassidy, 2007). The Burtville Terrane is a rifted fragment from the Youanmi Terrane, sharing similar Mesoarchaean and early Neoarchaean stratigraphy (Figure 4; Pawley et al., 2008; Korsch et al., 2013). Further east across the Yamarna shear zone, the Yamarna Terrane has juvenile Sm-Nd model age signatures and back-arc setting of equivalent age to the Kalgoorlie and Kurnalpi terranes (Cassidy and Champion, 2004; Champion et al., 2006; Champion and Cassidy, 2007; Pawley et al., 2008). The Kalgoorlie and Kurnalpi terranes share roughly similar Neoarchaean stratigraphy and both incorporate Mesoarchaean and early Neoarchaean stratigraphy in their respective basement rocks (e.g. Barley et al., 2002, 2003). A similar arrangement of rifted cratonic margins blocks with Mesoarchaean precursors has been documented in the Slave Province (Ketchum et al., 2004; Bleeker and Hall, 2007). In contrast to the EYC, the WYC contains Palaeoarchaean precursors, is dominated by Mesoarchaean and early Neoarchaean stratigraphy, and is devoid of the middle Neoarchaean stratigraphy that dominates the EYC (Figure 4; e.g. Barley et al., 2002, 2003; Cassidy and Champion, 2004; Champion and Cassidy, 2007; Groenewald et al., 2006; Kositicin et al., 2008; Van Kranendonk and Ivanic, 2009; Wyche et al., 2013). The Kurnalpi, Kalgoorlie and Youanmi terranes share typical granite-gneiss dome and greenstone synformal basin architecture and regional-contact type metamorphism associated with granitoid influx. Whereas, the Southwest and Narryer terranes are gneissic terranes with the early granite-greenstone architecture largely obliterated by late-stage high-grade metamorphism. The Youanmi and Southwest terranes are interpreted to be contiguous, without a major crustal structure between them; the boundary marked by a transitional increase in metamorphic grade. The Narryer Terrane is a distinct Palaeoarchaean terrane with little affinity to the rest of the Yilgarn and was amalgamated with the Youanmi Terrane along the steep west-dipping crustal-scale Yalgar thrust (Korsch et al., 2013). Rifting events effecting the Palaeo- and Mesoarchaean core of the Yilgarn established a crustal architecture that was reactivated throughout the Neoarchaean and partitioned deformation and metamorphism. The Kurnalpi, Kalgoorlie and eastern Youanmi terranes are characterized by a pattern of inter-connected laterally extensive crustal-scale, curviplanar shear zone systems up to 800 km long and 0.1-5.0 km wide, some of which penetrate the lithospheric mantle (Drummond et al., 2000; Chen et al., 2001a,b; Blewett et al., 2010; Korsch et al., 2013). Four crustal-scale shear zone systems; Youanmi, Ida, Ockerburry and Hootanui divide the central Yilgarn into N-S elongate terranes/domains; Southern Cross, Kalgoorlie and Kurnalpi, that are further divided into elongate lenticular-shaped domains (Figure 1; Cassidy et al., 2006). Deep seismic profiling show these are east-dipping listric shear zones, steepest at current crustal level (40-70º), ~20º dips in the mid-crust and shallow in the lower-crust (Goleby et al., 1993, 2002a,b; Drummond et al., 2000; Korsch et al.,

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2013). This large-scale crustal architecture is primarily the result of early lithospheric extension of the Youanmi Terrane during Kalgoorlie-Kurnalpi crust formation and volcanism (Goleby et al., 2002a; Blewett and Czarnota, 2007a,c). The Ida and Hootanui shear zones were reactivated during accretion of the Youanmi, Kalgoorlie-Kurnalpi and Burtville terranes during a phase of late felsic volcanism and high-Ca granite magmatism, which established the N-trending grain of the EYC (Blewett and Czarnota, 2007c). The Ockerburry shear zone between Kalgoorlie and Kurnalpi terranes was reactivated by late-stage lithospheric extension that produced rift basins in the hanging wall (Blewett and Czarnota, 2007c). The Laverton, Hootanui, Ida, Bardoc, Celia, Kunanalling and Keith-Kilkenny shear zones also underwent extensional reactivation forming rift basins (Figure 1; Blewett and Czarnota, 2005). The late-stage lithospheric extension established a NNW-trending grain in the EYC, which was subsequently overprinted by N-trending dextral transpressive shear zones that reactivated the pre-existing N-trending grain (Blewett and Czarnota, 2007c). 3. STRATIGRAPHY AND DEFORMATION HISTORY

3.1. Early Extension and Volcanism

The Kalgoorlie, Kurnalpi and Burtville terranes experienced broadly similar Neoarchaean deposition, deformation and magmatic histories (Figure 2 and 4; e.g. Blewett and Czarnota, 2005, 2007a, b, c; Blewett et al., 2010; Czarnota et al., 2010a). Middle Neoarchaean tectono-magmatic events can be broadly correlated with the latest events in the Youanmi Terrane, which overprint the older Mesoarchaean and early Neoarchaean tectono-stratigraphic history (Korsch et al., 2013). The Southwest and Youanmi terranes preserve evidence for two early cycles of extension with greenstone sequences and associated TTG magmatism. Middle Mesoarchaean ~2900-2960 Ma and late Mesoarchaean to early Neoarchaean ~2735-2820 Ma events (Figure 4; e.g. Groenewald et al., 2006; Kositicin et al., 2008; Van Kranendonk and Ivanic, 2009; Zibra et al., 2017) are probably related to subduction at the eastern margin of the Youanmi Terrane (Figure 4; Korsch et al., 2013). Early extension and volcanism in the Youanmi Terrane was terminated by contraction and an unconformity at ~2730 Ma (Zibra et al., 2017), which coincides with 2727±8 Ma magmatic arc collision at the Kurnalpi-Burtville margin further to the east (Goscombe et al., 2007, 2009; Korsch et al., 2013). These Meso- and early Neoarchaean greenstone sequences are also recognised in the EYC, indicating that middle Neoarchaean greenstone sequences of the EYC formed on extended Youanmi crust (e.g. Cassidy and Champion, 2004; Czarnota et al., 2010a). Extension and rifting of the eastern margin of the Youanmi Terrane is consistent with progressively deeper crustal levels being preserved by regional metamorphic parageneses from west to east: 5.25-10.5 km in the Murchison domain, 12.25 km in the Southern Cross domain and 14.0-15.75 km in the Kalgoorlie and Kurnalpi terranes (Goscombe et al., 2009). Neoarchaean volcanic sequences spanning from ~2690 to ~2720 Ma formed in a protracted and probably episodic period of ENE-directed back-arc extension (D1), with plume-related basalt and komatiite volcanism in the Kalgoorlie Terrane (e.g. Barley et al., 2003; Groenewald et al., 2006). At the same time, calc-alkaline magmatic arcs of ~2680-2700 Ma age such as the Gindalbie arc formed in the Kurnalpi Terrane (e.g. Champion et al., 2001; Brown et al., 2002; Groenewald et al., 2006), in response to west-dipping subduction further to the east (e.g. Blewett et al., 2004a; Mary Gee and Swager, 2007; Champion and Cassidy, 2007; Czarnota et al., 2010a; Korsch et al., 2013). The Youanmi Terrane was also extended at this time forming greenstone and clastic sequences between ~2700-2730 Ma (e.g. Krapez et al., 2000; Groenewald et al., 2006), and calc-alkaline magmatic arc complexes such as the Manyutup tonalite (Witt, 1999; Korsch et al., 2013). Extension of the EYC continued through ~2670-2690 Ma, accompanying bimodal volcanism and voluminous high-Ca granite intrusion, interpreted to be related to subduction slab shallowing and hinge rollback (e.g. Blewett et al., 2004a; Czarnota et al., 2010a). Bimodal volcanism and high-Ca granite bloom of identical ~2670-2690 Ma age in the Slave Province, have also been interpreted to be due to a shallow subduction system (Bleeker and Hall, 2007). Extensional stress in the Yilgarn back-arc setting produced the volcanic basinal architecture, crustal-scale, listric to shallow dipping extensional shear zones (Blewett et al., 2004a; Blewett and Czarnota, 2005), and extensional growth

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faults documented by stratigraphic thickness variation (Gauthier et al., 2007; Squire et al., 2007). Extension of the EYC is constrained by basalt and komatiite volcanics with radiometric dates of ~2665-2705 Ma (Nelson, 1997; Claoue-Long et al., 1988; Krapez et al., 2000) and gneissic fabrics in granitoids with dates of 2670-2681 Ma (e.g. Nelson, 1996, 1997). All subsequent deformation periods rotated these early extensional structures into steep orientations and produced relatively steep, sub-parallel transposed foliations that are strongly partitioned in shear zones. 3.2. Contraction and Termination of Volcanism

Termination of subduction and bimodal volcanism in the Youanmi Terrane occurred between ~2700 and ~2715 Ma, immediately followed by formation of an unconformity formed by contraction, erosion and deposition of clastic sequences of ~2690-2710 Ma age (detrital zircon maximum deposition ages of ~2710, ~2702 and ~2693 Ma) (e.g. Groenewald et al., 2006; Kositicin et al., 2008; Van Kranendonk and Ivanic, 2009). Termination of subduction and volcanism in the EYC coincides with a major regional unconformity at ~2673 Ma at the top of the Black Flag Formation, followed by coarse clastic sedimentation of the White Flag Formation. This unconformity marks the first major contraction (D2), exhumation and erosion period with diachronous deformation progressing from ~2670 Ma in the Kurnalpi Terrane to ~2665 Ma in the Kalgoorlie Terrane (Czarnota et al., 2010a). Peak high-Ca granite magmatism at ~2660-2675 Ma and associated regional-contact metamorphism (M2), coincides with the cessation of volcanism and D2 contraction (Champion et al., 2001). ENE-WSW shortening produced large-scale tight folds and thrusts without pervasive foliation and mesoscopic fold development, at moderate though highly partitioned strains (Figure 6; Blewett et al., 2004a; Blewett and Czarnota, 2005). D2 contraction did not affect postvolcanic fluvial late basins, limiting deformation to before ~2660-2665 Ma (Blewett et al., 2004a,b). Similarly, deformed and cross cutting granites constrain D2 contraction between 2664-2668 Ma in the Burtville Terrane and 2660-2667 Ma in the Kurnalpi Terrane (Champion et al., 2001; Blewett et al., 2004a). Termination of subduction and the stress switch to ENE-WSW contraction is interpreted as either an advancing subduction system or choking of the subduction zone by an oceanic plateau or magmatic arc (Blewett and Czarnota, 2007b). 3.3. Late Rift Basins

There are two generations of post-volcanic, rift basins that filled with clastic sediments; the first generation of late-stage basins contains marine turbidite fill and second generation are fluviatile filled grabens (Painter and Groenewald, 2001; Krapez et al., 2000). The post-volcanic clastic sequences are crucial time markers across the whole craton that develop pristine M3 metamorphic assemblages. They do not exhibit the earlier metamorphic events that pre-date these late basins, allowing clear separation of M2 and M3 metamorphic parageneses. Turbiditic late basins contain local and distal sources of detritus including from granite domes and examples include; Merougil basin in the Kalgoorlie Terrane, Kanowna Belle, Granny Smith, Wallaby and Belches basins in the Kurnalpi Terrane and Diemal basin in the Youanmi Terrane. Turbiditic late basins in the EYC formed by NW-SE extension (D3a), which produced rifts in the hanging wall of extensional detachments that wrap around the margin of granite batholith domes. Growth faults controlling these turbiditic basins have reactivated older growth structures that controlled the development of synvolcanic basins (Squire et al., 2007). NW-SE trending crustal-scale shear zones such as Celia, Zuleika and Keith-Kilkenny are transfer structures linking the arcuate extensional detachments wrapping around the granite domes (Williams and Whitaker, 1993; Blewett and Czarnota, 2005). Extension was at high-angles to the D2 shortening direction, producing basins oblique to the tectonic grain and compatible with the D2 stress field. Consequently, turbiditic late basins are interpreted as oblique extensional rifts coeval with D2 contraction in an overall dextral transpressional system (Blewett et al., 2004a; Blewett and Czarnota, 2005). Coeval timing is confirmed by D2 folding of turbidite sequences and deposition constrained between 2663-2670 Ma by detrital zircon maximum deposition ages and cross cutting intrusions (Krapez et al., 2000; Squire, 2006, 2007; Blewett and Czarnota, 2007c). Oldest post-

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volcanic turbidites are the White Flag Formation in the Kalgoorlie Terrane, followed by Granny Smith and Wallaby basins in the Kurnalpi Terrane, with 2669 Ma and 2665-2666 Ma maximum deposition ages respectively (Krapez et al., 2000; Squire, 2006, 2007). The Belches basin has 2666 Ma detrital zircon maximum deposition ages and overlaps some early extensional structures indicating it may be slightly younger than other turbiditic basins (Hall, 2006). Fluviatile late basins formed by ENE-directed extension (D3b) producing NNW-trending elongate rifts such as Scotty Creek, Pig Well and Kurrawang grabens, with detrital zircon maximum deposition ages between 2657-2662 Ma (Krapez et al., 2000). Maximum ages for extension are also constrained by 2660-2664 Ma granites overprinted by extensional fabrics (Blewett and Czarnota, 2007c). D3 extensional events are closely associated with mantle-derived syenite and mafic granitoids that give ages of ~2655-2670 Ma (Figure 4; Champion and Sheraton, 1997), and 2658±4 Ma Au-mineralization at Sunrise Dam (Brown et al., 2002). The turbidite and fluvial basins formed during a major lithospheric extension event that resulted in structural incision by crustal-scale extensional detachments with amphibolite facies metamorphism in the footwall to greenschist metamorphism in the hanging wall (Williams and Whitaker, 1993). A seismic reflection profile across the Kilkenny shear zone and hanging wall Pig Well basin shows a low-angle extensional shear zone and half graben geometry (Blewett et al., 2004a). A seismic reflection profile across the Ockerburry shear zone shows a 4 km wide, high-strain extensional detachment with listric geometry that rolls over a domal granitic core complex (Blewett and Czarnota, 2007c). The Ockerburry shear zone has extensional kinematics along its entire length and is the predominant crustal-scale extensional detachment that defines the boundary between Kalgoorlie and Kurnalpi terranes. Burial and rotation of grabens by listric faulting resulted in local shortening, folding and thrusting within late basin sequences (Blewett and Czarnota, 2007c). NNW-trending extensional detachments are the locus of formation of M3 metamorphic parageneses with anticlockwise P-T paths, both within the extensional shear zone fabrics and late basin sequences in the hanging wall (Goscombe et al., 2007, 2009). Extreme incision of the crust by these detachments buried late basin sequences in the hanging wall to depths of up to 12-16 km, such as in the Belches basin (Goscombe et al., 2007, 2009). Prior to D3 extension, M2 regional-contact metamorphic parageneses formed at low-pressures of ~4-5 kbar, indicating the crust was not overthickened and thus D3 extension cannot be the result of gravitational collapse (Goscombe et al., 2007, 2009). Lithospheric extension related to late basin formation and M3 metamorphism is best understood as an external stress field, probably related to cessation of subduction, sag of the subducted plate and retreat of the hinge (i.e., slab roll back). This stress switch accompanies the transition from D2 ENE–WSW contraction during subduction to D4 NNW–SSE contraction during reactivation and cratonization. The oldest Au-mineralization events dated between 2658-2675 Ma are coeval with this stress switch and restricted to the EYC such as at Golden Mile, Kanowna and Jundee (Yeats et al., 2001; Gauthier et al., 2004; Ross et al., 2004). Post-volcanic clastic sequences overlapping in age with late clastic basins in the EYC have been documented in the Narryer, Southwest and Youanmi terranes, mostly on the basis of detrital zircon maximum deposition ages (Figure 4). Extensional basins containing conglomerate in Marymia inlier, north Youanmi Terrane, are associated with 2675±15 Ma syn-extensional porphyry (Bagas, 1999). Turbiditic Diemal basin in central Youanmi Terrane contains detrital zircon populations of 2729±9 Ma (Nelson, 2001) and 2702±17 Ma (Thebaud and Miller, 2009). However, maximum deposition ages may be significantly younger, and are interpreted as ~2675 Ma by correlation with Marymia inlier conglomerates (Bagas, 1999; Chen, et al., 2003). Late clastic sequences in the Southwest and Narryer terranes are reworked and overprinted at high metamorphic grades obliterating the original basin geometries. Those in the Southwest Terrane have detrital zircon maximum deposition ages of 2663-2676 Ma in the Lake Grace and Jimperding domains (Table 1), and ~2646 Ma in the Balingup domain on the western margin (Sircombe et al., 2007). Late clastic sequences from the Jack Hill belt in the Narryer Terrane are constrained by detrital zircon maximum

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deposition age of 2665 Ma and a diagenetic monazite minimum deposition age of 2660 Ma (Rasmussen et al., 2010). These geochronological constraints indicate deposition of late clastic sequences started from ~26632675 Ma in the WYC, overlapping in age with post-volcanic late basins in the EYC, and suggest that D3 lithospheric extension propagated across the entire craton (Figure 4). Extension of the WYC is supported by M3 metamorphic parageneses showing anticlockwise P-T paths (section 8.4.2), mapped stratigraphic incisions in the Cheritons Find greenstone belt (Bagas, 1994) and extensional kinematic indicators and flat-lying fabrics in the Southern Cross greenstone belt. Extensional indicators and the Diemal turbiditic basin are concentrated along the Youanmi-Cocanarup shear zone system, suggesting extensional reactivation. Extensional strain was relatively low in the WYC, as indicated by relatively minor occurrences of late clastic sequences, and the general absence of syenite and mafic granitoids. 3.4. Reactivation, Au-Mineralization and Cratonization

Subsequent to cessation of subduction and the ensuing lithospheric extension, the Yilgarn craton was reactivated by far-field stress producing low-strain and highly partitioned deformation events during cratonization (Figure 6). The first of these involved NNW-SSE contraction (D4) at 2645-2655 Ma, resulting in sinistral transpressive reactivation of NNW- and N-trending faults and reverse movements on ENE-trending structures across the EYC and Youanmi Terrane (Chen et al., 2001a,b; Blewett and Czarnota, 2007c). Post-kinematic metamorphic mineral growth overprints sinistral fabrics, indicating the M3 thermal anomaly persisted to at least 2645 Ma in the EYC (Ross et al., 2004; Salier et al., 2004). Sinistral transpression pre-dates almost all low-Ca granite magmatism. D4 sinistral transpression was responsible for the second and predominant Au-mineralization event in the EYC, by focusing fluid flow and opening precipitation sites. Major deposits such as Kalgoorlie, St Ives, Wallaby, Chalice, Lancefield, Mount Morgans, Granny Smith and Sunrise Dam have been dated or interpreted at ~2650 Ma (Salier et al., 2004; Miller, 2005; Brown et al., 2002; Bucci et al., 2004; Blewett and Czarnota, 2007c; Vielreicher et al., 2003; Ojala et al., 1997). Contraction and minor crustal thickening are inferred in the Narryer, Southwest and western Youanmi terranes on the basis of shallow clockwise P-T paths accompanying the late-stage regional thermal pulse and highgrade M4 metamorphism (Ridley et al., 1997). Seismic reflection profiles across Narryer and western Youanmi terranes show listric NW-dipping crustal-scale thrust systems responsible for shortening across this region (Korsch et al., 2013; Wyche et al., 2013). Stacked thrust geometries have been mapped in the Ravensthorpe belt (Witt, 1998). These thrusts overprint earlier crustal structures and are late in the history of the Yilgarn craton. It is probable that collision of the Narryer Terrane in the northwest Yilgarn craton was responsible for the D4 stress field that reactivated structures as far as the EYC. The second reactivation event involved a stress switch to NE-SW contraction (D5) and protracted reactivation of N- to NNE-trending faults by strike-slip, dextral transtension (Blewett and Czarnota, 2005, 2007c; McNaughton et al., 2005). This stress switch coincides with initiation of a protracted period of craton-wide low-Ca granite emplacement with most intruded between 2630 and 2655 Ma (Champion et al., 2001). Deformed low-Ca granites constrain a 2647-2652 Ma maximum age for dextral shearing, and cross cutting dykes in the Burtville Terrane constrain a minimum age of 2638±2 Ma (Blewett and Czarnota, 2005, 2007c). Dextral transtensional reactivation of EYC shear zones was at greenschist to sub-greenschist facies and resulted in appreciable decompression to ~1.0-1.8 kbar (Section 8.6). Small shifts in the stress field late in the orogenic cycle or during cratonization, is a common scenario resulting in transtensional reactivation and significant exhumation (Foster et al., 2009; Goscombe et al., 2017b). Crustal-scale terrane boundaries such as the Hootanui and the Ida-Waroonga shear zone systems were reactivated by dextral shear and high flattening strains with multiple stretching directions. Shear strain was locally very significant and was partitioned into the core of reactivated shear zones that hosted late-stage Au-mineralization (Figure 6). Younger Au-mineralization events at ~2610-2640 Ma accompanied D5 dextral transtension, and are documented at Golden Mile, Revenge, Chalice and Victory-Defiance in the

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EYC and deposits throughout the WYC (Bucci et al., 2004; Clark et al., 1989; McNaughton et al., 2005; Vielreicher et al., 2003; Tripp et al., 2007). Some EYC deposits are associated with <26362640 Ma lamprophyres indicating renewed mantle input along conduits opened by dextral transpression (Tripp et al., 2007). The youngest deformation (D6) in the EYC produced scattered, sub-horizontal crenulations and kinkbands formed by minor vertical shortening (<<10%) at shallow upper-crust levels of ~3 km depth (Swager, 1997; Davis and Maidens, 2003). Vertical shortening typically overprints low-Ca granites and is interpreted to be ~2635 Ma (Blewett and Czarnota, 2007c). Exposed rocks in the EYC were largely exhumed by this stage, so that these structures do not indicate lithospheric extension, and their formation at shallow crustal levels preclude crustal over-thickening (Blewett and Czarnota, 2005, 2007c). Probable mechanisms for heterogeneously distributed vertical shortening are late-stage tectonic unloading in response to diminishing tectonic driving forces (Duclaux et al., 2007), or flattening above low-Ca granite plutons (Blewett and Czarnota, 2007c). Metamorphic response to events subsequent to cratonization is beyond the scope of this paper (e.g. Fitzsimons, 2003; Lu et al., 2015). The craton was affected by extensional events such as mafic dyke swarms at 2410 Ma, 1210 Ma and 1075 Ma, and marginal rift basins at 1650-1800 Ma. Craton margins were variably reactivated and reworked during the Capricorn Orogeny (~1780-1856 Ma), Albany-Fraser events (1660-1690 Ma, 1130-1255 Ma and 1025-1120 Ma), Chittering-Balingup belt (650-680 Ma) and Pinjarra Orogeny (510-615 Ma). 4. GRANITOID MAGMATISM

Large volumes of Neoarchaean felsic magma intruded into the upper-crust of the Yilgarn craton with all granite types together constituting 65-70% of current exposure. High field-strength element (HFSE) granites and granodiorites form only a minor (5-10%) component of the EYC and are almost absent from the remainder of the Yilgarn craton (Champion and Cassidy, 2007). These are high-level intrusives associated with early magmatic arcs and volcanic complexes restricted to juvenile crust within the Kurnalpi Terrane (Champion and Cassidy, 2007). HFSE granitoids are associated with the first middle Neoarchaean magmatic peak in the EYC between 2680-2700 Ma and young westward, indicating a west-dipping subduction zone east of the Kurnalpi Terrane (Champion et al., 2001; Champion and Cassidy, 2007). High-Ca granitoids constitute the greatest volume (60-70%) of all granite types, are widely distributed and form the bulk of granite-gneiss domes. The earliest cycle of Neoarchaean high-Ca granite ranges ~2700-2730 Ma in age and is largely restricted to rifted older crust of the Youanmi and Burtville terranes. The subsequent and most predominant cycle of high-Ca granite magmatism ranges ~2640-2700 Ma in age and is widespread across Youanmi, Kalgoorlie, Kurnalpi and Burtville terranes (Champion et al., 2001). This cycle of high-Ca granite magmatism was diachronous, involved multiple peaks and large volumes emplaced both during subduction, and after termination of volcanism (e.g. Nelson, 1996, 1997; Fletcher and McNaughton, 2001; Champion et al., 2001; Dunphy et al., 2003). There is a general pattern of two high-Ca granite peaks in each terrane (Figure 4). The first starts with felsic volcanism at ~2700 Ma, increasing to a peak at ~2675 Ma in the Kalgoorlie, Kurnalpi and Burtville terranes, before termination of all volcanism between ~26702675 Ma (Champion et al., 2001). Slightly earlier peaks occurred in the Youanmi Terrane between ~2680-2690 Ma. The second high-Ca granite peak was at ~2660 Ma, and expressed in Youanmi, Kalgoorlie and Burtville terranes and is absent from the Kurnalpi Terrane (Figure 4; Champion et al., 2001). Early high-Ca granite magmas have low-Y and required melting of mafic rocks at high pressure leaving garnet-mafic or eclogite restite or involved a two-stage process of remelting initial highpressure melts. Granitic magmas generated by the melting of mafic rocks in the eclogite field, can form from subducting oceanic crust, or alternatively the base of crustal lithosphere with accreted mafic material (Champion and Sheraton, 1997; Champion and Cassidy, 2007). The tectonic setting

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for early high-Ca granite melts is interpreted to be subduction at a continental margin with a component of crustal recycling, probably as the crust thickened or subduction dip changed (Cassidy et al., 2002). Later high-Ca granite magmas of the ~2660 Ma peak, have high-Y indicating garnetfree restite and shallower melting, interpreted to form after cessation of subducted slab melting and with a higher component of crustal melt. The two granite peaks bracket the termination of volcanism, and probably signify a major switch in tectonic setting from subduction processes and crustal thickening, to lithospheric extension and shallower melting. Voluminous high-Ca granite magmatism thermally stratified the crust by advecting heat into the upper-crust and by radiogenic heat production from Th-enriched granites (Beakhouse, 2007), resulting in the widespread regionalcontact type metamorphic pattern during M2. Mafic granitoids (sanukitoids) and syenite make up a small component of the EYC, 5% and <1% respectively, and are almost absent from the Youanmi Terrane. Mafic granitoids are dark, amphibole-bearing diorites, granites, granodiorites and tonalite with age range ~2650-2680 Ma and peak at 2665 Ma in Kalgoorlie and Kurnalpi terranes (Champion et al., 2001). Syenites span a longer period and peak at 2665 Ma in the Kalgoorlie and Burtville terranes and ~2650 Ma in the Kurnalpi Terrane. Syenite melts are derived by partial melting of a metasomatized mantle source and thus related to earlier magmatic arcs concentrated in the juvenile Kurnalpi Terrane (Champion and Cassidy, 2007). Other mantle-derived magmas in the EYC include ~2636-2640 Ma lamprophyres and carbonatite (Groenewald et al., 2006; Tripp et al., 2007). These mantle derived melts are associated with dry CO2-rich mantle fluids and carbonate veins, and are closely associated with gold deposits and possibly related to transfer of gold from mantle sources into the upper-crust (Blewett and Czarnota, 2007c; Walshe et al., 2008a,b). Low-Ca biotite granite and granodiorite were emplaced craton-wide with greatest concentration in the eastern Youanmi Terrane, and in significant volumes that constitute 30-35% of all granitoids. Intrusions range in size from dykes to large sheet-like bodies that typically post-date penetrative deformation events (i.e., D1, D2 and D3) and crustal growth processes such as volcanism and arc magmatism. Most were emplaced between ~2630-2655 Ma, and young westward across the craton: 2630-2660 Ma in Burtville, 2620-2655 Ma in Kurnalpi, 2600-2655 Ma in Kalgoorlie, 2600-2650 Ma in Youanmi and 2570-2645 Ma in Southwest terranes (Figure 4; Champion et al., 2001). Low-Ca melt was generated by partial melting of the middle- to lower-crust dominated by high-Ca granite, meta-basalt and meta-sediment (Champion and Sheraton, 1997). These melts do not contain a mantle component, nor involve subduction processes, and are interpreted to be associated with lithospheric extension, footwall decompression below extensional detachments, and lower-crust and mantle lithosphere delamination (Champion et al., 2001; Champion and Cassidy, 2007; Czarnota et al., 2010a). Low-Ca crustal melts transferred heat from the lower-crust to shallow crustal levels aiding cratonization. 5. NEOARCHAEAN METAMORPHIC HISTORY

The Yilgarn craton experienced a protracted middle Neoarchaean orogenic cycle spanning from ~2750 to ~2620 Ma that included: rifting of cratonic margins, subduction, arc accretion, basin closure, lithospheric extension, lower-crust and mantle lithosphere delamination, reactivation and finally cratonization (e.g. Blewett et al., 2008, 2010; Czarnota et al., 2010a). At largest scale and time frame, these events progressed from predominating in the EYC, to being felt craton-wide (e.g. Korsch et al., 2013). The whole orogenic cycle involved at least five temporally distinct metamorphic periods, developing characteristic parageneses in different tectono-stratigraphic settings and experiencing contrasting P-T evolutions (Table 2; Figures 3 and 4; Goscombe et al., 2007, 2009). [1] Ma low-P/high-T granulite facies metamorphism within magmatic arcs, probably limited to ~2690-2730 Ma. [2] M1 high-P/moderate-T metamorphism associated with accretion events and partial burial of magmatic arc margins at 2748±19 Ma, 2727±8 Ma and 2706±10 Ma. [3] M2 low-P/moderate-T regional-contact type metamorphism between 2665-2685 Ma, associated with high-Ca granite bloom and contraction. [4] M3 low-P/moderate-T anticlockwise P-T path

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metamorphism at 2656±5 Ma, associated with lithospheric extension and post-volcanic clastic basins. [5] M4 low-P/high-T regional metamorphism in the WYC between 2615-2650 Ma was associated with renewed contraction and arrival of a diffusion-delayed thermal anomaly resulting from lower-crust and mantle lithosphere delamination. [6] Wanning thermal flux continued to reset isotopic systems between 2610 and 2545 Ma, with cratonization complete before ~2410 Ma intrusion of Widgiemooltha mafic dykes. The five metamorphic events are defined on the basis of simple overprinting criteria between structural, magmatic and metamorphic mineral growth events, and contrasting patterns and conditions of metamorphic response in different tectono-stratigraphic settings (Table 2; Goscombe et al., 2007, 2009). These tectono-stratigraphic and petrology-based constraints have proven robust and consistent, and were built on a well-established foundation of age constraints for clastic and volcanic stratigraphic units, granite magmatism and deformation events in the literature (e.g. Section 3; Krapez et al., 2000; Champion et al., 2001; Blewett et al., 2010; Czarnota et al., 2010a; Korsch et al., 2013; etc.). Tectono-stratigraphic age constraints on the different metamorphic events are tested and quantified further by direct dating of metamorphic minerals (Section 9). 5.1. Ma Magmatic Arc Metamorphism

Ma granulites formed in small domains of low-P/high-T static metamorphic conditions typical of magmatic arcs (Hallberg, 1985; Goscombe et al., 2007, 2009). These localities are spatially restricted to known magmatic arc settings, such as the Gindalbie magmatic arc in the Kurnalpi Terrane, which has associated HFSE granites of ~2675-2715 Ma age (Champion et al., 2001; Champion and Cassidy, 2007; Barley et al., 2008). Other localities are in the western Burtville Terrane associated with rifted old lithosphere and volcanic stratigraphy of ~2730-2810 Ma age (e.g. Champion and Cassidy, 2010). It is probable that Ma parageneses also occur in other magmatic arc complexes, such as the Melita, Welcome Well and Yandal complexes and magmatic arc complexes in the Duketon domain along the western margin of the Burtville Terrane (Groenewald et al., 2006; Champion and Cassidy, 2007). Metamorphic age determinations are not available, though the relict, scattered distribution and association with old stratigraphy and early HFSE magmatic events indicate they formed early. The magmatic arc settings are also consistent with being early-formed and later accreted during M1 arc docking events. Multiple magmatic arc accretion events dated from M1 parageneses at ~2748 Ma, ~2727 Ma and ~2706 Ma (Section 9.3), indicate Ma metamorphism was in all probability diachronous between different magmatic arcs, with low-P, high-grade parageneses forming prior to each of these arc docking events (Korsch et al., 2013). 5.2. M1 High-Pressure Metamorphism

M1 metamorphism was at high-pressures (>7.0 kbar) and high-P/moderate-T (Barrovian-series) thermal regimes (<25 ºC/km), distinct from all other events in the Yilgarn craton. M1 parageneses are characterized by relict, early-formed garnet-clinopyroxene-amphibolite or garnet-staurolitemetapelite assemblages preserved in low-strain shear lenses (1-1000 m) within crustal-scale shear zones (Hallberg, 1985; Swager, 1994c; Swager and Nelson, 1997; Goscombe et al., 2007, 2009). These shear lenses are typically reworked and enveloped by foliations with lower-P parageneses typical of the subsequent metamorphic events (i.e. M2 and M3). Most high-P occurrences are restricted to shear zones within the Kalgoorlie and Kurnalpi terranes, with few from shear zones in the eastern Youanmi Terrane and western Burtville Terrane (Figure 5). All major shear zones that define terrane boundaries preserve these high-P assemblages, such as the Ida, Ockerburry and Hootanui shear zones. High-P assemblages are also concentrated within linking shear zones throughout the EYC such as the Celia, Laverton and Perseverance shear zones, and in arcuate shear zones wrapping elongate granite-gneiss domes, such as the Laverton, Gindalbie, Norseman and Boorara domes. Rare high-P localities in the Youanmi Terrane are concentrated in the Westonia shear zone at the boundary with the Southwest Terrane or scattered along the Youanmi-Cocanarup shear zone system, and in thrust sheets on the west and northeast margin of the Ravensthorpe belt

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(Figure 5). Moderately high-P amphibolites that formed between 5-9 kbar and 520-620 ºC have been documented from the west Marymia Inlier, northern Youanmi Terrane (Gazley et al., 2011). 5.3. M2 Regional-Contact Type Metamorphism

M2 metamorphism is characterized by widespread low-P/moderate-T (Buchan-series) conditions that formed static to foliated parageneses in association with heterogeneous D2 contractional deformation and clockwise P-T paths (e.g. Goscombe 2007, 2009). Metamorphism is intimately associated with voluminous high-Ca granite magmatism, giving regional-contact type metamorphic patterns with lowest grades in the core of greenstone synforms and highest grades adjacent to granite-gneiss domes (e.g. Binns et al., 1976; Hallberg, 1985; Ahmat, 1986; Wilkins, 1997; Mikucki and Roberts, 2003). Isograds are sub-parallel to granite batholith margins and lowest-grade domains are most distant from the granites. Metamorphic grades range from sub-greenschist and greenschist facies in synform cores, to middle- and upper-amphibolite facies gneisses within granite-gneiss domes (Figure 7). M2 regional-contact metamorphic patterns predominate across the Kalgoorlie, Kurnalpi and Burtville terranes and much of the northeast Youanmi Terrane (Figure 7; e.g. Binns et al., 1976; Ahmat, 1986; Watkins and Hickman, 1990; Mikucki and Roberts, 2003; Goscombe et al., 2009). The Narryer and Southwest terranes preserve no evidence of regional-contact metamorphic patterns; and high-Ca granite magmatism and M2 metamorphic age determinations are almost absent, suggesting the M2 thermal anomaly did not extend this far west (Figure 4). The Narryer, Southwest and western Youanmi terranes were later metamorphosed at middle amphibolite to granulite facies grades, variably over-growing and obliterating earlier parageneses with new M4 matrix parageneses. The age of M2 metamorphism is constrained by correlation with the peak of high-Ca granite magmatism, D2 contraction and termination of volcanism (Section 3.2; Blewett et al., 2004a), and minimum age limited by the start of lithospheric extension at ~2665 Ma (Section 3.3). The near craton-wide thermal anomaly associated with all types of high-Ca granite was long-lived and probably spanned 2650-2690 Ma, leading to high thermal gradients throughout M2 and M3 metamorphic events. The greater volume of early high-Ca granites (low-Y), associated with contraction and termination of volcanism, were emplaced between 2665-2685 Ma. This is considered the peak of M2 metamorphism (Hall, 1997; Goscombe 2007, 2009), and consistent with M2 metamorphic age determinations averaging 2671±6 Ma (Section 9.4). 5.4. M3 Extensional Metamorphism

M3 metamorphism is characterized by low-P/moderate-T (Buchan series) conditions that formed parageneses with anticlockwise P-T paths, localized to post-volcanic late basins or D3 extensional shear zones (e.g. Goscombe 2007, 2009). All post-volcanic, late basins have been metamorphosed to biotite-grade or higher (Krapez et al., 2000; Blewett and Czarnota, 2005), and reach middleamphibolite facies in the Belches Formation (Painter and Groenewald, 2001; Goscombe et al., 2007, 2009). M3 parageneses with anticlockwise P-T paths are documented at localities along the length of the extensional Ockerburry shear zone at the Kalgoorlie - Kurnalpi terrane boundary (Section 6.4.2). This major crustal structure was reactivated by extensional top down to the east transport along its entire length, and is possibly the main controlling extensional structure during D3 (Czarnota et al., 2010a,b). M3 parageneses potentially developed in other major shear zones reactivated by D3 extension, such as the Ida and Hootanui, and extensional shear systems on the flanks of granitegneiss domes (Blewett and Czarnota, 2005, 2007c). Significant decrease in metamorphic grade across extensional structures has been documented in the Ockerburry (Williams and Currie, 1993), Keith-Kilkenny (Williams and Whitaker, 1993), Kunanalling, Ida, Hootanui and Laverton shear zones (see Figure 22). Because post-volcanic late basins and extensional shear zones post-date D2 contraction and M2 metamorphism, the metamorphic parageneses developed within these settings must be M3 or younger. M3 mineral growth may have also occurred in the underlying volcanic sequences, but

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remains obscure and unrecognised unless entirely reworked by extensional shearing or where there is compelling evidence for anticlockwise P-T paths. As a consequence, the minimum pattern and extent of M3 thermal anomalies is arcuate domains associated with late basin sequences and extensional shear zones, where they have been recognised (Figure 8). M3 anomalies are in all probability wider than these erosional-structural relicts and interpreted to extend to regions with significant volumes of late high-Ca granite (high-Y), mafic granite and syenite. Where late basin sequences have been removed by erosion, or detailed kinematic analysis is not available, M3 metamorphism will go unrecognised. Almost all expressions of M3 metamorphism are restricted to the Kalgoorlie and Kurnalpi terranes, which were strongly extended during D3 (e.g. Blewett and Czarnota, 2005, 2007c; Czarnota et al., 2010a). Extensional basin development and M3 metamorphism are inferred in the Youanmi and Southwest terranes on the basis of clastic sequences deposited after ~2663-2676 Ma (Bagas, 1999; Chen, et al., 2003; Sircombe et al., 2007; Rasmussen et al., 2010). M3 metamorphism is inferred in the Youanmi Terrane on the basis on relict, earlyformed low-P cordierite, and anticlockwise P-T paths in the Diemals Formation and Murchison, Southern Cross and eastern Ravensthorpe belts (Figure 8). The age of M3 metamorphism is constrained by correlation with the peak of late high-Ca granite (high-Y) magmatism, D3 extension and post-volcanic late clastic basins, all of which are associated with lithospheric extension (Section 3.3). Late high-Ca granites are associated with crustal melting and were emplaced between 2650-2665 Ma in all terranes, this being considered the peak of M3 metamorphism (Figure 4). Earliest post-volcanic late basins are turbidite sequences with maximum deposition ages range 2665-2669 Ma (Krapez et al., 2000; Squire, 2006, 2007). Renewed extension after 2660-2664 Ma (Blewett and Czarnota, 2007c) led to fluviatile sequences with maximum deposition ages that range 2657-2662 Ma (Krapez et al., 2000). The onset of extension at ~2665 Ma coincides with the peak of mantle-derived mafic granite and syenite intrusions (Champion et al., 2001; Champion and Cassidy, 2007). Termination of M3 metamorphism was controlled by dissipation of the thermal anomaly resulting from lithospheric extension and thinning. The minimum age limit for the peak of M3 conditions coincides with the cessation of deposition in late clastic basins at ~2650 Ma, followed by decay of the thermal anomaly. All of these constraints indicate the thermal peak of M3 metamorphism was between ~2550-2555 Ma, consistent with metamorphic age determinations averaging 2656±5 Ma (Section 9.5). 5.5. M4 High-Grade Metamorphism

M4 metamorphism is characterized by late-stage, low-strain regional metamorphism at low-P/high-T conditions with parageneses showing clockwise P-T paths indicating contraction (i.e. D4 and D5) and minor crustal thickening (e.g. Goscombe 2009). The Narryer and Southwest terranes experienced upper-amphibolite to granulite facies metamorphism that completely obliterated earlier parageneses. Metamorphic grade is lower- to middle-amphibolite facies in the western Youanmi and northeast Narryer Terranes, such as in the Ravensthorpe, Forrestania, Cheritons, Southern Cross, Bullfinch, Perenjori, Yalgoo, Dalgaranga, Murchison and Jack Hills belts (Figure 9). In these regions, M4 mineral growth variably developed new low-strain matrix assemblages, post-kinematic porphyroblasts and corona textures, leaving remnant regional-contact metamorphic map patterns (Figure 9), and relict M2 and M3 mineral parageneses (Section 5.4). At highest grades in the Bullfinch, Southern Cross, Cheritons and Forrestania belts, M4 mineral growth totally replaced earlier parageneses and the M2 regional-contact metamorphic pattern is largely obliterated (Figure 9). M2 and M4 in the Youanmi Terrane experienced similar metamorphic grades and clockwise P-T paths, making it impossible to accurately map the extent of M4 metamorphism. In the absence of widespread stratigraphic units deposited between these two events, M2 and M4 parageneses are largely indistinguishable except where they have been directly dated, or relict M3 minerals are preserved, inferring the matrix assemblage is M4. Age determinations show that M4 metamorphic parageneses predominate in the western Youanmi Terrane such as in the Southern Cross, Cheritons, Dalgaranga and Murchison belts (Figure 9), and relict M3 parageneses are identified in the Ravensthorpe, Southern Cross and Dalgaranga belts (Section 6.4.3). Regional metamorphic parageneses in the northeast and west Ravensthorpe belt overprint relict M1 (high-P) and M3 (low-

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P) parageneses and are interpreted to have formed during M4 metamorphism. The western margins of the Ravensthorpe, Forrestania and Southern Cross belts show steep, smooth metamorphic gradients from middle- to upper-amphibolite facies, and grade into granulite facies in the Southwest Terrane. M4 regional metamorphic parageneses have not been recognized in the northeast Youanmi Terrane and EYC (Figure 9). However, patches of greenschist facies alteration and Au4mineralization of 2610-2640 Ma age occur in these regions, indicating low-grade thermal anomalies overlapping in age with M4, extended as far east as the Ockerburry shear zone (Section 9.7; Figure 4). The age of M4 metamorphism is constrained to being younger than all expressions of D3-M3 lithospheric extension, such as late high-Ca granite magmatism and post-volcanic late clastic sequences (Section 5.4). The Southwest and Youanmi terranes show a relatively sharp break between high-Ca and low-Ca granites at ~2650 Ma, and this is considered the onset of M4 metamorphism. Low-Ca granites were generated by crustal melting and required a craton-wide highT anomaly at the base of the crust to produce the large volume of melt (Champion and Cassidy, 2007). These granites range in age from ~2615-2650 Ma across the Southwest and Youanmi terranes, indicating a protracted M4 thermal anomaly. A minimum age for M4 metamorphism is constrained by a post-kinematic, unmetamorphosed aplite dyke of 2620±6 Ma age in the Southern Cross belt (Bloem et al., 1995). Metamorphic dates from high-grade regional metamorphic parageneses in the Narryer, Southwest and Youanmi terranes range from ~2613 to ~2649 Ma (Section 9.6), consistent with the magmatic and stratigraphic age constraints. 6. METAMORPHIC PETROLOGY

Extensive sampling of metamorphic rocks from across the Yilgarn craton (n~1030) was guided by GA and GSWA field databases and focused on typical meta-tholeiite and metapelite bulk compositions. A select number of samples representing key stratigraphic units, tectono-metamorphic settings and metamorphic events are described below, summarized in Tables (3 to 6), and augmented by descriptions in the literature (e.g. Archibald et al., 1981; Bickle and Archibald, 1984; Ridley et al., 1997; Witt, 1998; Dalstra et al., 1998, 1999; Mikucki and Roberts, 2003; Mikucki, 1997). Comprehensive petrographic descriptions and microphotographs are available in Goscombe et al. (2009), and sample locations are available in Appendices (3 and 5). Interpretation of textural relationships and sequence of mineral growth is developed further with garnet zoning and PT determinations, in constraining P-T paths (Section 8). 6.1. Ma Magmatic Arc Mafic Granulite

Mafic granulite has been identified in at least four clusters in the Burtville Terrane and others in the core of the Gindalbie magmatic arc, Kurnalpi Terrane (Figure 5). All have medium-grained polygonal granoblastic matrix assemblages, with relict ophitic textures consisting of inter-locking recrystallized plagioclase laths and large amoeboid and interstitial aggregates of mafic minerals that pseudomorphed primary igneous pyroxenes (Figure 10e). The matrix assemblage consists of bytownite plagioclase, green-brown to brown magnesio-hornblende, orthopyroxene, salite to augitic clinopyroxene, ilmenite and minor rutile (Table 3). Quartz has not been confirmed. Orthopyroxene and clinopyroxene occur as large amoeboid poikiloblasts. Hornblende and orthopyroxene also occur as polygonal granoblastic grains interstitial to the large poikiloblasts. Orthopyroxene rarely develops thin coronas of hornblende(2). 6.2. M1 High-Pressure Amphibolite

High-P upper-amphibolite facies M1 assemblages have only been identified within low-strain shear lenses preserved within crustal-scale shear zones. These relatively high-grade assemblages were formed early and are overprinted and reworked to varying degrees by lower-grade fabrics in the enveloping shear zone. M1 assemblages develop medium- to coarse-grained polygonal granoblastic textures with aligned amphibole, clinopyroxene and plagioclase grains, defining an early grain shape

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fabric (Figure 10a,b,c). Three types of high-P mafic assemblages are developed: clinopyroxene, garnet±clinopyroxene and garnet-grunerite±clinopyroxene amphibolites (Table 3). Clinopyroxene amphibolites develop matrix assemblages of salite clinopyroxene, green-brown ferroan pargasitic or edenitic hornblende, andesine to labradorite plagioclase and minor quartz, magnetite, ilmenite and titanite (Table 3). Retrograde actinolitic hornblende partially replaces clinopyroxene, and epidote occurs as coronas around hornblende or quartz in contact with hornblende. Garnet±clinopyroxene amphibolites develop matrix assemblages of green ferroan pargasitic hornblende, bytownite to andesine plagioclase, Ca-almandine garnet and minor biotite, titanite, Mnilmenite, hematite and magnetite (Figure 10c). Clinopyroxene-bearing samples are uncommon (Table 3), and where present clinopyroxene is in textural equilibrium with the matrix assemblage and has salite compositions with Al-rich cores. Garnets typically form irregular-shaped synkinematic poikiloblasts (2-6 mm) that are in equilibrium with the matrix. Two generations of garnet are common: early syn-kinematic poikiloblasts and post-kinematic idioblastic overgrowths or small idioblasts over-growing the main foliation. Inclusion trails containing quartz, hornblende, ilmenite, magnetite and biotite are typically aligned with the matrix foliation. Most garnets have irregular margins showing minor resorption, and coronas of labradorite plagioclase(2) with idioblastic hornblende(2) are uncommon. Epidote is uncommon and overprints the main foliation, forming small post-kinematic porphyroblasts or coronas enclosing garnet and titanite. Garnet-grunerite±clinopyroxene amphibolites develop matrix assemblages of green ferro-pargasitic or ferro-edenitic hornblende, albite to oligoclase plagioclase, Ca-almandine garnet, grunerite and minor brown biotite, ilmenite and quartz (Figure 10d). There are three generations of grunerite: abundant large grunerite(1) inclusions within garnet porphyroblasts, minor grunerite(2) in the main foliation and grunerite(3) corroding garnet margins and overprinting the main foliation. Clinopyroxene-bearing samples are uncommon (Table 3), and where present clinopyroxene has hedenbergite compositions. Hornblende-grunerite aggregates are probable pseudomorphs after early clinopyroxene. Garnet occurs as post-kinematic idioblastic porphyroblasts (1-3 mm) that overprint the main foliation (Figure 10d). A few garnets are irregular shaped relict grains with resorbed margins and oligoclase coronas. Second generation hornblende(2) and biotite(2) occur as postkinematic laths that overprint the main foliation and corrode garnet and grunerite(2) margins. Barrovian-series metapelite schists are exceedingly rare, known only from a thin 30 cm wide band within mafic gneiss from the Ida shear zone. This sample (BG6-183a) has a medium-grained schistose matrix of muscovite, chlorite, quartz, and minor plagioclase, ilmenite and rutile. Small idioblastic staurolite grains overprint the matrix foliation and often closely associated with rutile. Garnet occurs as large (10 mm) idiomorphic poikiloblasts that are elongate and aligned with the main foliation (Figure 10f). Garnets occlude staurolite grains, are post-kinematic and overprint the main foliation that is preserved as linear inclusion trails. Calcareous-pelite schists have matrix assemblages of quartz, plagioclase, orange-brown biotite, muscovite and minor grunerite and ilmenite. Porphyroblasts (3-5 mm) of green magnesio-hornblende and poikiloblastic garnet are typically pre-kinematic, have resorbed margins and are enveloped by the main foliation. In highly schistose samples the garnet and hornblende are stretched into elongate lenticular shapes. 6.3. M2 Metapelite

Though M2 metamorphic assemblages predominate across the EYC and northeast Youanmi Terrane, aluminous and metapelite rocks with diagnostic mineral parageneses useful for metamorphic analysis are rare. Metapelites develop a medium- to fine-grained schistose matrix of varying intensity defined by aligned micas and quartz and plagioclase grain shapes. Chloritoid, staurolite, garnet, biotite, cordierite and andalusite porphyroblasts formed at different stages with respect to the main foliation (Table 4; Archibald et al., 1981; Bickle and Archibald, 1984). Three types of parageneses are developed: garnet, garnet-staurolite and garnet-free metapelites (Table 4). Garnet metapelites have matrix foliations consisting of quartz, oligoclase plagioclase, brown biotite, chlorite, ilmenite and minor muscovite, graphite, hematite and tourmaline (Table 4). Earliest formed

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phases are pre-kinematic chloritoid porphyroblast fish that are deformed, stretched and enveloped by the main foliation. Garnets are small (0.5-4.0 mm) rounded to idiomorphic porphyroblasts with minor quartz, biotite and ilmenite inclusions preserving early high-angle foliations. Most garnets formed relatively early and are enveloped by the main foliation with margins partially resorbed and corroded. Post-kinematic cordierite or andalusite poikiloblasts occur in a few samples, and overprint the main foliation, occlude garnets and are kinked and flattened by ongoing strain. Latest generation mineral growth is post-kinematic red-brown biotite(2), muscovite(2) or chlorite(2) laths that overprint the main foliation. Typical sequence of mineral growth in garnet metapelite is summarized as: Ctd > Grt > Qtz-Pl-Bt-Chl-Ilm±Ms foliation > Crd + And poikiloblasts > Bt2±Ms2±Chl2 laths Garnet-staurolite metapelites have matrix foliations consisting of quartz, andesine to bytownite plagioclase, orange biotite, ilmenite and minor chlorite and muscovite (Table 4). Large chloritoid laths are aligned with the main foliation in some samples (Bickle and Archibald, 1984). Relict chloritoid laths are overprinted by garnet, staurolite and biotite growth on margins, and corroded by cordierite moats that were later pseudomorphed to chlorite-muscovite (Table 4; Bickle and Archibald, 1984). Staurolite forms small to large idioblastic poikiloblasts that overprint both the main foliation and chloritoid laths (Figure 11a; Table 4; Archibald et al., 1981; Bickle and Archibald, 1984). Garnets are typically small (<2 mm) post-kinematic porphyroblasts with idiomorphic overgrowths that overprint the main foliation, chloritoid laths and staurolite grains (Figure 11a). Thin partial coronas of bytownite plagioclase(2) are rarely developed on garnet. Latestages of porphyroblast growth are biotite(2) laths that overprint the main foliation, and large andalusite crystals that occlude staurolite (Table 4; Bickle and Archibald, 1984). All cordierite growth is late, either as coronas enclosing chloritoid or as large poikiloblasts that occlude garnets and overprint andalusite (Table 4; Bickle and Archibald, 1984). Typical sequence of mineral growth in garnet-staurolite metapelite is summarized as: Bt-Pl-Qtz-Ilm±Chl±Ms±Ctd foliation > St > Grt > Bt2, and > Crd, Pl moats > Ms-Chl Garnet-free, staurolite, andalusite or sillimanite metapelites have matrix foliations consisting of quartz, oligoclase plagioclase, brown biotite, muscovite, ilmenite, chlorite and small staurolite grains (Table 4). Medium-grained, red-brown blocky laths of biotite(2) overprint the main foliation. Andalusite occurs as large skeletal poikiloblasts that over-grow the main foliation and biotite(2), and contain numerous quartz, ilmenite and biotite inclusions and irregular-shaped resorbed staurolite (Bickle and Archibald, 1984). Andalusite margins are embayed by partial moats of post-kinematic cordierite with relict ilmenite and biotite inclusions (Mikucki and Roberts, 2003; Bickle and Archibald, 1984). Andalusite is also corroded by partial moats of inter-grown muscovite, sillimanite and biotite(3), which are overprinted by fine-grained fibrolite and muscovite. Andalusite and cordierite margins are retrograded by muscovite and chlorite. Few samples have sillimanite, biotite and quartz matrix assemblages over-grown by cordierite porphyroblasts and late-stage biotite, chlorite and muscovite (Mikucki and Roberts, 2003). Typical sequence of mineral growth in garnetfree metapelite is summarized as: Bt-Ms-Ilm-Qtz-Pl-St±Chl > Bt2 > And poikiloblasts > Crd±Bt3±ms2±sill moats > Fib-Ms3 > MsChl 6.4. M3 Metapelite

6.4.1. Belches Formation, Kurnalpi Terrane Post-volcanic, turbiditic Belches Formation was deposited subsequent to M2 metamorphism and developed diagnostic low-P/moderate-T M3 metamorphic parageneses (Painter and Groenewald, 2001; Goscombe et al., 2007, 2009). Metapelites developed a fine-grained, bedding-parallel foliation that was overprinted by post-kinematic porphyroblasts of almandine garnet, staurolite, biotite,

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cordierite and andalusite (Figure 11b; Table 5). The matrix foliation consists of quartz, albite to oligoclase plagioclase, brown to orange-brown biotite, ilmenite and minor muscovite, chlorite, graphite and tourmaline. The matrix foliation is overgrown by early cordierite poikiloblasts that are flattened in the main foliation and over-grown by post-kinematic biotite(2), garnet, andalusite and staurolite porphyroblasts (Table 5; Mikucki and Roberts, 2003). Andalusite occurs as large (<10 mm) post-kinematic poikiloblasts across the main foliation and as inclusions within staurolite porphyroblasts. Andalusite is overgrown by garnet porphyroblasts and partially pseudomorphed by muscovite and biotite moats. Staurolite is common and occurs as large (<6 mm) post-kinematic idiomorphic poikiloblasts. Second-stage biotite(2) occurs as blocky laths at high-angles to the main foliation and as large porphyroblasts that contain main foliation inclusion trails (Figure 11b). Coarse ilmenite(2) grains also overprint the main foliation. Garnet is common and occurs as small (0.2-2.3 mm) idioblasts with fine-grained inclusion trails of quartz, monazite and ilmenite (Figure 11b), and overprint biotite(2) and andalusite. Late-stage plates of muscovite(2) and chlorite overprint the main foliation and biotite(2) laths. Fine-grained muscovite and chlorite forms on the margins of cordierite, staurolite, andalusite and biotite. Typical sequence of mineral growth in metapelite is summarized as: Qtz-Bt-Pl-Ilm±Ms±Gr±Chl foliation > Crd > And > St-Bt2-Ilm > Grt > Ms-Chl 6.4.2. Ockerburry Shear Zone, Kalgoorlie Terrane Diagnostic M3 parageneses showing anticlockwise P-T paths are also developed in aluminous schists within the extensional Ockerburry shear zone at the boundary between Kalgoorlie and Kurnalpi terranes. Mineral relationships are remarkably similar at widely separated localities at Mount Martin near Kalgoorlie, and Kyanite Hill near Leonora that have been described by Purvis (1978 and 1984). Most samples developed an intense fine-grained foliation of quartz aggregate ribbons, muscovite and fine skeletal andalusite(1) with minor sillimanite, biotite, chlorite, graphite, ilmenite and sanidine (Table 5). The main foliation enveloped, stretched and boudinaged prekinematic andalusite(1) porphyroblasts (Figure 11c). Post-kinematic idioblasts of kyanite or chloritoid overprint the main foliation and early andalusite and are undulose and kinked or remain undeformed (Figure 11d). Late-stage, second generation andalusite(2) formed a very fine-grained skeletal mesh along foliation planes and on kyanite margins (Figure 11d). Other aluminous schists from Mount Martin developed main foliations of quartz, muscovite, biotite, sillimanite, chlorite and graphite; overprinted by kyanite laths that were in turn over-grown by andalusite poikiloblasts (Table 5). A second generation of sillimanite(2), kyanite(2), muscovite and biotite were developed in late-stage shearbands that boudinaged and overprint the andalusite porphyroblasts. Aluminous schists from the Ockerburry shear zone on the margin of the Penny Dam late basin; developed a main foliation consisting of quartz, muscovite and rutile, overprinted by post-kinematic andalusite porphyroblasts, in turn overgrown by cross-cutting chloritoid laths (Mikucki and Roberts, 2003). Typical sequence of mineral growth in aluminous schists is summarized as: And > Qtz-Ms-And±Sill±Bt±Chl±Rt±Ilm foliation shearbands > Ms

> Ky + Ctd > And2 > Sill2-Ky2-Ms-Bt

6.4.3. Youanmi Terrane A small number of metapelite samples from widely scattered localities across the Youanmi Terrane have low-P prograde parageneses and anticlockwise P-T paths indicative of M3 metamorphism. Lower amphibolite facies metapelite schists from the post-volcanic Diemals Formation (BG9-27) develop a fine-grained bedding-parallel foliation of quartz, muscovite replaced by pyrophyllite and biotite replaced by chlorite (Table 5). Early-formed cordierite porphyroblasts are flattened parallel with the main foliation and overprinted by randomly oriented post-kinematic andalusite porphyroblasts and late chlorite plates. Similar low-P metapelite schists from the Dalgaranga belt in the Murchison domain (BG10-59), develop matrix foliations of quartz, oligoclase plagioclase, Mgmuscovite, orange-brown siderophyllitic biotite, graphite and ilmenite (Table 5). Large cordierite

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poikiloblasts (8 mm) occlude early foliations and were strongly flattened and rotated into low-angles with the enveloping main foliation before being partially pseudomorphed to chlorite, sericite and graphite. Post-kinematic andalusite formed as large (<10 mm) porphyroblasts that overprint the main foliation and are mortared by fine-grained muscovite(3) and chlorite. Second generation muscovite(2) occurs as large plates that overprint the cordierite and the main foliation. Secondary plagioclase(2) forms moats around biotite and muscovite plates and merge into sieve-like plagioclase poikiloblasts. Late-stage flattening of the main foliation is indicated by boudinaged andalusite, aligned pseudomorph minerals and crosscutting shearbands containing fine-grained muscovite(3). Typical sequence of mineral growth in metapelite is summarized: Early foliations > Crd > Qtz-Ms-Bt-Pl-Gr-Ilm foliation > And > Ms2 > Pl2 moat > Ms3-Chl-Gr flattening Four localities within the Ravensthorpe belt (BG6-197, BG6-198, BG9-164 and BG9-169) preserve early low-P parageneses overprinted by higher-P matrix assemblages, indicating either anticlockwise P-T paths or parageneses from two distinct metamorphic events. Aluminous schists contain prekinematic cordierite porphyroblasts that are boudinaged, stretched and enveloped by the main foliation and shearbands that contain quartz, oligoclase, biotite, chloritoid, chlorite, ilmenite, graphite and sillimanite (Figure 11e). Cordierite is overprinted by syn-kinematic andalusite, small (<1 mm) post-kinematic garnet and plagioclase poikiloblasts. More typical metapelite schists develop matrix assemblages of quartz, orange-brown siderophyllitic biotite, Na-muscovite, oligoclase plagioclase, chloritoid, ilmenite, graphite and minor rutile, chlorite, magnetite and tourmaline (Table 5). Chloritoid growth is early and precedes all other porphyroblastic phases: forming large pre- to syn-kinematic, resorbed and kinked laths (10-20 mm), aligned within or at lowangles to the main foliation (Figure 11f). All other porphyroblastic phases overprint the main foliation to varying degrees and relationships between them indicate a progression from synkinematic andalusite to post-kinematic biotite(2), muscovite(2), staurolite and finally peak metamorphic Mn-almandine garnet and late-stage chlorite plates. Andalusite occurs as large (10-20 mm) ellipsoid poikiloblasts that range from being syn-kinematic and preserving early oblique foliations, to post-kinematic and overprinting the main foliation. Large biotite(2) and muscovite(2) plates overprint the main foliation and typically pre-date staurolite. Staurolite occurs as postkinematic, idiomorphic poikiloblasts across the main foliation, and as elongate and kinked skeletal aggregates aligned with the foliation. Garnet is late-stage and occurs as small to medium (1-7 mm) post-kinematic idioblasts with numerous quartz, ilmenite and minor chloritoid inclusions. Garnet post-dates the main foliation, staurolite and andalusite, and is never occluded within andalusite. Retrograde chlorite and muscovite are aligned across the main foliation and partially replace andalusite margins. The typical sequence of mineral growth in metapelites is summarized by: Qtz-Ms-Bt-Pl-Gr-Ilm-Ctd±Rt±Chl±Mag foliation > Ctd2 > And > Ms2, Bt2 > St > Grt > Ms3-Chl 6.5. M4 Metapelite

6.5.1. Youanmi Terrane Middle-amphibolite facies assemblages in the western Youanmi Terrane, such as in the Southern Cross, Cheritons, Ravensthorpe, Dalgaranga and Murchison belts are interpreted to be M4 parageneses (Section 5.5). Metapelite schists from the Southern Cross and Cheriton belts show similar mineral relationships. Main foliation assemblages consist of quartz, siderophyllitic biotite, graphite, tourmaline, magnetite, ilmenite, andesine plagioclase, staurolite and rare muscovite (Table 4). Staurolite formed early and occurs as small highly resorbed relict grains aligned within the main foliation or preserved as inclusions within garnet or cordierite (Figures 12a,c,d). Almandine garnets are typically small (1.0-2.6 mm) post-kinematic idioblasts that overprint the main foliation and have few staurolite, chloritoid, biotite and ilmenite inclusions (Figures 12a,c). All andalusite growth is late-stage, forming large (>10 mm) idiomorphic post-kinematic porphyroblasts that occlude garnets without resorption of the grain margins (Figure 12b). Higher-grade samples contain sillimanite and

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biotite(2) laths that overprint andalusite and form in garnet pressure shadows. Cordierite is uncommon and when present is the latest phase, forming poikiloblasts or moats around andalusite (Table 4; Ridley et al., 1997). Cordierite is partially pseudomorphed to aggregates of fawn chlorite. The typical sequence of mineral growth in metapelites is summarized by: Qtz-Bt-Pl-St-Gr-Ilm-Mag-Tur±Ctd foliation > Grt > And > Sill-Bt2 > Crd > Ms-Chl The predominant metapelite assemblages in the Ravensthorpe and Dalgaranga belts are devoid of staurolite and chloritoid and show a common sequence of early garnet followed by late cordierite and andalusite. Matrix foliations consist of quartz, andesine to oligoclase plagioclase, orange-brown siderophyllitic biotite, ilmenite, graphite and minor tourmaline and Na-muscovite. Secondgeneration plagioclase occurs as poikiloblasts with round quartz inclusions and is devoid of garnet inclusions. Garnets occur as small (~1 mm) rounded to idiomorphic grains that are often occluded within cordierite and andalusite porphyroblasts (Figure 12f). Garnets are syn- to late-kinematic with cores that occlude early foliations trails, and inclusion-free idioblastic rims enveloped by the main foliation. Andalusite porphyroblasts are late-kinematic, weakly deformed and partially enveloped by the main foliation (Figure 12f). Andalusites commonly develop biotite coronas and are partially pseudomorphed by muscovite leaving relict andalusite grains. Cordierite occurs as large postkinematic poikiloblasts that over-grew garnet, andalusite and the main foliation, with no later flattening of the foliation around the cordierite (Figure 12f). Second-generation biotite(2) and chlorite occur as post-kinematic plates across the main foliation. The typical sequence of mineral growth in metapelite is summarized as: Qtz-Bt-Pl-Gr-Ilm foliation > Pl2-Grt1 > Grt2 rims > And > Crd, Bt2 > Chl-Ms Low-P metapelite schists from the Mount Magnet belt in the Murchison domain develop matrix foliations of quartz, muscovite, albite plagioclase, ilmenite and elongate skeletal andalusite aligned with the foliation (Table 4). Andalusite and plagioclase continued to grow and formed late-kinematic porphyroblasts that both overprint and are enveloped by the main foliation. Staurolite was formed early and preserved only as small sub-idioblastic inclusions within andalusite and overprinting early ilmenite inclusion trails (Figure 12e). Second-generation muscovite(2) and chloritoid laths overprint the main foliation and andalusite porphyroblasts. Late-stage flattening across the foliation kinks these laths. The sequence of mineral growth is summarized by: Early foliation > St > Pl-Qtz-And-Ms-Ilm foliation > Pl2 + And2 porphyroblasts > Ms2-Ctd 6.5.2. Southwest Terrane Aluminous granulites from the Southwest Terrane developed two types of assemblages: garnetcordierite metapelite granulite and orthopyroxene-garnet±cordierite granulites (Table 6). Garnetfree, two-pyroxene mafic granulites are common and described elsewhere (e.g. Nemchin et al., 1994; Rennie, 1998). All rock types developed medium- to coarse-grained polygonal granoblastic textures, and weak grain shape elongation and medium-grained domains with aligned phlogopite indicating an annealed deformation fabric. Matrix assemblages in metapelite granulites consist of quartz, orange-brown phlogopite, Mgcordierite, andesine plagioclase with oligoclase rims, unzoned pyrope-almandine garnet and minor sillimanite, ternary alkali-feldspar, Cr-Zn-hercynitic spinel, ilmenite and magnetite (Table 6). Sillimanite and hercynitic spinel are uncommon matrix phases, spinel typically enclosed by ilmenite and biotite, and sillimanite mostly occurring as inclusions within garnet. Garnet occurs as large (10 mm) rounded and resorbed porphyroblasts with inclusions of quartz, biotite, ilmenite and sillimanite. Garnets typically formed early, are variably resorbed and in extreme cases are irregular-shaped relict grains enclosed by thick cordierite(2)±biotite moats (Figure 13c). Thin plagioclase(2)-biotite(2) coronas corrode garnet margins in psammopelite samples (Figure 13b). Cordierite occurs as large ellipsoid porphyroblasts, typically in textural equilibrium with matrix garnet and phlogopite.

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Cordierite continued to grow after the peak of metamorphism, producing cordierite(2) moats around ilmenite, phlogopite and garnet. Late-stage cordierite(2) formed at the expense of garnet that is typically resorbed to varying degrees. Large, late-stage Mg-muscovite plates form at high-angles to the gneissic fabric, overprinting both plagioclase and garnet, and occluding small relict garnet grains. The typical sequence of mineral growth in metapelite granulite is summarized as: Qtz-Bt-Ilm±Sill inclusions > Qtz-Phl-Pl-Crd-Grt-melt±Kfs±Ilm±Spl±Mag > Crd2±Bt±Pl moats > Ms-Chl Orthopyroxene granulites have coarse equigranular matrix assemblages in textural equilibrium that consist of quartz, Al-rich Mg-orthopyroxene, pyrope-almandine garnet, orange-brown annitic biotite with phlogopite rims and andesine plagioclase with labradorite rims (Table 6). Some samples contain minor Mg-cordierite, graphite, sulphides and alkali-feldspar exsolution lamellae in matrix plagioclase and Mn-ilmenite inclusions within orthopyroxene and garnet. These granulites are devoid of sillimanite, spinel and magnetite. Orthopyroxene occurs both as large (<20 mm) porphyroblasts and smaller elongate grains, both in textural equilibrium with garnet, plagioclase and biotite (Figure 13d). Orthopyroxene porphyroblasts are either inclusion-free or sieve-like poikiloblasts containing garnet, quartz, plagioclase, biotite and ilmenite inclusions, indicating orthopyroxene growth outlasted garnet and biotite. Elongate orthopyroxene(2) grains and aligned biotite(2) define a weak grain shape fabric that envelops garnet porphyroblasts (Figure 13a). Thin partial coronas of quartz and labradorite(2) occur on orthopyroxene margins in contact with biotite. Garnet occurs as medium-sized (4 mm) rounded porphyroblasts with quartz, biotite, orthopyroxene and ilmenite inclusions in the cores. The smaller garnet grain size and rounded to irregular shapes are due to resorption, probably in response to late-stage cordierite and orthopyroxene growth. Thin quartz and biotite coronas develop on garnet margins. When present, cordierite occurs as large porphyroblasts, and second generation cordierite(2) occurs as medium-grained interstitial shapes and moats around garnet. The typical sequence of mineral growth in orthopyroxene granulite is summarized by: Qtz-Bt-Pl±Ilm > Opx-Grt-Qtz-Phl-Pl-melt±Crd±Ilm±Gr > Opx2-Bt2 fabric > Crd2 moats > Pl2, Bt2 > Chl 6.5.3. Narryer Terrane Metapelite granulites from Mount Narryer preserve textural relationships and a simple sequence of mineral growth in common that indicates a single metamorphic cycle, with no evidence for relict older parageneses, with the exception of relict zircon and monazite grains (Section 9.2). Metapelite granulites have a medium- to coarse-grained polygonal granoblastic matrix with moderate grain shape alignment of biotite, quartz, plagioclase and cordierite. Quartz and plagioclase are typically strained. Matrix assemblages consist of quartz, orange-brown siderophyllitic biotite, Fe-cordierite, bytownite to labradorite plagioclase with andesine rims, Mg-almandine garnet and minor microcline, apatite, zircon and monazite (Table 6). Ilmenite and magnetite are absent from most samples, and hercynitic spinel is recognised in only one sample. Sillimanite formed in early foliations that are preserved within garnet and cordierite cores, and is typically absent from the matrix. Fine-grained second-generation sillimanite(2) forms in reaction textures that corrode garnet and cordierite margins. Cordierite forms large elongate porphyroblasts (<10 mm) that are boudinaged with interboudin biotite growth and partially replaced by symplectites of biotite(2)-plagioclase(2). Garnet occurs as small (1-2 mm), rounded and resorbed porphyroblasts that overprint matrix biotite and contains minor inclusions of quartz, sillimanite, biotite and monazite. Garnets are commonly resorbed by cordierite(2) growth in moats and embayments, and extreme cases leave relict garnet grains within large cordierite(2) pools (Figures 13e,f). Garnet margins are also corroded by cordierite(2)-biotite(2) symplectites, and matrix biotite is partially replaced by worm-like growths and thin coronas of cordierite(2). Garnets are also enclosed by thin coronas of andesine plagioclase(2). The typical sequence of mineral growth in metapelite is summarized as:

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Sill-Bt-Qtz inclusions > Qtz-Bt-Pl-Crd-Grt±kfs±Ilm±Mag matrix > Crd2 moats > Pl2-Sill2-Bt2 > Ms-Chl 7. METAMORPHIC METHODOLOGIES

7.1. Mineral Chemistry

Mineral analyses were undertaken on the Cameca S51 electron microprobe at Adelaide University, with an operating voltage of 15 kV and 30 nA for all phases except micas (10 nA) and feldspar (15 nA), and a beam radius of 2 m for most phases and 4 m for micas and feldspars. These analyses have been augmented by silicate mineral analyses compiled from literature (Purvis, 1978; McQueen, 1981; Gole and Klein, 1981; Phillips and Groves, 1982; Bickle and Archibald, 1984; Spray, 1985; Neall and Phillips, 1987; Clark et al., 1989; Knight et al., 1993, 1996, 2000; Williams and Currie, 1993; Bloem et al., 1994; Dalstra, 1995; Mikucki, 1997; Witt and Davy, 1997; Witt, 1998; Dalstra et al., 1999). Non-ideal mineral end-member activities were calculated by the method of Holland and Powell (1990) using the program AX (Holland et al., 1998). Ideal mineral activities of 1.0 are assumed for sillimanite, kyanite, andalusite, rutile, titanite and quartz. Garnet compositional maps were undertaken on the Cameca S51 at a current of 100 nA and voltage of 15 kV, and average spacing between spots was 5 m. Mg counts were collected on the TAP crystal, Ca on PET crystal and Mn and Fe on LiF crystal. All mineral analyses and garnet compositional maps are available in Goscombe et al. (2009), mineral chemistry is summarized in Appendix (4), and representative garnet maps are presented in Figure (14). 7.2. PT Determinations by THERMOCALC

PT determinations by the statistically robust average-PT approach of Powell and Holland (1994) were undertaken using THERMOCALC (Powell and Holland, 1988, 1990) with the v3.25 thermodynamic dataset (Powell et al., 1998). By using the same internally consistent thermodynamic dataset, calculation method and error propagations for all PT calculations, the results are considered directly comparable between different rock-types and metamorphic grades (Powell and Holland, 1988, 1994). Average-PT calculations by THERMOCALC utilize the entire equilibrium assemblage by calculating all independent reactions between mineral end-members, and the results are independent of bulk composition. This approach avoids uncertainties associated with estimating bulk composition to calculate PT pseudosections (Section 7.5), and the many different thermodynamic calibrations used by conventional geothermobarometry (Section 7.3). The average-PT calculation methods used for regional data sets are described in Goscombe et al. (2017a, 2018). All results give good quality statistics with low e* values and satisfy the chi-squared test with sigfit value being less than the cut off value for 95% confidence (Appendix 5). Error propagation by THERMOCALC incorporates uncertainties on mineral end-member activities and errors on thermodynamic parameters. Average-PT calculations have been tested against the stability field of the matrix assemblage in generic petrogenetic grids (Section 7.5). In all samples the results overlap directly with the stability field and are thus considered plausible estimates of equilibration conditions (Appendix 5). The most robust average-PT results (n=287) are listed in Table (7), and these constitute the core metamorphic dataset for this project (Section 8). All calculations assume that the mineral assemblage was in equilibrium and require estimates of fluid composition. For lower- and middle-amphibolite facies samples, XH2O=1.0 has been generally assumed. At these metamorphic grades, variation in XH2O does not appreciably affect PT calculations and assemblages are largely free of re-equilibration effects, the results being considered generally reliable. For upper-amphibolite facies garnet- and clinopyroxene-amphibolites, XH2O=0.5 has been generally assumed. In amphibolite facies rocks with porphyroblastic garnet, peak metamorphic conditions were calculated using garnet rims in conjunction with the cores of matrix feldspar, staurolite and micas (Spear, 1993). Garnet rim analyses were taken from the furthest extent of prograde growth, typically 15-50 m from the garnet edge to avoid post-peak re-equilibrated margins (Spear, 1993; Kohn and Spear, 2000; Goscombe et al., 2005). Post-peak metamorphic

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conditions were calculated using the outer-most garnet rim in conjunction with rim analyses from matrix phases or core analyses from minerals in late-stage foliations. The conditions of formation of reaction textures and symplectites were calculated using the cores of symplectite phases in conjunction with the outer-most re-equilibrated rims of the reactant phases. For granulite facies samples, fluid absent conditions, and aH2O between 0.2 and 0.3 are generally assumed (Goscombe et al., 2017a, 2018). At these grades, PT calculations can be significantly influenced by aH2O. In these cases, results are based on the intersection between the P-T-aH2O array calculated using THERMOCALC, and temperatures from Fe-Mg exchange reactions or average-PT calculated independent of H2O and CO2, both also calculated using THERMOCALC (Goscombe et al., 2017a, 2018). Most results by this method are consistent with the phase stability field of the matrix assemblage, and aH2O estimates based on these intersections range between 0.1 and 0.4 (Table 7). Granulite minerals typically show little core to rim compositional zoning, and assumed to have been homogenized during, or immediately after the peak of metamorphism (Tuccillo et al., 1990; Spear, 1993; Goscombe et al., 1998, 2005). Consequently, peak metamorphic conditions are best represented by analyses from the cores of mineral grains, potentially away from the effects of post-peak mineral re-equilibration at grain margins. At granulite facies grades there is the potential for adjacent mineral grains to continue to re-equilibrate by mineral reactions and cation exchange during cooling (Frost and Chako, 1989). As a result, average-PT calculations from mineral cores represent peak metamorphic conditions at best, or more likely, conditions at some stage during postpeak cooling (e.g. Spear, 1993; Goscombe et al., 1998, 2017a, 2018). Most average-PT calculations from granulites in the Yilgarn craton (80%) are consistent with the phase stability field of the matrix assemblage. The remaining 20% were significantly re-equilibrated during cooling and return average-PT results at temperatures lower than expected for the matrix assemblage. In these cases, alternative peak metamorphic conditions (Table 7) were estimated on the basis of the stability field of the matrix assemblage within published PT pseudosections (e.g. Tinkham et al., 2001; White et al., 2001; Storm and Spear, 2005; Groppo et al., 2007; Kosulicova and Stipska, 2007; Diener et al., 2007; Ziaja et al., 2010; Longridge et al., 2017). 7.3. PT Determinations by Conventional Geothermobarometry

The vast majority of metamorphic rocks in the Yilgarn craton have insufficient mineral endmembers to define a set of independent reactions suitable for average-PT calculations using THERMOCALC. Consequently, to expand the dataset further for the purposes of metamorphic field gradient mapping, a wide range of conventional geothermometers and geobarometers were used to calculate peak metamorphic conditions from the remaining samples (n=330). Accepted peak metamorphic conditions are based on tight clustering and intersection of results by a range of different calibrations and mineral pairs (Appendix 5). This data was augmented by PT determinations from literature (n=422), most of which were based on conventional geothermobarometry (76%) and the remainder were estimates from petrogenetic grids. Approximately 67% of literature PT determinations assumed either pressure or temperature (Appendix 5; e.g. Gole and Klein, 1981; Bickle and Archibald, 1984; Williams and Currie, 1993; Bloem et al., 1994; Dalstra, 1995; Mikucki, 1997; Witt, 1998; Dalstra et al., 1999; and many others). Conventional geothermobarometry methods utilize a large range of mineral pairs and different thermodynamic calibrations and therefore, results vary widely and are not directly comparable. Furthermore, these methods utilize only a subset of the matrix assemblage and because error propagations are not calculated, equilibration cannot be tested and the results are considered less robust in general (e.g. Kohn and Spear, 2000; Dasgupta et al., 2009; Goscombe et al., 2018). Consequently, results from conventional geothermobarometry have only been utilized for metamorphic mapping, and not used to draw conclusions nor characterize metamorphic conditions for the different terranes and metamorphic events (Section 8). A wide range of different Fe-Mg exchange geothermometers, and net transfer geobarometer reactions were calculated using THERMOCALC v3.25 (Appendix 5). A large number of different conventional geothermobarometers have also been applied, as listed in Appendix (8). Alteration and

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mineralization conditions have been determined in literature by fluid inclusion studies (e.g. Ho, 1987; Clark et al., 1989; Ho et al., 1990; Bloem, 1994; Mikucki, 1997; and many others). 7.4. Estimation of Prograde Conditions

Prograde conditions have been estimated by two methods using garnet core analyses from samples where growth zoning patterns and original garnet compositions are still preserved (Table 7). [1] Average-PT calculations using THERMOCALC were undertaken on inclusion assemblages and garnet core analyses. [2] Estimates of prograde conditions were also made using garnet core analyses compared to compositional isopleths in published PT pseudosections for typical metapelite compositions (Spear, 1993; Storm and Spear, 2005; Gaidies et al., 2006, 2008; Kosulicova and Stipska, 2007; Jerabek et al., 2008). These are considered semi-quantitative estimates, because the pseudosections will not accurately resemble the bulk composition of all samples, and results can vary by ±10-20 ºC and ±0.5-1.0 kbar between different bulk compositions (Goscombe et al., 2017a, 2018). Given the inherent limitations in estimating bulk composition during prograde metamorphism, this approximate method is appropriate for large-scale, regional analysis, and shown to give reasonable estimates for prograde conditions in Phanerozoic orogenic belts (e.g. Goscombe et al., 2017a, 2018). The prograde PT results cluster with other samples from the same tectonometamorphic zone and are consistent with P-T trajectories constrained by evolving mineral parageneses (Sections 7.5 and 8). 7.5. Semi-Quantitative P-T Paths

The simplified method used to determine semi-quantitative P-T paths, from large numbers of samples in regional-scale projects, is described in Goscombe et al. (2017a, 2018). A combination of calculated PT determinations (Sections 7.2 and 7.4) and the sequence of mineral growth from petrology (Section 6) have been used to constrain the most probable P-T paths. Specifically, prograde conditions from garnet cores, together with peak and post-peak PT calculations using THERMOCALC, frame prograde and retrograde trajectories (Table 7). These PT pinning points and the trajectories between them are compared with petrology constraints from the sequence of mineral growth. For example, prograde inclusions, matrix assemblages, sequence of overprinting porphyroblasts and post-peak reaction textures, have been interpreted within generic petrogenetic grids based on published PT pseudosections with similar bulk composition to the sample. Because these pseudosections were not calculated specifically for Yilgarn samples, they are utilized only as generic petrogenetic grids representing the most probable phase relationships for samples of similar bulk composition. XRF whole rock analyses from a select number of metapelite and mafic samples were made for comparison with the bulk composition of PT pseudosections available in the literature (Figure 15). Metapelite samples show moderate compositional variation in AFM and fall into geographical clusters that overlap with published pseudosections for a range of typical metapelite compositions (Figure 15a). Pseudosections have been chosen that represent sample clusters and fields divided by garnet-chlorite and staurolite-biotite tie lines in AFM space (Figure 15a). Similarly, the mafic samples studied are representative of Yilgarn craton mafics in general, and these cluster around published pseudosections for typical tholeiite compositions (Figure 15b). Typically, the difference in bulk composition between the studied rock sample and pseudosection is less than 10% for most components. The generic topology of these PT pseudosections are expected to be broadly similar to one calculated specifically for the sample (e.g. Tinkham et al., 2001). The most utilized metapelite petrogenetic grids are MnNCKFMASH±TO model systems in different parts of P-T space, with bulk compositions appropriate for: [1] aluminous high-Fe metapelite (Powell et al., 1998; Storm and Spear, 2005 figure 6; Gaidies et al., 2006 figure 14; Gaidies et al., 2008 figure 9), [2] aluminous low-Fe metapelite (White et al., 2001; Johnson et al., 2003; Martinez et al., 2004 figures 6a and 7; Tinkham et al., 2001 figure 8), [3] sub-aluminous high-Fe metapelite (Jerabek et al., 2008, figure 8; Johnson et al., 2003, figure 2c; Tinkham and Ghent, 2005), and [4]

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sub-aluminous low-Fe metapelite (Tinkham et al., 2001, figure 5a; Longridge et al., 2017 figure 6). Partial petrogenetic grids for mafic rocks were constructed from diagnostic phase relationships in published PT pseudosections, based on different tholeiite compositions in the model systems: NCKFMASHTO (Groppo et al., 2007 figure 9; Phillips and Powell, 2010, figure 7), NCFMASH (Diener et al., 2007, figure 18a; Mahan et al., 2008 figure 11d; Clarke et al., 2006 figure 4a; Elmer et al., 2006 figure 1b), CFMASH (Pattison, 2003; Green and Ringwood, 1967), and CFASH (Zeh et al., 2005). Details of the different petrogenetic grids are summarized in the figure captions. PT calculations and the petrology of selected samples representing different tectono-metamorphic zones and metamorphic events; have been plotted on generic petrogenetic grids to illustrate typical constraints on the P-T paths (see Figures 16, 19, 20, 23, 25, 26, 28, 29 to 30). Similar plots for all samples analysed in this project are available in Goscombe et al., (2009). Large-scale projects elsewhere have shown this method gives robust characterization of P-T paths across different orogenic systems (Goscombe et al., 2017a, 2018). Comparison between PT calculations and the petrogenetic grids show strong concordance between peak metamorphic results and phase stability field of the matrix assemblage, with ~97% of all samples overlapping directly and ~80% of granulite samples overlapping. P-T paths from the same tectono-metamorphic zone and metamorphic event are mutually similar, and when pooled together these give a robust characterization of metamorphic evolutions in the different zones (see Figures 17, 18, 21, 24 to 27). This simple approach is considered preferable for large-scale programs because of the inherent problems in determining the effective bulk composition needed to calculate sample specific pseudosections (e.g. Gaidies et al., 2015), and has the potential to generate large relational datasets across orogenic systems and thus increase accuracy and confidence in the results (Goscombe et al., 2017a, 2018). 8. INTEGRATED METAMORPHIC RESULTS

Generalized metamorphic evolutions for different tectono-metamorphic zones and metamorphic events have been synthesised from petrology (Section 6), garnet zoning (Appendix 4), PT determinations (Table 7), semi-quantitative P-T paths (Section 7.5) and metamorphic age constraints (Sections 5 and 9). This synthesis also incorporates petrology constraints available in the literature (e.g. Archibald et al., 1981; Bickle and Archibald, 1984; Ridley et al., 1997; Witt, 1998; Dalstra et al., 1998, 1999; Mikucki and Roberts, 2003). Diagnostic metamorphic features, pooled average results and characteristic P-T paths for each of the Neoarchaean metamorphic events are summarized in Table (2) and described below. 8.1. Ma Metamorphism (~2715-2750 Ma)

Ma metamorphism produced low-P/high-T upper-amphibolite to granulite facies assemblages in a small number of widely scattered locations from the Gindalbie magmatic arc and western Burtville Terrane (Figure 5). Ghosted ophitic textures are preserved in metamorphosed gabbros, indicating static metamorphism. Polygonal granoblastic matrix assemblages are clinopyroxene-hornblendeplagioclase±orthopyroxene, with rutile-ilmenite pairs indicating reducing fO2 conditions (Table 3). Peak metamorphic conditions range ~680-780 ºC at pressures between 3.7-5.0 kbar, indicating very high T/depth ratios of 51-55 ºC/km (Table 7). Highest grades averaging ~740 ºC and 4.2 kbar are recorded in the Burtville Terrane (Table 2). There is an absence of relict, early garnet-bearing parageneses, or the development of post-peak garnet or plagioclase reaction textures, limiting prograde and retrograde metamorphism to low-P/T heating and cooling paths (Figures 16 and 17). The low-P/high-T static metamorphic conditions and low-P/T P-T paths are consistent with magmatic heat sources such as magmatic arc settings, and tight anticlockwise loops have been modelled for these settings due to prograde burial during magmatic arc growth (Spear, 1993; Bedard et al., 2003).

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8.2. M1 Metamorphism (2700-2750 Ma)

M1 metamorphism produced high-P (>7.0 kbar) upper-amphibolite facies parageneses at highP/moderate-T conditions, with annealed medium- to coarse-grained polygonal granoblastic gneissic textures. M1 parageneses formed early in the geological history and all occurrences are relict assemblages within low-strain shear lenses in major crustal-scale shear zones (Figure 5). Diagnostic metamorphic assemblages are garnet±clinopyroxene±grunerite amphibolites and garnet-staurolitemuscovite-chlorite-rutile metapelite schists (Table 3), and range down to lower-grade epidote amphibolite facies assemblages. Moderate grain shape fabrics developed within the shear lenses matrix assemblage, indicating dynamic metamorphism at high metamorphic grades. These matrix assemblages are variably overprinted by low-grade foliations developed during subsequent reactivation of the shear zones, particularly during D2-M2 and D3-M3. Peak metamorphic conditions determined by PT calculations from garnet porphyroblast rims and matrix minerals cores, range 580-620 ºC at pressures between 7.0-8.8 kbar, indicating uniquely low T/depth ratios of 18-25 ºC/km (Tables 2 and 7). Metamorphic conditions do not overlap with M2 regional-contact metamorphism (see Figures 16 and 19). Peak metamorphic conditions in the Burtville Terrane, average 619±27 ºC and 7.7±1.1 kbar, Hootanui shear zone averages 584±42 ºC and 7.5±0.7 kbar, Ida and other shear zones in the Kalgoorlie Terrane average 597±49 ºC and 7.5±0.8 kbar, and Youanmi Terrane high-P samples average 598±43 ºC and 7.4±0.5 kbar (Table 2; Figure 17). Peak metamorphic conditions (i.e. maximum-T) were attained at maximum-P, with no evidence from relict parageneses for higher pressures during the prograde history. Almandine garnets in high-P metapelite schists from the Ida shear zone, Kurnalpi Terrane and Ravensthorpe belt show typical growth zoning, with Fe2+, Mg and Ca increasing and Fe/(Mg+Fe) and Mn decreasing with growth (Figure 14). Some samples show irregular Ca patterns and thin outer rims rich in Mn and Fe/(Mg+Fe), and poor in Ca and Mg indicating resorption. Ca-almandine garnets from a range of high-P amphibolite assemblages, show typical growth zoning, with Fe2+ and Mg increasing and Fe/(Mg+Fe), Ca and Mn decreasing with growth. Few samples show thin outer rims rich in Mn and Ca and poor in Fe and Mg indicating resorption. Inverse growth zoning patterns are uncommon and these show Ca and Mg increasing, Fe/(Mg+Fe) and Fe2+decreasing and Mn either increasing or decreasing with growth (Figure 16). Prograde conditions based on compositional isopleths from metapelite garnet cores in the Ida shear zone, give estimates of ~4.5 kbar at ~525 ºC and ~8.2 kbar at ~480-500 ºC (Table 7; Figure 17a). Prograde conditions from high-P amphibolites based on THERMOCALC average-PT calculations using mineral cores and inclusions, give estimates between 3.7 and 7.9 kbar at 440-553 ºC (Table 7; Figure 17b). Together these results show a steep P/T array of prograde conditions ranging from low-pressures to pressures equivalent to the peak of metamorphism, indicating near isothermal burial and hairpin clockwise P-T loops (Figure 17). Average-PT calculations in THERMOCALC using variably resorbed outer rims of garnets, matrix mineral rims and plagioclase coronas, define steep P/T arrays for post-peak re-equilibration. Multiple post-peak PT determinations from high-P metapelite in the Ida shear zone, define an array between 7.8 kbar at 590 ºC and 5.0 kbar at 530 ºC, with lowest pressure results between 2.8 and 3.9 kbar at 480-550 ºC (Figure 17a). Numerous post-peak PT determinations from high-P amphibolites show similar steep arrays from sub-peak conditions at 5.7-7.3 kbar and 530-594 ºC, down to 2.8-5.0 kbar and 455-550 ºC conditions typical of regional M2 metamorphism (Figure 17). Collectively, these samples define near isothermal decompression paths from the peak of metamorphism. These near isothermal decompression paths are consistent with post-peak mineral growth in amphibolites, such as secondary grunerite, hornblende, actinolitic hornblende, and corona epidote and labradorite plagioclase (e.g. Green and Ringwood, 1967; Pattison, 2003; Groppo et al., 2007). The steep prograde paths, hairpin clockwise loops and near isothermal decompression is consistent with relatively fast burial and exhumation of buoyant crust, such as partial burial of magmatic arc margins during accretion.

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8.3. M2 Metamorphism (2665-2685 Ma)

M2 metamorphism produced regionally extensive, static to dynamic, low-P/moderate-T matrix assemblages in pre-late basin stratigraphy across the Kalgoorlie, Kurnalpi, Burtville and northeast Youanmi terranes. Extensive metamorphic mapping of the Yilgarn craton show metamorphic grades range from prehnite-pumpellyite to upper-amphibolite facies, corresponding to a temperature range of ~350-620 ºC (e.g. Mikucki and Roberts, 2003, Hallberg, 1985; Ridley et al., 1997; Goscombe et al., 2009). Peak metamorphic pressures are consistently low, typically range ~3.0-5.0 kbar and indicate elevated T/depth ratios between 30-50 ºC/km (Figure 18a,b; Ridley et al., 1997; Goscombe et al., 2007, 2009). Metamorphic temperatures increase toward granite-gneiss domes and metamorphic isograds parallel the batholith margins (e.g. Ahmat, 1986; Ridley et al., 1997; Mikucki and Roberts, 2003). These map patterns and elevated thermal gradients are consistent with regionalcontact type of metamorphism (e.g. Binns et al., 1976; Wilkins, 1997). The regional metamorphic anomaly is interpreted due to advection of large volumes (60-65%) of early high-Ca granite into the upper-crust (e.g. Ridley, 1993; Goscombe et al., 2007, 2009). This resulted in an inverted thermally stratified crust due to neutral buoyancy pooling of large volumes of granite at crustal depths immediately below the Kambalda Sequence (e.g. Czarnota et al., 2010a). Diffusion of magmatic heat was further enhanced by the anomalously high thorium contents in these granites, which also contributed significant radiogenic heat (Champion and Cassidy, 2007). Direct dating of M2 metamorphic minerals constrain peak metamorphism at 2671±6 Ma (Section 9.4), consistent with the early high-Ca granite peak at ~2675 Ma (Champion et al., 2001) and estimates for the age of D2 contraction at ~2665-2673 Ma (Blewett et al., 2004a,b). Peak metamorphic PT determinations from middle- to upper-amphibolite facies samples, offer strong controls on crustal depth and average thermal gradient variation across the craton. AveragePT calculations using THERMOCALC were undertaken on garnet rim analyses and matrix mineral cores (Table 7). Pooled average peak metamorphic results range from ~570 ºC at ~5.3 kbar in the Burtville and Kurnalpi terranes, ~558 ºC at ~4.8 kbar in the Kalgoorlie Terrane, and ~560 ºC at ~3.9 kbar in the Youanmi Terrane (Table 2). This systematic decrease in crustal depth westward across the craton is also repeated at higher resolutions with pooled averages from individual greenstone belts and sub-domains across the craton (see Figures 18a,b and 22; Goscombe et al., 2007, 2009). Lowest pressure results range 1.5-3.2 kbar at 470-600 ºC in Murchison greenstone belts, and 2.0-3.8 kbar at 502-604 ºC in greenstone belts from northeast Youanmi Terrane (Table 7). Average thermal gradient (T/depth ratios) are consistently high and average ~31 ºC/km in Burtville-Kurnalpi terranes, ~35 ºC/km in Kalgoorlie Terrane and ~45 ºC/km in northeast Youanmi Terrane, indicating elevated thermal regimes across the whole craton during M2. Al-in-hornblende geobarometry of granitoids within greenstone belts also document a decrease in pressure across the craton from 3.0-4.0 kbar in Kalgoorlie and Kurnalpi terranes, to 3.0-4.0 kbar in the Southern Cross region and 1.8-2.8 kbar in the Murchison region (Figure 18a,b). Metamorphic field gradients across individual greenstone belts show small increases in pressure from synformal cores (<550 ºC) to high-grade gneisses (>570 ºC) on the margin of granite-gneiss domes (e.g. Archibald et al., 1978; Ahmat, 1986; Ridley et al., 1997; Goscombe et al., 2007, 2009). Average pressures vary by only ~1.0-1.4 kbar across greenstone belts in the Youanmi Terrane, and ~0.4-0.6 kbar across belts in the Kalgoorlie and Kurnalpi terranes. These pressure increases are in concert with temperature and result in almost no variation in T/depth ratio from synforms to batholith margins. Specifically, within greenstone belts in the Kurnalpi Terrane, average pressures in synformal cores are ~4.9 kbar at 480-545 ºC, and ~5.3 kbar at 570-626 ºC in gneissic margins. In Kalgoorlie Terrane greenstone belts, average pressures in synformal cores are ~4.4 kbar at 500-545 ºC, and ~5.0 kbar at 570-640 ºC in gneissic margins. Across belts in the northeast Youanmi Terrane, average pressures in synformal cores are ~3.4 kbar at 470-565 ºC, and ~4.8 kbar at 575-635 ºC in gneissic margins. In Murchison greenstone belts, average pressures in synformal cores are ~3.3 kbar at 500-565 ºC, and ~4.3 kbar at 570-612 ºC in gneissic margins.

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In the Southern Cross region, Al-in-hornblende geobarometry from granitoids show a similar pressure gradient, from 3.2-4.0 kbar within greenstone belts up to 4.6 kbar on the margin of granitegneiss domes (see Figure 18a,b and 24a,b; Dalstra et al., 1998). P-estimates from the core of the Ghooli granite-gneiss dome extend this field gradient further and document pressures up to 5.4-6.2 kbar in the core of the dome (see Figure 18a,b and 24a,b; Dalstra et al., 1998). The Manyutup tonalite in the Ravensthorpe belt crystallized across a similar range in pressure of 2.5-5.9 kbar (see Figure 24a). Pressure gradients within the domes are not systematic and pressure determinations from margins can be up to 6.0 kbar and in cores are up to 6.2 kbar (Ridley et al., 1997; Dalstra et al., 1998). Consequently, pressure field gradients from both granite and metamorphic rocks, document only moderate amounts of post-peak exhumation of granite-gneiss domes, with margins exhumed ~1.2-1.4 kbar relative to the greenstones, and domal cores by up to ~2.0-2.8 kbar (Dalstra et al., 1998, 1999). This pattern shows differential exhumation of granite-gneiss domes subsequent to the peak of metamorphism amounted to only ~0.8-1.6 kbar, which is consistent with core complex development during D3 extension and insufficient for this to be caused by diapirism (Blewett and Czarnota, 2007c; Goscombe et al., 2009, 2017a; Czarnota et al., 2010a). This pattern may also be in part due to older granites in domal cores (Dalstra et al., 1998), these being emplaced at high pressures during prograde metamorphism, and lower-P marginal granites during decompression through the peak of metamorphism. Almandine garnets in metapelite schists and gneisses from pre-late basin stratigraphy of the Kalgoorlie, Kurnalpi and Burtville terranes, show typical growth zoning patterns with Fe2+ and Mg increasing and Fe/(Mg+Fe), Ca and Mn decreasing with growth (Figure 14). Idioblastic garnet overgrowths grown at the peak of metamorphism have high Fe and Ca and low Mn (Figure 19). Resorbed garnets have thin outer rims with high Mn and Fe/(Mg+Fe), and low Mg. A small number of samples show atypical zoning with decreasing Mg, Ca and Mn, and increasing Fe/(Mg+Fe) towards rims. Almandine garnets in metapelite schists in the northeast Youanmi Terrane and Murchison region, show atypical growth zoning patterns with increasing Fe2+ and Fe/(Mg+Fe), decreasing Mg and Mn and flat to decreasing Ca towards rims (Figure 14). Prograde conditions based on compositional isopleths from metapelite garnet cores in the Kalgoorlie and Kurnalpi terranes, give estimates of ~5.5-6.5 kbar at ~480-550 ºC (Table 7). Prograde PT determinations are typically slightly higher-P than peak metamorphic conditions in individual samples (Figure 19) and consistently higher as a population (Figure 18a). Metapelite garnet core isopleths in the Youanmi Terrane show similar prograde conditions of 4.0-5.6 kbar at 500-560 ºC, also typically at higher-P than peak metamorphic conditions (Table 7; Figure 18b). Prograde pressures are 0.2-2.3 kbar higher than peak metamorphic conditions in individual samples, and on average are 0.9 kbar higher in the Youanmi Terrane and 1.4 kbar higher in the Kalgoorlie-Kurnalpi terranes, indicating only moderate crustal thickening during D2 contraction. These results show maximum-P was attained prior to the thermal peak, and late-prograde heating accompanied minor decompression to peak metamorphic conditions, indicating shallow clockwise P-T loops (Figures 18a,b and 19). Similar clockwise P-T paths showing only minor burial during prograde metamorphism have previously been documented from the south Kalgoorlie Terrane (Archibald et al., 1981; Bickle and Archibald, 1984). Post-peak P-T paths are constrained by average-PT calculations in THERMOCALC using variably resorbed outer rims of garnets, matrix mineral rims, coronas and phases in reaction textures. Postpeak PT determinations are lower-P than peak metamorphic conditions, and multiple results from samples define moderate-P/T arrays tracking post-peak re-equilibration (Figures 18a,b and 19; Goscombe et al., 2009). Post-peak PT determinations from Kalgoorlie, Kurnalpi and Burtville terranes, range through 535-550 ºC and 3.4-4.9 kbar and down to 1.4 kbar at 518 ºC, starting from peak metamorphic conditions between 545-610 ºC and 5.1-6.4 kbar (Table 7). These results constrain a moderate-P/T decompressive cooling trajectory of -4.6 to -11.3 kbar/100 ºC. Postpeak P-T determinations in the Youanmi Terrane are similar to peak metamorphic conditions, and the little cooling recorded indicates low-P/T cooling paths of -1.3 to -4.0 kbar/100 ºC. Taken together, shallow clockwise P-T paths with low-P/T prograde paths during burial amounting to only ~0.2-2.3 kbar, and moderate-P/T decompressive cooling paths, are consistent with a

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protracted thermal anomaly accompanying minor crustal thickening (Thompson, 1989; Thompson et al., 1997). Decompression through the peak of metamorphism is the result of cessation of burial and D2 contraction, and destruction of compressive deviatoric stress around large magma bodies (Stüwe and Sandiford, 1994), resulting in isostatic compensation of the moderately thickened crust. These tight, shallow clockwise P-T paths, and universally low-P and high-T/depth ratios are indicative of only minor crustal thickening (e.g. Ridley, 1993; Goscombe et al., 2007, 2009), and inconsistent with crustal over-thickening and collisional models (i.e. Groves and Phillips, 1987). Sequence of mineral growth from regional-contact metamorphic parageneses is consistent with prograde heating at relatively higher pressures, decompression through the peak of metamorphism and clockwise P-T paths (Bickle and Archibald, 1984; Goscombe et al., 2007, 2009). Prograde mineral growth accompanied heterogeneous D2 strain and foliation development, and peak metamorphic mineral growth typically post-dates main phase foliations at lower-strain conditions. The typical sequence of mineral growth in M2 metapelites is early biotite-muscovite-chlorite foliations containing prograde chloritoid and staurolite, overprinted by peak metamorphic postkinematic idioblasts of garnet and biotite, followed by post-peak crystallization of late-stage andalusite and cordierite poikiloblasts. Prograde chloritoid is recognised as pre-kinematic laths that are aligned in, stretched and enveloped by the main foliation. Chloritoid laths are partially resorbed and overprinted by garnet, staurolite and biotite growth (Table 4; Archibald et al., 1981). Staurolite poikiloblasts overprint chloritoid and the main foliation, and were later over-grown by peak metamorphic garnet, and later sillimanite, andalusite and cordierite parageneses, indicating decompression through prograde and peak metamorphism (Table 4; Archibald et al., 1981; Bickle and Archibald, 1984). This sequence from prograde chloritoid to staurolite, followed by peak garnetbiotite-ilmenite±andalusite parageneses, indicates the staurolite field was traversed at mediumpressures during prograde metamorphism (Figure 19). This trajectory crossed steep-P/T reactions such as; chloritoid+chlorite = staurolite+chlorite+garnet, chloritoid+chlorite = staurolite+biotite, and chlorite+garnet = garnet+staurolite+biotite into the staurolite field, and moderate-P/T reaction garnet+staurolite+biotite = garnet+biotite+aluminosilicate to form the peak metamorphic parageneses (Figure 19; e.g. Martinez et al., 2004; Tinkham et al., 2001; Johnson et al., 2003; Gaidies et al., 2006). Prograde garnet growth also started early with porphyroblasts commonly enveloped by the main foliation. In some samples, garnet growth continued after foliation development and resulted in peak metamorphic idioblastic garnet overgrowths (Figure 19), formed at the same time as post-kinematic biotite and ilmenite plates. Garnet margins are commonly resorbed by post-peak mineral reactions occurring in the rock. Post-peak mineral growth followed a typical sequence, starting with large andalusite idioblasts that occlude peak metamorphic phases such as garnet, biotite, ilmenite and highly resorbed staurolite grains (e.g. Bickle and Archibald, 1984). In low-P rocks devoid of garnet, andalusite is corroded by muscovite-sillimanite-biotite aggregates and later muscovite-fibrolite. In all samples, cordierite formed late, either as coronas replacing chloritoid and andalusite, or large poikiloblasts that occlude peak metamorphic garnet, biotite, sillimanite and andalusite (Mikucki and Roberts, 2003). Latest formed are post-kinematic, randomly oriented plates of biotite, muscovite and chlorite and retrogressive muscovite and chlorite growth. The sequence from staurolite to garnet, andalusite-sillimanite and finally cordierite parageneses indicates the crossing of moderate-P/T reactions by decompression through the peak of metamorphism (Figure 19). Reactions crossed are; garnet+staurolite+biotite = garnet+biotite+aluminosilicate and garnet+biotite+aluminosilicate = garnet+cordierite+biotite+plagioclase (e.g. Tinkham et al., 2001; Johnson et al., 2003; Martinez et al., 2004; Gaidies et al., 2006). Other reaction textures indicating decompression through the peak are thin plagioclase coronas on matrix garnet in both metapelite and amphibolite rocks (e.g. Spear, 1993).

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8.4. M3 Metamorphism (2650-2665 Ma)

M3 occurred during a second period of low-P/moderate-T metamorphism restricted to smaller-scale thermal anomalies superimposed on the regional-contact M2 metamorphic pattern (Goscombe 2007, 2009). M3 parageneses have only been identified in settings associated with D3 extension; turbiditic and fluvial post-volcanic rift basins in the EYC, post-volcanic clastic sequences in the WYC and hanging walls of crustal-scale extensional shear zones (Figure 8). These tectono-stratigraphic overprinting criteria and fortuitous deposition of late-stage sedimentary sequences after the peak of M2 metamorphism, have allowed recognition of the M3 thermal anomaly. This secondary metamorphic anomaly produced mineral parageneses in the underlying volcanic stratigraphy of similar metamorphic grade indistinguishable from the matrix forming M2 assemblages. As a consequence, only a minimum estimate of the scale and distribution of M3 thermal anomalies can be documented. These thermal anomalies are widely distributed, elongate and arcuate domains of moderate-scales between 10-50 km wide and 40-200 km long, centred on erosional/structural remnants of late basins, and arcuate bands of ~10-20 km width centred on crustal-scale extensional shear zones (Figure 8). The true width of these anomalies is probably wider and others may go unrecognised where late basin sequences have been eroded or detailed kinematic analysis is absent. Despite similar lower- to middle-amphibolite facies conditions and low-P/moderate-T thermal regimes, the contrast with extensive regional-contact patterns of M2 metamorphism is categorical. In addition to a spatially restricted pattern of metamorphism, P-T paths are anticlockwise, consistent with extension, and overprinting criteria and direct age dating show M3 metamorphism post-dates the peak of M2 and regional-contact metamorphic patterns. On these criteria, almost all expressions of M3 metamorphism are restricted to the highly extended Kalgoorlie and Kurnalpi terranes. A small number of localities in the Youanmi and Southwest terranes preserve evidence for anticlockwise P-T paths, M3 metamorphic age determinations and late clastic sequences, suggesting D3 extension may have propagated across the entire craton. Early parageneses and evidence for anticlockwise P-T paths has been obliterated in the Southwest Terrane due to whole-scale recrystallization during subsequent high-grade M4 metamorphism. The Diemal Formation late basin and most M3 localities are sited in the hanging wall of the Youanmi-Cocanarup shear zone system, consistent with extensional reactivation, suggesting this structure was the probable focus of D3 extension in the WYC (Figure 8). Extensional fault offset of stratigraphy and extensional kinematic indicators are concentrated along this shear system in the Southern Cross and Cheritons region (Bagas, 1994). Consequently, along with the Ida, Ockerburry, Keith-Kilkenny, Kunanalling, Ida, Hootanui and Laverton shear zones in the EYC, these crustal-scale shear systems partitioned D3 extensional strain and were the focus of most M3 thermal anomalies (Williams and Currie, 1993; Williams and Whitaker, 1993; Blewett and Czarnota, 2005, 2007c; Czarnota et al., 20010a,b). The superimposed M3 thermal anomaly started with the inception of lithospheric extension and development of post-volcanic rift basins, constrained to ~2665-2669 Ma in the EYC (Krapez et al., 2000; Squire, 2006, 2007). Detrital zircon maximum deposition ages from late clastic sequences in the WYC constrain a similar age range for initiation of extension between ~2663 and ~2675 Ma (Table 1; Bagas, 1999; Chen, et al., 2003; Sircombe et al., 2007; Rasmussen et al., 2010). Direct dating of M3 metamorphic minerals constrain peak metamorphism at 2656±5 Ma (Section 9.5), ~15 m.y. after the peak of M2 metamorphism, and consistent with the ~10-15 m.y. delay in heat conduction after the start of extension that is typical of rift settings (Mawby et al., 1999; Foden et al., 2006; Bleeker, 2003). This age range overlaps with the peak of late high-Ca granites involving a greater component of crustal melt, and mafic granite and syenites derived by partial melting of a metasomatized mantle (Champion et al., 2001; Champion and Cassidy, 2007). 8.4.1. Kalgoorlie and Kurnalpi Terranes Post-volcanic late basins in the EYC where metamorphosed at lower- to middle amphibolite facies conditions (Krapez et al., 2000; Painter and Groenewald, 2001; Blewett and Czarnota, 2005). The largest of these basins, the turbiditic Belches Formation, experienced high vertical flattening strains

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early in its history, producing a strong penetrative bedding-parallel schistosity that was over-grown by garnet, staurolite, biotite and andalusite porphyroblasts in static conditions at the peak of metamorphism. Peak metamorphic conditions were determined by average-PT calculations in THERMOCALC using garnet rims and matrix mineral cores (Table 7). The pooled average of n=18 peak metamorphic determination is 569±19 ºC and 3.9±0.5 kbar, indicating a high T/depth ratio of 42.2±6.2 ºC/km (Table 2). Small almandine garnets in metapelite from the Belches Formation show typical growth zoning patterns of increasing Fe2+ and Mg, decreasing Fe/(Mg+Fe) and Mn and decreasing to flat Ca towards rims (Figure 14). Most garnets are idiomorphic without resorption or corrosion of the margin, though half of the samples show very thin (<20 ) outer rims of low-Mg and high-Mn. Prograde conditions based on compositional isopleths from metapelite garnet cores, give estimates between 1.7 and 3.5 kbar at ~513-565 ºC (Table 7). All estimates of prograde conditions are at lower-pressures (P~0.4-2.2 kbar) and similar temperatures to the peak of metamorphism, indicating isothermal burial during prograde metamorphism (Figures 20 and 21a). Post-peak conditions were determined by THERMOCALC average-PT calculations using thin reequilibrated garnet edges and matrix mineral rims. Results range 500-568 ºC and 2.4-5.0 kbar, overlapping substantially with peak metamorphic conditions and at slightly lower-T, with pressures consistently higher than prograde conditions (Figures 20 and 21a). The sequence of mineral growth developed in Belches Formation metapelites is consistent with the anticlockwise P-T path outlined by prograde, peak and post-peak PT determinations (Figure 20). The bedding-parallel foliation is over-grown by flattened cordierite poikiloblasts, all of which formed early and were later overprinted by post-kinematic ilmenite, biotite, garnet, andalusite and staurolite porphyroblasts. Andalusite is also formed relatively early and was occluded by staurolite poikiloblasts and overprinted by garnet. Staurolite, biotite and garnet porphyroblasts are idiomorphic, late-formed, entirely post-kinematic and constitute the peak metamorphic assemblage. This ubiquitous sequence of mineral growth from early cordierite to andalusite and finally garnet, staurolite and biotite, indicate steep burial trajectories across moderate-P/T reactions such as: garnet + cordierite + biotite + plagioclase = garnet + biotite + andalusite, and garnet + biotite + andalusite = garnet + staurolite + biotite (e.g. Tinkham et al., 2001; Johnson et al., 2003; Martinez et al., 2004; Gaidies et al., 2006). Corona, retrogression and pseudomorph assemblages indicate postpeak cooling across steep muscovite- and chlorite-forming reactions (Figure 20). Near-isothermal burial paths culminating in static peak metamorphism, terminated by relatively shallow (low-P/T) cooling paths, is consistent with the metamorphic evolution of rift basins (e.g. Mawby et al., 1999; Webb, 2001; Alias et al., 2002). Steep burial trajectories also suggest high sedimentation rates typical of turbiditic basins (Foden et al., 2006; Bleeker, 2003). Thermal diffusion rates are high in extensional settings and culmination in peak metamorphism occurring within 10-15 m.y. of initiation of extension is typical of turbiditic rift basins (Mawby et al., 1999; Foden et al., 2006; Bleeker, 2003). Metamorphic parageneses consistent with extension have been documented in the hanging wall of the crustal-scale Ockerburry and Keith-Kilkenny extensional shear zones in the Kurnalpi Terrane (Figure 8). Sequence of mineral growth constrained by aluminosilicate relationships in the hanging wall of the Ockerburry shear zone show anticlockwise P-T paths (Goscombe et al., 2007, 2009). Local metamorphic field gradients across the Ockerburry and Keith-Kilkenny shear zones show a systematic decrease in temperature into the hanging wall, also consistent with extensional telescoping (Williams and Currie, 1993; Williams and Whitaker, 1993). Regional metamorphic field gradients show significant decreases in metamorphic grade into the hanging walls across all of the major extensional structures in the EYC; Ockerburry, Keith-Kilkenny, Kunanalling, Ida, Hootanui and Laverton shear zones (Figure 22). Peak metamorphic conditions based on THERMOCALC average-PT calculations from three samples showing anticlockwise P-T evolutions near the Ockerburry shear zone, give a pooled average of 539±35 ºC and 3.4±0.7 kbar, indicating a high T/depth ratio of 45.9±7.8 ºC/km (Table 2). Additional samples along the length of the Ockerburry shear zone give a broad spread of THERMOCALC, geothermobarometry and phase stability results range 480-610 ºC and 3.2-4.8 kbar, overlapping with conditions in the Belches Formation (Figure

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21b). The spread in temperature estimates at pressures centred around ~4.0±0.6 kbar indicate lowP/T cooling paths. Small almandine garnet porphyroblast from the hanging wall of the Ockerburry shear zone, show typical growth-zoning patterns of increasing Fe2+ and Mg, and decreasing Mn, Ca and Fe/(Mg+Fe) towards the rims, without development of resorbed or reequilibrated outer rims. Prograde conditions based on compositional isopleths from garnet cores range ~2.0-2.4 kbar at ~490-535 ºC (Table 7; Figure 21b), and are at significantly lower-pressures (P~1.4-2.4 kbar) than the peak of metamorphism, indicating isothermal burial (Figure 21b). The sequence of mineral growth in hanging wall aluminous schists at Mount Martin, Penny Dam, Kyanite Hill (Purvis, 1978, 1984), Gordon Mine and New Celebration Mine (Hodge, 2007), are consistent with anticlockwise P-T paths (Figure 21b). Pre- to syn-kinematic skeletal and poikiloblastic andalusite was formed early and accompanied main phase extensional foliation development. Andalusite was deformed, stretched and boudinaged by ongoing progressive shear and overprinted by sillimanite growth in the enveloping main foliation. Late-kinematic idiomorphic kyanite and chloritoid grew across the main foliation (Purvis, 1984; Mikucki and Roberts, 2003). A second generation of andalusite forms post-kinematic poikiloblasts and fine-grained skeletal beards on kyanite. This sequence of mineral growth around the aluminosilicate triple junction, from early andalusite to sillimanite, kyanite±chloritoid and back into the andalusite field, constrains an anticlockwise loop with decompressive cooling, showing similar relationships to those described from low-P aluminous schists during extensional events in the Mount Isa Province (Rubenach, 1992). Late-kinematic extensional shearbands indicate ongoing extensional shear and develop second-generation sillimanite and kyanite that overprint primary kyanite and andalusite porphyroblasts. These late shearband parageneses indicate ongoing extension and multiple cycles of burial and heating complicated the P-T evolution as the shear zone evolved and stepped down into the footwall. Metapelite schists in the hanging wall of the Bardoc shear zone have relationships similar to the Belches Formation of early cordierite, peak metamorphic staurolite and garnet and plagioclase coronas, indicating an anticlockwise P-T path. 8.4.2. Youanmi Terrane Relict metamorphic parageneses and PT determinations showing anticlockwise P-T paths, consistent with extension, have been documented in the Youanmi Terrane from; Diemal Formation late basin, and scattered localities in the Murchison, Gum Creek, Southern Cross and Ravensthorpe belts (Figures 8 and 21c). Diemal Formation was metamorphosed at greenschist to lower-amphibolite facies conditions and produced an early bedding-parallel flattening foliation similar to the Belches Formation late basin. All cordierite growth was early, forming flattened porphyroblasts aligned within the main foliation, and overprinted by large post-kinematic andalusite crystals and finally chlorite plates. These relationships are consistent with prograde burial to peak metamorphic conditions estimated at ~550 ºC and ~2.4 kbar, followed by shallow cooling (Figure 21c). Similar mineral relationships; from early cordierite overprinted by peak metamorphic andalusite or garnet occurs in low-P metapelite schists in the Murchison belt, and higher-P metapelite in the Gum Creek belt (Table 5). Prograde, peak and post-peak PT determinations are consistent with anticlockwise PT paths in these samples (Figure 21c). Peak metamorphic conditions based on average-PT calculations in THERMOCALC average 570±70 ºC at 2.4±1.1 kbar in the Murchison belt, and are 612±14 at 5.6±1.4 kbar in the Gum Creek belt (Table 2). Small almandine garnets in these rocks show typical growth zoning patterns of increasing Fe2+ and Fe/(Mg+Fe) and decreasing Mn and Ca towards rims. Prograde conditions based on garnet core compositional isopleths in the Gum Creek belt are ~3.2-3.6 kbar at ~525-560 ºC (Table 7), at pressures substantially lower (P~2.0-2.4 kbar) than peak metamorphic conditions. Post-peak PT determinations from the Murchison belt are ~529 ºC at 2.2 kbar, and ~475 ºC at ~5.0 kbar in the Gum Creek belt, defining shallow-P/T decompressive cooling paths between -0.4 and -1.1 kbar/100 ºC (Table 7; Figure 21c). A small number of samples in the Southern Cross belt give prograde, peak and post-peak PT determinations consistent with anticlockwise P-T paths (Figure 8). In these, prograde conditions are estimated at ~2.0 kbar and ~500 ºC, peak metamorphism between 540-560 ºC and 3.0-4.5 kbar and post-peak PT

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determinations between 2.8-4.8 kbar and 475-520 ºC, indicating cooling trajectories between -0.2 and -0.8 kbar/100 ºC. A small number of metapelite schists from the Ravensthorpe belt preserve relict M3 parageneses indicative of early low-P foliations, steep prograde loading and anticlockwise P-T paths through the peak of metamorphism. Aluminous metapelite from Annabelle volcanics in the northwest part of the belt; grew early cordierite overprinted by syn-kinematic andalusite, and post-kinematic fibrolitebiotite-plagioclase or garnet-biotite-plagioclase peak metamorphic assemblages. Typical metapelite from Chester Formation (Witt, 1998) in the northeast part of the belt, show similar prograde burial trajectories (Figure 23). Main phase foliations of muscovite-chlorite-plagioclase-rutile-chloritoid formed at very low-pressures of ~2.0 kbar below ~500 ºC (Tinkham and Ghent, 2005). Chloritoid formed early and was overprinted in a sequence from syn-kinematic andalusite to post-kinematic biotite-staurolite-muscovite, and finally garnet growth that overprinted staurolite and andalusite. This sequence of mineral growth indicates crossing of steep chlorite- and chloritoid-consuming reactions, followed by steep burial across moderate-P/T reactions such as: chloritoid + chlorite = staurolite + chlorite, staurolite + chlorite + muscovite + plagioclase = garnet + biotite + andalusite, garnet + biotite + aluminosilicate + plagioclase = garnet + staurolite + biotite, staurolite + biotite + aluminosilicate = muscovite + biotite + staurolite, and muscovite + staurolite + biotite + plagioclase = garnet + muscovite + staurolite + biotite (e.g. Tinkham and Ghent, 2005; Johnson et al., 2003). Prograde, peak and post-peak PT determinations are consistent with the anticlockwise P-T paths documented by sequence of mineral growth in the Ravensthorpe samples (Figures 21c and 23). Peak metamorphic conditions based on average-PT calculations in THERMOCALC, give a pooled average from n=8 determination of 571±27 ºC and 4.6±0.6 kbar, indicating a high T/depth ratio of 36.0±6.6 ºC/km (Table 2). PT determinations span a large range in pressures from 3.9 to 5.2 kbar, reflecting the steep burial path through peak metamorphic conditions (Figure 23). Garnet porphyroblasts show typical growth-zoning pattern of increasing Fe2+, Mg and decreasing Fe/(Mg+Fe) and Mn towards the rims and both decreasing and increasing Ca patterns. Few garnets are resorbed and have thin outer rims (<20-30 ) showing lower Mg, and higher Mn and Fe/(Fe+Mg). Prograde conditions based on garnet core compositional isopleths, give estimates between 2.3 and 4.3 kbar at ~528-590 ºC (Table 7; Figure 21c), at significantly lower-pressures (P~0.3-2.3 kbar) than the peak of metamorphism, indicating near isothermal burial (Figures 21c and 23). Post-peak conditions based on THERMOCALC average-PT calculations using garnet outer rims and matrix mineral rims; range 515-565 ºC and 2.1-5.9 kbar (Table 7), overlapping substantially with peak conditions and indicating shallow cooling paths. 8.5. M4 High-Grade Metamorphism (2615-2650 Ma)

M4 metamorphism occurred during a protracted, late-stage, crustal-scale thermal anomaly that produced a broad region (~300-500 km wide) of low-strain low-P/high-T parageneses across the Narryer, Southwest and western Youanmi terranes (e.g. Bosch et al., 1996; Rennie 1998; Nemchin et al., 1994). Metamorphic grade varies from lower- to middle amphibolite facies in the western Youanmi Terrane, across steep transitional gradients to granulite facies gneisses and migmatites in the Southwest Terrane, and a sharp transition across the Yalgar shear zone to lower-grade granulites in the Narryer Terrane (Figure 9). In granulite-grade regions all earlier metamorphic parageneses were recrystallized and lost. Whereas, at amphibolite facies conditions in the western Youanmi Terrane, relict domains of earlier parageneses (i.e. M2 and M3) and regional-contact metamorphic patterns are variably preserved (Figures 8 and 9). Regional-contact metamorphic patterns are preserved in the Marda greenstone belt and remainder of the northeast Youanmi Terrane (Figure 7; Ridley et al., 1997). Consequently, the M4 regional metamorphic overprint either did not occur in the northeast Youanmi Terrane and EYC or was at too low grade to be recognised. However, patchy development of late-stage alteration and mineralization assemblages with dates between ~2610 and ~2645 Ma age (Section 8.6) are widespread in the Youanmi and Kalgoorlie terranes. This suggests that the M4 thermal anomaly had a broad low-grade margin (~180-300 km wide) extending as far

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east as the Ockerburry shear zone. M4 parageneses and P-T determinations consistently show evidence for tight, shallow (low-P/T) clockwise P-T paths with only moderate amounts of burial (e.g. Ridley et al., 1997; Dalstra et al., 1999; Goscombe et al., 2007, 2009). Evidence for steep prograde loading and anticlockwise P-T paths are conspicuously absent. These P-T paths indicate that the thermal anomaly was associated with only minor contraction and crustal thickening, probably during D4 and D5 reactivation (Section 3.4). Metamorphic age determinations from high-grade regional metamorphic parageneses indicate highgrade conditions persisted for ~36 m.y. between 2649 and 2613 Ma and involved two distinct thermal peaks at 2644±4 Ma and 2629±7 Ma (Section 9.6). These dates indicate that M4 metamorphism was late and post-dates all other metamorphic events, D3-M3 lithospheric extension, late clastic basins and high-Ca magmatism (Table 1). There is no systematic spatial variation in the age of metamorphism, indicating that these thermal pulses were of regional extent and affected the middle- to upper-crust at the same time across much of the craton. M4 metamorphism overlaps in age with large volumes of low-Ca granite of ~2615-2650 Ma age intruded craton-wide and amounting to ~30-35% of all granite in the Yilgarn craton and ~20-24% of the upper-crust (e.g. Champion et al., 2001). Consequently, the M4 thermal anomaly was responsible for extensive partial melting of the middle- to lower-crust and advective transfer of heat by plutons into the upper-crust (e.g. Champion and Sheraton, 1997; Champion and Cassidy, 2007). The regional pattern of metamorphism indicates diffusion of the thermal anomaly on a broad-scale front and was significant enough to generate whole-scale granulite facies metamorphism in the middle- to upper-crust. The high intensity of the thermal anomaly along with the craton-wide scale, suggest an asthenospheric heat source. All features of the M4 thermal anomaly: (1) low-strain, (2) late-stage, (3) protracted, (4) broad-scale, (5) high intensity, and (6) low-P/high-T thermal regimes, are consistent with resulting from delamination of eclogitic lower-crust and/or mantle lithosphere and upwelling of the asthenosphere (e.g. Smithies and Champion, 1999; Blewett, et al., 2004a, 2008; Czarnota et al., 2010a; Korsch et al., 2013). The near craton-wide extent and similar age of M4 metamorphism over such a large region, indicates that the most probable trigger for delamination was M3 lithospheric extension at ~2665-2668 Ma, which was also of craton-wide extent (Section 3.3; Blewett et al., 2008; Czarnota et al., 2010a; Korsch et al., 2013). Metamorphic age constraints indicate conduction of heat into the middle- to upper-crust took at least 15-18 m.y., followed by thermal peaks ~21-24 m.y. and ~36-39 m.y. after delamination triggered by lithospheric extension. These delays are consistent with modelling of heat diffusion through ~40-47 km of crust (e.g. Houseman et al., 1981; Campbell and Hill, 1988; Houseman and Molnar, 1997; Rey, 2006; Czarnota et al., 2010a; Ducea, 2011). 8.5.1. Youanmi Terrane Middle amphibolite facies assemblages interpreted to be M4 parageneses on the basis of direct age determinations or overprinting criteria (Sections 5.5, 6.5.1 and 9.6; Figure 9), are widespread across the western Youanmi and northeast Narryer terranes; being documented in the Ravensthorpe, Forrestania, Cheritons, Southern Cross, Bullfinch, Perenjori, Yalgoo, Dalgaranga, Murchison and Jack Hills belts (Figure 9). Peak metamorphic assemblages are characterized by well-annealed, static to moderate-strain matrix foliations with polygonal granoblastic textures and post-kinematic porphyroblastic phases. Typical metapelite assemblages are: garnet-staurolite-andalusitebiotite±sillimanite, garnet-cordierite-andalusite-biotite, and cordierite-andalusite-biotite-muscovite (Section 6.5.1; Table 4). Peak metamorphic conditions based on THERMOCALC average-PT calculations give similar pooled averages in different parts of the Youanmi Terrane: 568±42 ºC and 4.1±1.1 kbar for n=40 samples in the Cheriton-Southern Cross region, 574±44 ºC and 4.7±1.2 kbar for n=20 samples in the Forrestania and Ravensthorpe belts, and 564±40 ºC and 3.7±1.3 kbar for n=7 samples in the southwest Murchison region (Tables 2 and 7; Ridley et al., 1997; Dalstra et al., 1998). All samples record high T/depth ratios of ~38-48 ºC/km, and these thermal regimes are comparable to granulites in the Southwest Terrane ~34-46 ºC/km (n=39), and Narryer Terrane ~48

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ºC/km (n=13) (Table 2). PT determinations from the same region cluster relatively closely (Figure 24a,b), and show almost no systematic variation across individual greenstone belts (Figure 22; Dalstra et al., 1998; Ridley et al., 1997). Small almandine garnets from the Southern Cross region and Ravensthorpe belt, consistently show typical growth-zoning patterns of increasing Fe2+ and Mg, and decreasing Mn, Ca and Fe/(Mg+Fe) towards the rims, with almost no development of resorbed or re-equilibrated outer rims (Figures 14, 25 and 26: Appendix 4). Prograde conditions based on compositional isopleths from garnet cores in the Cheriton, Southern Cross and Dalgaranga belts cluster around ~4.6-5.6 kbar at ~485-570 ºC (Table 7). These pressures are typically higher (P~0.5-2.3 kbar) than the peak of metamorphism, indicating late-prograde decompressive-heating trajectories and clockwise P-T paths (Figures 24b and 26). Pressure estimates from garnet cores in the Ravensthorpe and Forrestania belts are typically ~3.0-4.3 kbar (Table 7), and lower than peak metamorphic conditions, indicating loading during prograde metamorphism to peak metamorphic conditions (T-max) at the maximum-P attained (Figures 24a and 25). Post-peak conditions in all belts, determined by THERMOCALC average-PT calculations using the outer rims of garnet and matrix mineral rims, are consistently lower (P ~0.22.1 kbar) than peak metamorphic conditions (Table 7). These constrain moderate-P/T decompressive-cooling paths of -1.2 to -11.8 kbar/100 ºC (Figures 24a,b) and average -2.5 kbar/100 ºC (n=10). Taken together these PT constraints document tight, shallow (low-P/T) clockwise P-T paths with only minor prograde burial. Similar tight clockwise P-T paths showing only minor burial and decompression through the peak have been documented from the Southern Cross belt (Ridley et al., 1997; Dalstra et al., 1999). The sequence of M4 mineral growth in metapelite from different belts in the Youanmi Terrane are consistent with the clockwise P-T paths outlined by prograde, peak and post-peak PT determinations (Figures 24a,b, 25 and 26). Prograde metamorphism accompanied deformation, main foliation development and syn-kinematic growth of garnet cores. In contrast the peak of metamorphism coincides with post-kinematic porphyroblast growth at lower-strain conditions. Sequence of mineral growth in typical metapelites from the Southern Cross, Cheriton, Dalgaranga and Ravensthorpe belts follow similar progression from; prograde chloritoid and staurolite parageneses, overprinted by peak metamorphic garnet and biotite idioblasts, in turn occluded by late-stage andalusite and cordierite poikiloblasts (Figures 25 and 26; Ridley et al., 1997). Chloritoid is earliest-formed and strongly resorbed by subsequent mineral growth, leaving little remaining within the biotite-plagioclaseilmenite±muscovite foliation or preserved within garnet porphyroblasts. Almost all staurolite occurs as small, highly resorbed relict grains aligned within the main foliation and occluded within synkinematic garnet cores, and andalusite and cordierite porphyroblasts. Peak metamorphic assemblages are represented by post-kinematic growth of idioblastic garnet rims, biotite plates, andalusite idioblasts and rare sillimanite, all of which overprint the main foliation. Garnets are overgrown by sillimanite and occluded by andalusite and cordierite but remain idiomorphic without resorption of grain margins. Cordierite is latest-formed as coronas and poikiloblasts on andalusite margins, indicating decompression through the peak of metamorphism (Ridley et al., 1997). The progression from chloritoid to staurolite and eventual garnet-andalusite-biotite±sillimanite parageneses indicate the staurolite field was traversed at medium-pressures during prograde metamorphism (Figure 26; Tinkham and Ghent, 2005; Johnson et al., 2003; Gaidies et al., 2006, 2008). The prograde path crossed steep-P/T reactions: chloritoid+chlorite = staurolite+chlorite+garnet, chlorite+garnet = garnet+staurolite+biotite, and garnet+muscovite+staurolite+biotite = garnet+biotite+andalusite±sillimanite to form the peak metamorphic parageneses. These reactions were followed by decompression across moderate-P/T reactions: sillimanite = andalusite, garnet+biotite+andalusite = garnet+cordierite+biotite, and garnet+cordierite+biotite = andalusite+cordierite+biotite+plagioclase to form post-peak parageneses, and cooling across biotite+cordierite+andalusite = chlorite+muscovite+andalusite to form retrograde parageneses (Figures 25 and 26; Tinkham and Ghent, 2005; Johnson et al., 2003; Gaidies et al., 2006, 2008).

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8.5.2. Southwest Terrane Metamorphic age determinations in the 2650-2615 Ma age range are widespread across the Southwest Terrane (Section 9.6) and the greater part of the metamorphic pattern is interpreted to be M4 (Figure 9). The central region is dominated by a broad NW-trending expanse (220 x 400 km) of low- to medium-P, high temperatures >750 ºC granulites, including the Lake Grace and Jimperding Belts and eastern region. Peak metamorphic assemblages are characterized by a coarse-grained, well-annealed polygonal granoblastic matrix, either unaligned or with a weak grain shape fabric. Typical diagnostic matrix assemblages are: garnet-cordierite metapelite granulite, orthopyroxenegarnet±cordierite granulites and garnet-free two-pyroxene mafic granulites (Section 6.5.2; Table 6). Non-diagnostic garnet-biotite felsic orthogneiss and paragneiss predominate. Graphite plates in the matrix assemblage of orthopyroxene granulites in the Lake Grace belt indicate reducing conditions. All rock types are migmatized. Partial melt pools and stromatic segregations in metapelite and felsic granulites contain garnet or cordierite porphyroblasts. Partial melt segregations in mafic granulites contain orthopyroxene, limiting metamorphic conditions to >800 ºC (Spear, 1993), and the absence of garnet limit conditions to below 7.0 kbar (e.g. Pitra et al., 2010; Ziaja et al., 2010). To the north and east are broad, patchy transition zones with mixed granulite and upper-amphibolite facies. Across the Westonia and Cocanarup shear zone systems are relatively steep transitions to middle amphibolite facies in the Youanmi Terrane. The western margin including the Boddington greenstone belt, has a relatively steep transition through upper- to middle- and low-amphibolite facies, with isolated patches of low-P mafic granulite (~3.5 kbar and ~798 ºC) along the northwest margin (Table 2). In the southwest, the Balingup domain has been reworked substantially in the Cryogenian due to collisional events in the Pinjarra Orogen, and consequently original Neoarchaean metamorphic conditions cannot be confidently characterized. Peak metamorphic conditions based on THERMOCALC average-PT calculations and phase stability constraints, give similar pooled averages for granulites in different parts of the Southwest Terrane: 772±27 ºC and 4.9±0.8 kbar in the eastern region (n=15), 801±34 ºC and 6.2±1.1 kbar in the Lake Grace belt (n=11), and 807±35 ºC and 6.0±1.2 kbar in the Jimperding belt (n=13) (Tables 2 and 7). All samples record high T/depth ratios of ~38-46 ºC/km, comparable to the Narryer Terrane ~48 ºC/km (n=13) and Youanmi Terrane ~38-48 ºC/km (n=67) (Table 2). Peak metamorphic PT determinations from each region cluster relatively closely (Figure 27b,c). These regional populations and multiple PT determinations from individual samples, both show steep-P/T arrays of ~3.3-20.0 kbar/100 ºC and 4.4-42.5 kbar/100 ºC respectively, indicating decompression through the peak of metamorphism. There is near systematic variation in metamorphic conditions along the NW-SE length of the central granulite region (Figures 22 and 27b,c). Temperatures are highest in the Jimperding belt in the northwest (~800-850 ºC), lowest in Lake Grace belt in the centre (~750-820 ºC) and moderate in the southeast region (~770-830 ºC). Pressures decrease from ~5.0-7.5 kbar in the northwest to ~4.0-5.8 kbar in the southeast. Large Mg-almandine garnet porphyroblasts from both metapelite and orthopyroxene granulites preserve little of the original growth zoning, most being homogenized to flat compositional patterns. Garnet cores that do preserve weak zoning have increasing Fe2+ and Fe/(Mg+Fe) and decreasing Mg and Ca towards rims, with Ca and Mn most typically showing no zoning. Almost all garnets are strongly corroded and resorbed (typically by cordierite), leaving irregular shaped margins and broad re-equilibrated rims with significantly lower Fe and Mg, and higher Mn (Figures 14 and 28; Appendix 4). Prograde conditions based on compositional isopleths from metapelite garnet cores that preserve primary growth zoning, give estimates between ~6.0-8.4 kbar (Table 7). These pressures are slightly higher (P~0.0-2.8 kbar) than the peak of metamorphism, indicating isobaric to decompressive heating prograde paths (Figures 27b,c). These peak-P conditions constrain a lowP/T clockwise P-T path and indicate only moderate burial and crustal thickening during M4 (Figure 28). Post-peak conditions based on average-PT calculations using resorbed margins of garnet and matrix mineral rims, range between ~3.3-7.8 kbar, and at pressures lower (P +0.3 to -3.1 kbar) than the peak of metamorphism (Table 7). Numerous pairs (n=17) of peak and post-peak PT

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determinations constrain isobaric to low-P/T decompressive-cooling paths of +0.5 to -3.1 kbar/100 ºC (Figures 24a,b) and averaging -1.0 kbar/100 ºC. The sequence of mineral growth in metapelite and orthopyroxene granulites from different parts of the Southwest Terrane are consistent with the clockwise P-T paths outlined by prograde, peak and post-peak PT determinations (Figures 27b,c, 28 and 29). Mineral growth follows a similar sequence in most of these samples (Section 6.5.2; Table 6). Prograde parageneses are characterized by synkinematic garnet porphyroblast cores with inclusions of sillimanite, plagioclase, ilmenite and biotite, and absence of early cordierite and orthopyroxene. Early garnet-sillimanite parageneses are overgrown by peak metamorphic matrix assemblages of garnet-cordierite-melt or garnetorthopyroxene-melt±cordierite. Peak metamorphic assemblages are overprinted by orthopyroxene poikiloblasts, secondary biotite, orthopyroxene and cordierite in coronas and symplectites and latestage fabrics, all grown at the expense of corroded and resorbed garnet. Latest-stage mineral growth is secondary plagioclase and alkali feldspar in coronas and exsolution lamellae and retrogressive chlorite and muscovite. The progression from prograde garnet-sillimanite to peak garnetcordierite±orthopyroxene and post-peak cordierite±orthopyroxene and plagioclase parageneses, indicate decompression across lowto moderate-P/T reactions: garnet+sillimanite+biotite+plagioclase = garnet+sillimanite+cordierite+melt, garnet+sillimanite+cordierite+biotite = garnet+cordierite+melt, and garnet+cordierite+biotite = garnet+cordierite+orthopyroxene+melt, to form post-peak parageneses. This was followed by decompressive cooling across steep-P/T reactions: garnet+cordierite+orthopyroxene = cordierite+orthopyroxene+plagioclase, cordierite+orthopyroxene+melt = cordierite+biotite+plagioclase, and garnet+cordierite+melt = cordierite+biotite+ plagioclase to form retrograde parageneses (Figures 28 and 29; e.g. White et al., 2001; Kelsey, 2008; Longridge et al., 2017). 8.5.3. Narryer Terrane The greater part of the Narryer Terrane is composed of low-P granulite facies gneisses formed at moderately high temperatures of ~725-800 ºC. Upper-amphibolite facies gneisses dominate the eastern and western margins of the terrane (Figure 9). There is a sharp drop in metamorphic grade across the Yalgar shear zone into the Youanmi Terrane. Metamorphic age determinations from the granulites indicate peak metamorphism between 2620 and 2659 Ma and overlap in age with M3-M4 metamorphic events elsewhere (Section 9.6; Figure 32). Lower-amphibolite facies parageneses were developed in widespread retrograde shear zones (m- to km-scale), and in the northeast region the Jack Hills belt was reworked at middle-amphibolite conditions and developed clockwise P-T paths (Occhipinti et al., 1998). These lower-grade retrogressive parageneses formed during reactivation, reworking and widespread thermal overprint along the northern margin of the Narryer Terrane during the Palaeoproterozoic Capricorn Orogeny (Occhipinti et al., 1998). Middle-amphibolite parageneses in the Jack Hills belt yield U-Pb monazite ages of ~1820±20 Ma and 1856±6 Ma (Rasmussen et al., 2010), and Ar-Ar muscovite ages range ~1776-1827 Ma and ~1736-1760 Ma (Spaggiari, 2007; Spaggiari et al., 2008). Metapelite and meta-psammopelite granulites develop a well-annealed polygonal granoblastic matrix with moderate grain shape fabric of elongate biotite, quartz, plagioclase and cordierite. Matrix assemblages are garnet-cordierite-biotite-quartz-plagioclase, and K-feldspar and partial melt segregations are uncommon (Section 6.5.3; Table 6). Peak metamorphic PT determinations from metapelite granulites, based on THERMOCALC average-PT calculations and phase stability constraints, give a pooled average of 758±37 ºC and 4.6±0.7 kbar (n=13), indicating high T/depth ratios of 48±9 ºC/km (Tables 2 and 7). These results range between ~730-800 ºC and ~4.1-6.0 kbar and are consistent with the stability field of the garnet-cordierite-biotite-plagioclase-K-feldspar matrix assemblage (Figure 30; e.g. White et al., 2001; Tinkham et al., 2001; Storm and Spear, 2005). Peak metamorphic PT determinations cluster along a moderate-P/T array indicating decompressive cooling through the peak at ~2.7 kbar/100 ºC (Figure 27a). A subset of average-PT

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determinations from metapelite granulites is between 650-700 ºC and 4.1-5.0 kbar (Table 7; Figure 27a). These are ~30-80 ºC lower than peak metamorphic conditions and only partially overlap with the stability field of the matrix assemblage (Figure 30), indicating re-equilibration of these samples during cooling (Frost and Chako, 1989; Goscombe et al., 1998; Kohn and Spear, 2000). Almandine garnet porphyroblasts from metapelite granulites preserve little growth zoning, most being homogenized to flat compositional patterns. Garnet cores that do preserve weak zoning have increasing Fe2+ and Fe/(Mg+Fe), and decreasing Mg and Ca towards rims, with Ca and Mn most typically showing no zoning. Garnet margins are variably resorbed and develop thin re-equilibrated rims with low-Mg and high-Mn where corroded by cordierite (Figure 14; Appendix 4). Prograde conditions based on compositional isopleths from metapelite garnet cores that still preserve growth zoning, give pressure estimates between ~4.0-7.3 kbar (Table 7). These pressures are higher (P~0.1-3.0 kbar) than the peak of metamorphism, indicating decompression during late-prograde metamorphism (Figures 27a). The peak-P conditions constrain a shallow (low-P/T) clockwise PT path and indicate only moderate burial and crustal thickening during M4 (Figure 30). Post-peak average-PT calculations using garnet outer rims and matrix mineral rims range between ~2.6-4.8 kbar, and pressures are typically lower (P +0.4 to -2.6 kbar) than peak metamorphic conditions (Table 7; Figure 30). Peak and post-peak PT determination pairs (n=10) constrain isobaric to lowP/T decompressive-cooling paths of +0.3 to -1.3 kbar/100 ºC (Figures 27a) and average -0.7 kbar/100 ºC. The sequence of mineral growth in metapelite granulites from the Narryer Terrane is consistent with the tight clockwise P-T paths outlined by prograde, peak and post-peak PT determinations (Figures 27a and 30). Mineral growth follows a similar sequence in most of these samples (Section 6.5.3; Table 6). Prograde parageneses are characterized by growth of garnet porphyroblast (now cores) and early sillimanite and biotite preserved as inclusions within garnet and cordierite porphyroblasts. Sillimanite is absent from the peak metamorphic matrix assemblage of garnet-cordierite-plagioclasebiotite±K-feldspar. Cordierite growth outlasted garnet porphyroblasts and secondary cordierite±biotite developed as moats and embayments that corroded garnet margins. Cordierite is partially replaced by late-stage biotite-plagioclase-sillimanite, and garnet margins have secondary sillimanite growth and plagioclase coronas. This sequence of mineral parageneses indicate decompression across low- to moderate-P/T reactions: garnet+sillimanite+biotite+K-feldspar = garnet+sillimanite+biotite+melt, and garnet+sillimanite+biotite = garnet+cordierite+biotite+Kfeldspar to form peak parageneses. These reactions were followed by cooling across low-P/T reactions: garnet+cordierite+biotite = cordierite+biotite+plagioclase, and cordierite+biotite+garnet = cordierite+biotite+sillimanite+plagioclase to form post-peak parageneses (Figure 30; e.g. White et al., 2001; Tinkham et al., 2001; Storm and Spear, 2005). 8.6. M4 Alteration and Mineralization (2610-2655 Ma)

Late-stage mineral assemblages associated with low-grade alteration and mineralization, are developed in relatively small patchy domains (~2-20 km) scattered across the Youanmi, Kalgoorlie and Kurnalpi terranes (Figure 31; Witt, 1993; Mikucki and Roberts, 2003). The two youngest alteration-mineralization events, of 2640-2655 Ma (Au3) and 2610-2640 Ma (Au4) age, overlap in age and distribution with the near craton-wide M4 thermal anomaly (Section 9.7). Mineralization and alteration younger than 2645 Ma is restricted to west of the Ockerburry shear zone, which is interpreted to be the eastern limit of the M4 thermal anomaly (Figure 4 and 32). The M4 thermal anomaly may have extended further east into the Kurnalpi Terrane, but at grades too low to drive appreciable alteration reactions and mineralization. A large dataset of PT determinations from alteration and mineralization parageneses in the Youanmi, Kalgoorlie and Kurnalpi terranes (n=370), based on geothermobarometry and fluid inclusion studies, was compiled from the literature (Appendix 1) (e.g. Ho et al., 1987, 1990; Neall and Phillips, 1987; Wang et al., 1993; Bloem, 1994; Bloem et al., 1994; Dalstra, 1995; Knight et al., 1996, 2000; Witt,

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1998; Mikucki, 1997; Cassidy et al., 1998; Dugdale and Hagemann, 2001; Vielreicher et al., 2002; White et al., 2003; Hagemann and Luders, 2003; Mueller et al., 2004; Salier et al., 2005; Evans et al., 2006; and others). These PT determinations define a broad linear array from: 450-550 ºC and 2.4-4.2 kbar (overlapping with M3 and M4 conditions), to 220-320 ºC and 1.4-3.0 kbar, and a lowPT at 175 ºC and 1.0-1.8 kbar (Figure 3; Goscombe et al., 2009). This PT-array has a low-P/T slope between -0.5 and -0.8 kbar/100 ºC (Figure 3), interpreted to represent a shallow decompressive cooling path for Youanmi, Kalgoorlie and Kurnalpi terranes from ~2645 to ~2610 Ma. These regions experienced a total of ~1.4-2.4 kbar of exhumation during this period, amounting to ~4.9-8.4 km of erosion. The greater part of this exhumation occurred during decompression through, or immediately after, the peak of regional metamorphic events. Highest-grade alteration and mineralization assemblages are contemporaneous with M3 metamorphism in the EYC, and M4 metamorphism in the Youanmi Terrane. These alteration assemblages formed at pressures ~0.1-1.4 kbar lower than the peak of regional M3 and M4 events, which is consistent with decompression during or immediately after the peak. A second exhumation event of ~0.8 kbar occurred late in the Au4-mineralization period, stepping from alteration conditions at ~1.4-3.0 kbar, to lowest-grade alteration at ~1.0-1.8 kbar (e.g. Ho et al., 1987, 1990; Knight et al., 1996; Cassidy et al., 1998; Dugdale and Hagemann, 2001; Salier et al., 2005). These final depths are comparable with Pestimates of ~2 kbar from late-stage granites using Al-in-hornblende geobarometers (Figure 18). This youngest-phase of exhumation and mineralization is associated with D5 dextral transtension between ~2620-2645 Ma (Blewett and Czarnota, 2005, 2007c). Average thermal gradients associated with alteration and mineralization assemblages were high throughout this whole period and range 28-54 ºC/km. For typical fluid-to-rock ratios and fluid flow rates, both fluid and host rock remain thermally equilibrated, and any thermal difference was transient and short lived. Furthermore, alteration assemblages vary on scales too small to have experienced significant differences in T and P, and thus were primarily controlled by fluid chemistry, pH and redox (e.g. Mikucki, 1997; Neumayr et al., 2007; Walshe et al., 2008a,b). Similarly, pressures recorded by fluid inclusions are continuously equilibrated with changing conditions and represent minimum estimates of crustal depth at entrapment (Hagemann and Luders, 2003). As a result, PT determinations from alteration and mineralization assemblages approximately represent background regional conditions at the time, within error of the different thermobarometric methods. High, background average thermal gradients were maintained throughout this whole period by a combination of: [1] multiple and protracted near craton-wide M4 thermal pulses that dissipated slowly (Section 8.5), and [2] advection of heat into the upper-crust by voluminous low-Ca granite magmatism (~20-25% of crustal volume). The patchy distribution of alteration and mineralization, therefore, does not reflect small-scale thermal anomalies, but is primarily due to high fluid-to-rock ratios superimposed on background conditions undergoing shallow decompressive cooling. Local, transient thermal pulses following isobaric heating and cooling paths are expected only in aureoles associated with low-Ca granite, mafic granitoids and syenite intrusions. Furthermore, the diversity of mineralization and alteration assemblages is primarily dependent on fluid chemistry and mixing of fluids from different sources. At least four fluid sources have been identified in association with alteration and mineralization: [1] saline, reduced-acid H2O fluids dewatered from post-volcanic turbiditic late basins, [2] reduced-acid H2O-CO2 fluids from prograde dehydration reactions in late basins and low-grade volcanosedimentary rocks, [3] large volumes of dry, oxidized-alkaline CO2-SO2 magmatic fluids exsolved from crystallizing granites, and [4] dry, reduced-alkaline H2-H2S-CH4 mantle-derived fluids associated with mafic granite, syenite and carbonic magmas (e.g. Mikucki, 1997; Neumayr et al., 2007; Sheldon et al., 2008; Walshe et al., 2008a,b). 9. GEOCHRONOLOGY OF METAMORPHIC EVENTS

Metamorphic age determinations undertaken for this project (n=52) include a range of different radiometric systems and mineral chronometers on parageneses representing each of the different

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metamorphic events (Tables 1 and 2). Sm-Nd and Lu-Hf garnet-whole rock isochron ages were determined on M1 amphibolites and M3 metapelites. Pb207-Pb206 weighted mean concordia ages were determined in situ from monazite in M3 and M4 metapelite parageneses using LA-ICPMS. Pb207-Pb206 isochron ages were determined in situ from titanites in M1, M2 and M4 amphibolite assemblages using MC-LA-ICPMS. Pb207-Pb206 weighted mean concordia ages were determined from zircon grains separated from M4 metapelite granulites using MC-LA-ICPMS and SHRIMP. The analytical methods and age determinations are summarized below, and full sample and analytical details are available in Appendices (1 and 2) and Goscombe et al. (2009). These data are augmented by metamorphic age determinations collated from literature (n=62) (Table 2; Appendix 1; GSWA geochronology database; Kinny et al., 1988; Pidgeon et al., 1990; Nemchin et al., 1994; McMillan, 1996; Wang et al., 1998; Schiotte and Campbell, 1996; Dalstra et al., 1998; Rennie, 1998; Campbell et al., 1998; Yumin et al., 1998; Nemchin and Pidgeon, 1999; Bagas, 1999; Mueller and McNaughton, 2000; Wilde, 2001; Fletcher and McNaughton, 2001; Mueller et al., 2004; Sircombe et al., 2007; Van Kranendonk and Ivanic, 2009; Muhling et al., 2008; Iizuka et al., 2008; Thebaud and Miller, 2009; Rasmussen et al., 2010). Metamorphic age data has been restricted to UPb zircon, monazite, titanite, allanite and xenotime, and Sm-Nd and Lu-Hf garnet age determinations from matrix parageneses (Figure 32; Table 1 and 2). Given the large size of the total dataset, the following summary focuses on middle Neoarchaean metamorphic events and discusses pooled ages from a number of sources (Table 2; Figure 32), with the details of specific age determinations available in Appendices (1 and 2). 9.1. Analytical Procedures

9.1.1. Lu-Hf and Sm-Nd Garnet Chronology Lu-Hf and Sm-Nd isotope analyses were undertaken at Melbourne University following the method of Maas et al. (2005). Garnet separates were prepared using heavy liquids, magnetic separator and handpicking. Garnet-poor rock matrix was powdered using an agate mill and two splits were dissolved at high pressure and low pressure. Each split was spiked with mixed 176Lu/180Hf and 149Sm-150Nd spikes. Garnet fractions were ground in an agate mortar and leached with hot 6M HCl for 2 hours to remove phosphates and raise the Sm/Nd ratio of the garnet residue. Garnets, garnet residues and leachates were spiked with high Lu/Hf ratio spike and dissolved on a hotplate (HFHNO3, HCl). LREE were extracted from all samples using 0.1 ml beds of EICHROM RE resin. Lu and Hf were extracted using a single pass over 1 ml columns of EICHROM LN resin. Sm and Nd were purified on different columns of the same resin. Isotopic analyses were carried out on a NU Plasma MC-ICPMS (Maas et al., 2005). Hf isotope ratios were corrected for interfering Lu, Yb and W, normalized to 179Hf/177Hf=0.7325 and reported relative to JMC475=0.282160. Spiked Lu fractions are corrected for Yb and mass bias based on the measured 173Yb/171Yb. Nd mass bias is corrected by normalizing to 146Nd/145Nd=2.0719425 and reported relative to La Jolla=0.511860. Accuracy was assessed through isotope dilution analysis of USGS standard basalts with both Lu-Hf and Sm-Nd spikes, with results consistent with other laboratories. Age calculations used the decay constant for 176Lu of 1.865 10-11/yr (Scherer et al., 2001), and 147Sm is 6.54 10-12/yr (Dickin, 1995). Age determinations are based on the 2 whole rock splits and variously treated garnet. 9.1.2. U-Pb Monazite Chronology Pb isotope analyses of monazite were undertaken at Adelaide University following the method of Payne et al (2008). U-Pb analyses were carried out on a UP-213 New Wave Nd-YAG laser attached to an Agilent 7500cs ICP-MS. Ablation was conducted in a helium atmosphere and argon gas added immediately to aid transport of material. Spot size was 12-15 micron, with a laser frequency of 4 Hz resulting in an average fluence of 6 J/cm2. A single analytical spot consisted of a 50s gas blank followed by 40s of data acquisition with the laser firing. Measured isotopes were 204Pb, 206Pb, 207Pb, and 238U with dwell times of 10, 15, 30, and 15 ms respectively. Age calculations were carried out using the data reduction software Glitter (Jackson et al. 2004). Down-hole element fractionation was corrected in the Glitter software via the use of the external monazite standard Madel. Accuracy

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was verified using an in-house monazite standard 94-222/Bruna-NW (Payne et al. 2008) and the distributed monazite standard 44069 (Aleinkoff et al. 2006). Due to the unresolvable 204Hg on 204Pb interference, uncorrected isotope ratios are used for all reported age calculations, with concordia plots generated using Isoplot/Ex 3.7 (Ludwig, 2008). Intercept ages were calculated via linear TeraWasserburg regression (Tera and Wasserburg, 1972), with decay-constant errors propagated. Weighted average age calculations (at 95% confidence) have been performed using uncorrected 206Pb/238U ratios. 9.1.3. U-Pb Titanite Chronology Pb isotope analyses of titanite were undertaken via LA-ICP-MS at the University of Florida using a Nu-Plasma multi-collector instrument with a U-Pb collector block and a New Wave 213 nm Nd-Yag laser (Kamenov et al., 2004). Titanite analyses were standardized against fused and homogenized Bancroft titanite glass and crystals calibrated by isotope dilution analyses of Pb isotopes via TIMS and ICP-MS (Coyner et al., 2005). Analyses of the standard were made between every three unknown analyses. In all samples, titanite was measured in situ on polished thin sections with the reference titanite attached to the thin section. Dates (±2) from titanite analyses were calculated from 207Pb/204Pb versus 206Pb/204Pb isochrons. 9.1.4. U-Pb Zircon Chronology U-Pb analyses of zircon were undertaken at the University of Florida following the method of Foster et al. (2009). Grain separates were hand picked, mounted in epoxy with the zircon standard (FC-1), and ground and polished to reveal internal surfaces. Pb isotope analyses were performed using a Nu Plasma multi-collector inductively coupled plasma source mass spectrometer (MC29 ICP-MS) equipped with three ion counters and 12 Faraday detectors, U-Pb collector block and a New Wave 213 nm Nd-Yag laser (Kamenov et al., 2004). Data calibration and drift corrections were based on multiple ablations of the reference zircons from the Duluth Gabbro (Paces and Miller, 1993) collected from the Forest Center location (FC-1). Analyses of the standard were made between every three unknown analyses. Analyses were plotted on conventional concordia diagrams and cumulative density diagrams using ISOPLOT (Ludwig, 1999) to assess discordance. Discordant grains that plotted along reliable discordia were assumed to be of the upper intercept age and included in probability plots. Discordant analyses that did not intersect the concordia curve or plot along discordia were not included in age populations because of the possibility of multiple stages of Pb loss. Pb isotope analyses of zircon were also undertaken using SHRIMP II at Curtin University following the method of Williams (1998). Grain separates were hand picked, mounted in epoxy with the GSWA zircon standard, and ground and polished to reveal internal surfaces. Photomicrographs in transmitted and reflected light and SEM cathodoluminescence images, were used to decipher the internal structures of grains to target areas for spot analysis. The data have been reduced in a manner similar to that described by Williams (1998) using the software SQUID (Ludwig, 2000). For zircon calibration the 206*Pb/238U ratios have been normalised relative to a value of 0.1859 for the reference zircons (Paces and Miller, 1993). U and Th concentrations were determined relative to the SL13 standard. Concordia plots and weighted mean concordia ages are calculated using the software SQUID (Ludwig, 2000) and given at 95% confidence limits and include standard calibration errors. 9.2. Relict (pre-M1) Metamorphism

Narryer, Southwest and Youanmi terranes preserve evidence for a number of Palaeoarchaean, Mesoarchaean and early Neoarchaean metamorphic events recorded by U-Pb age determinations on relict monazite, zircon and titanite grains. All of these age determinations are without the context of the original matrix assemblage due of extensive reworking and metamorphism at moderate to high metamorphic grades in the middle Neoarchaean. Consequently, metamorphic setting, PT conditions, P-T paths and distribution of these older events are largely unknown. Palaeoarchaean metamorphism

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in the Narryer Terrane is indicated by U-Pb age determinations from relict zircon grains of ~32843300 Ma age (Kinny et al., 1988; William and Myers, 1987; Kinny and Nutman, 1996; Wilde and Spaggiari, 2007), and are consistent with the detrital zircon maximum deposition age of 3297±5 Ma (William and Myers, 1987). Later Palaeoarchaean metamorphism in the Narryer Terrane is indicated by a single zircon age of 3195±18 Ma and upper intercept zircon age of 3195±81 Ma (Table 1), which overlap with a single zircon age of ~3180 Ma from the Southwest Terrane (Nieuwland and Compston, 1981). Mesoarchaean metamorphic age determinations are almost absent from the Yilgarn craton. It is probable that the Youanmi Terrane experienced Mesoarchaean metamorphism, given the widespread Mesoarchaean volcano-sedimentary sequences of ~2900-3050 Ma age, granites of ~2880-3023 Ma age (GSWA database; Pidgeon and Wilde, 1990; Wiedenbeck and Watkins, 1993; Wang et al., 1998; Nelson, 1997, 1999; Pidgeon and Hallberg, 2000; Mueller and McNaughton, 2000; Fletcher and McNaughton, 2001; Pawley, 2010), and probable metamorphic zircon ages of ~2853-2856 Ma from felsic extrusives in the Southern Cross domain (Wang et al., 1998). High-grade hornfels formed in the contact aureoles of the Windimurra and Lady Alma mafic-ultramafic igneous complexes of 2815-2825 Ma age (Wang et al., 1998; Nelson, 2001; Ivanic et al., 2010). THERMOCALC average-PT calculations on cordierite-orthopyroxene-quartz hornfels assemblages indicate formation at 709±25 ºC and 0.24±0.25 kbar (Appendix 5). Mafic-ultramafic igneous complexes overlap in age with Pb-Pb isochron ages of 2823±31 Ma and 2824±18 Ma from titanite grains in amphibolites from the Kalgoorlie and Kurnalpi terranes (Table 1). There were at least 13 other younger mafic-ultramafic igneous complex events spanning ~2812 to ~2673 Ma in the Youanmi Terrane and EYC (Ivanic et al., 2010; Hoatson et al., 2006; Thorne et al., 2014), any of which may have also developed thin contact aureole assemblages (see Figures 4 and 33). Seafloor alteration assemblages are expected to have originally been widespread in the back-arc settings (e.g. Phillips. 1986), and highly altered protolith chemistries attest to their existence prior to being obliterated by subsequent metamorphic events (Purvis, 1978, 1984). Three periods of early Neoarchaean metamorphism are recognised in different parts of the WYC: ~2772-2782 Ma in the Narryer Terrane and Southern Cross domain, ~2724-2728 Ma in the Murchison domain and ~2700-2715 Ma in the Narryer Terrane (Table 2). A Pb-Pb isochron from titanite inclusions in the Southern Cross domain give an imprecise age of 2782±37 Ma (Table 1), overlapping with granite and felsic porphyry U-Pb zircon ages of 2775±10 Ma and 2772±5 Ma from Copperhead mine (Miller and Rasmussen, 2006), and an unpublished U-Pb zircon age of ~2770 Ma from Corinthia mine. Elsewhere, this metamorphic event is known only by a single U-Pb monazite age of 2775±25 Ma in the Narryer Terrane (Iizuka et al., 2008). Pb-Pb isochrons from matrix titanite in amphibolites from the Murchison domain give ages of 2724±13 Ma and 2728±32 Ma (Table 1). These ages correlate with an unconformity between early and middle Neoarchaean volcanosedimentary sequences in the Murchison domain, and overlap in age with M1 metamorphic ages of 2727±8 Ma from the Hootanui shear zone between Kurnalpi and Burtville terranes (Section 9.3). The break in sedimentation infers contraction and erosion in the Murchison domain, possibly related to stress propagation from arc accretion events at the same time in the Kurnalpi and Burtville terranes to the east. Murchison titanite Pb-Pb ages also overlap with the ~2727-2720 Ma Dalgarangado igneous complex and alternatively may be related to local mafic-ultramafic intrusions. Metamorphic age determinations in the range of 2700 to 2715 Ma have only been identified in the Narryer Terrane; a poorly constrained U-Pb zircon age of ~2700 Ma (Kinny et al., 1990), and U-Pb monazite age of 2715±12 Ma (Table 1). 9.3. M1 Metamorphism

Pb-Pb titanite isochrons and Lu-Hf garnet-whole rock isochrons from high-P amphibolites fall into three spatially and temporally distinct groups with over-lapping errors (Figure 32; Table 2). The oldest is a single Pb-Pb titanite age of 2748±19 Ma from the western Burtville Terrane (Table 1). Two age determinations from high-P assemblages in the Hootanui shear zone between the Burtville and Kurnalpi Terranes, give a pooled age of 2727±8 Ma (Table 1). Five age determinations from

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other shear zones across the Kalgoorlie and Kurnalpi terranes give a pooled age of 2706±10 Ma (Table 1). M1 parageneses are interpreted to have formed during partial burial of buoyant magmatic arc margins during docking events, with diachronaity indicating accretion younging westward. Earliest docking events of 2748±19 Ma and 2727±8 Ma age are restricted to the Kurnalpi / Burtville margin, consistent with the high concentration of known older mafic stratigraphy ranging between ~2730 and 2810 Ma (e.g. Champion and Cassidy, 2010). Further west, the youngest phase of docking at 2706±10 Ma within both Kalgoorlie and Kurnalpi terranes, is consistent with mafic stratigraphy giving dates of ~2690-2720 Ma (Figures 4 and 32). M1 metamorphic dates are consistent with the overprinting relationships where M1 parageneses are typically reworked by D2 deformation fabrics and M2 metamorphism of ~2665-2685 Ma age (Section 5.3). Significantly, all M1 age determinations are older than high-Ca granite magmatism, and M1 parageneses are not associated with the margins of granite-gneiss domes, discounting their formation within sinking keels of vertical tectonics diapirism models (Goscombe et al., 2009). Furthermore, M1 parageneses are spatially restricted relict shear lenses, and not part the regional metamorphic pattern, which also precludes formation by collisional tectonics models. 9.4. M2 Metamorphism

U-Pb monazite, zircon and titanite age determinations from M2 assemblages in this study confirm the 2665-2685 Ma age range inferred previously based on the timing of high-Ca granite magmatism (Section 5.3). M2 metamorphic parageneses in volcanic stratigraphy of the Kalgoorlie Terrane give a monazite U-Pb date of 2675±6 Ma and zircon U-Pb date of ~2670 Ma from Saint Ives Mine (unpublished CET report, 2006), and other zircon ages of 2673±6 Ma (Clout et al., 1990), 2669±11 Ma (Compston et al., 1986) and 2663±7 Ma (Sircombe et al., 2007). The pooled average of all M2 age determinations from across the Yilgarn craton is 2671±6 Ma (n=11), overlapping the peak of early high-Ca granite magmatism (Figure 32). M2 metamorphic age determinations from the Youanmi Terrane span the same age range as the EYC (Figure 32). The Southern Cross domain returns monazite ages of ~2670 Ma and ~2680 Ma (unpublished CET report, 2006), and a Pb-Pb titanite isochron age of 2664±10 Ma from amphibolite (Table 1). Mylonite formed during the main phase of deformation in the Murchison domain gives a SHRIMP U-Pb age of 2668±8 Ma from zircons in partial melt pools (Van Kranendonk and Ivanic, 2009). Pb-Pb titanite analyses from amphibolites in the Murchison domain give dates of 2667±6 and 2681±16 Ma (Table 1). The Southwest Terrane has a single SHRIMP U-Pb metamorphic zircon age of ~2670 Ma (Rennie, 1998). Although the Youanmi Terrane returns metamorphic age determinations from matrix assemblages of the same age range as the EYC, it is probable that M2 regional-contact metamorphism was diachronous (Ridley et al., 1997), starting earlier and being longer-lived in the Youanmi Terrane. Youanmi Terrane high-Ca granite peaks are between 26802695 Ma, and bimodal volcanism was terminated between ~2700-2715 Ma, followed by an unconformity and clastic sedimentation (Figure 4; Champion et al., 2001; Groenewald et al., 2006; Kositicin et al., 2008; Van Kranendonk and Ivanic, 2009). Consequently, the regional-contact metamorphic pattern in the Youanmi Terrane may have started to develop ~5-10 m.y. earlier than the EYC, in association with a separate west-dipping subduction system at the eastern margin of the terrane as proposed by Korsch et al. (2013). 9.5. M3 Metamorphism

Metamorphic monazite in metapelite from the Belches Formation turbiditic late basin in the EYC, have been dated directly by in situ LA-ICPMS, returning a weighted mean Pb207/Pb206 age of 2649±8 Ma (Table 1). This age determination is consistent with the peak of M3 metamorphism estimated between ~2650-2655 Ma on the basis of maximum deposition age of the late basins and fast deposition and heat diffusion rates (Section 5.4). Dating of other metamorphic minerals in EYC late basins has been difficult because of the relatively low metamorphic grade (<550 ºC) and predominance of detrital titanite and zircon grains that were not reset (Appendix 2). Micro-zircon inclusions in garnet resulted in mixed Lu-Hf garnet isochron ages of 2695±9 Ma and 2688±8 Ma,

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and inherited titanites in calcsilicate give a Pb-Pb isochron of 2682±25 Ma, all significantly older than the sediment protolith (Table 1). A significant number of U-Pb monazite, xenotime and titanite age determinations from Au-mineralization and alteration parageneses in the Burtville and Kurnalpi Terranes overlap with M3 metamorphism between ~2650-2665 Ma (Salier et al., 2005; Mueller and Nemchin, 2003). M3 metamorphism in the Youanmi Terrane is constrained by one imprecise U-Pb metamorphic zircon date of 2653±46 Ma from the Murchison belt (Schiotte and Campbell, 1996). Nevertheless, the presence of a late high-Ca granite peak of 2655-2660 Ma age, metamorphosed post-volcanic late basin sequences in Marymia and Diemals belts of <2675 Ma age (Bagas, 1999; Chen et al., 2003), and anticlockwise P-T paths in the Diemals, Southern Cross, Ravensthorpe and Murchison belts, attest to minor M3 metamorphism. M3 metamorphism is also inferred in the Narryer and Southwest terranes on the basis of metamorphosed late clastic sequences deposited after ~2663-2676 Ma (Table 1; Sircombe et al., 2007). M3 metamorphism in the Narryer and Southwest terranes is confirmed by U-Pb metamorphic age determinations younger than onset of late clastic sedimentation, and with dates identical to those in the EYC. A 2660±10 Ma U-Pb xenotime age from late clastic sequences in the Jack Hills belt of the Narryer Terrane has been interpreted as authigenic (Rasmussen et al., 2010). Low-grade M3 metamorphism of these sediments has been confirmed by a U-Pb metamorphic monazite date of 2653±5 Ma (Rasmussen et al., 2010). Higher-grade metamorphic parageneses in the Mount Narryer region yield a SHRIMP U-Pb date of 2659±3 Ma for metamorphic zircon (Table 1). Granulites from the Southwest Terrane have SHRIMP U-Pb metamorphic zircon dates in the same range: 2663±8 Ma and 2659±4 Ma (Rennie, 1998) and 2655±11 Ma (Table 1). The pooled average of M3 metamorphic age determinations from across the craton is 2656±5 Ma (n=12), which overlaps with M3-Au3 alteration and mineralization in the highly extended Kurnalpi Terrane, which average 2655±5 Ma (n=14) (Figure 32). 9.6. M4 Metamorphism

A large dataset of U-Pb monazite, zircon and titanite age determinations (n=42) from M4 assemblages confirm protracted metamorphism spanning 2613-2649 Ma that overlaps closely with low-Ca granite magmatism (Table 1; Pidgeon et al., 1990; Barnicoat et al., 1991; Nemchin et al., 1994; Bosch et al., 1996; Campbell et al., 1998; Dalstra et al., 1998; Napier et al., 1998; Nemchin and Pidgeon, 1999; Mueller et al., 2004; Mueller and McNaughton, 2000; Wilde, 2001; Fletcher and McNaughton, 2001; Sircombe et al., 2007; Muhling et al., 2008; Thebaud and Miller, 2009). These age determinations group into two distinct metamorphic peaks: early-M4 at 2640-2649 Ma with pooled average of 2644±4 Ma (n=11), and peak-M4 at 2613-2639 Ma with pooled average of 2629±7 Ma (n=31). Early-M4 age determinations are only reported from the Murchison domain and Southwest Terrane (Figure 32). Early-M4 metamorphism is best represented by a well-constrained LA-ICPMS date of 2645±8 Ma from matrix monazite in the Lake Grace belt, central Southwest Terrane (Table 1). The predominant peak-M4 metamorphism was more extensive and experienced in the Narryer, Southwest and Youanmi terranes. Peak-M4 metamorphism varies little between each region, with pooled average of 2628±8 Ma (n=4) in the Narryer Terrane, 2630±7 Ma (n=13) in the Southwest Terrane, 2629±10 Ma (n=6) in the Murchison domain and 2629±5 Ma (n=8) in the Southern Cross domain. Peak-M4 metamorphism in the Youanmi Terrane is best represented by LAICPMS dates from metapelite matrix monazite of 2624±11 Ma age in the Southern Cross belt, and 2613±9 Ma in the Dalgaranga belt (Table 1). LA-ICPMS ages from matrix monazite in metapelite granulite from Mount Narryer are 2626±8 and 2624±7 Ma (Table 1). Peak-M4 metamorphism in the Southwest Terrane is best represented by SHRIMP titanite dates from partial melt segregations of 2615±3, 2625±4, 2631±4 and 2637±5 Ma (Nemchin and Pidgeon, 1999), and SHRIMP zircon ages of 2638±2 Ma from metapelite granulite (Table 1), and 2635±5 Ma from orthogneiss (Sircombe et al., 2007). The peak of M4 metamorphism in the WYC overlaps with M4-Au4 alteration and mineralization in the Kalgoorlie Terrane, which average 2630±8 Ma (n=18) (Figure 32).

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9.7. Alteration and Mineralization

A large dataset of age determinations in literature (n=70; e.g. Gazley et al., 2011; Vielreicher et al., 2002; Napier et al., 1998; Mueller and McNaughton, 2000; Mueller et al., 2004; Campbell et al., 1998; Barnicoat et al., 1991; Phillips and Miller, 2006; Kent and Hagemann, 1996; Roberts, 2004; McNaughton et al., 2005; Salier et al., 2004, 2005; Mueller and Nemchin, 2003), show that goldmineralization and alteration events are diachronous and group into four tectonically distinct periods that are variably represented in the different terranes (Figure 4; Robert et al., 2005; Blewett et al., 2008, 2010; Czarnota et al., 2010b). Oldest mineralization events (Au1) between 2665-2675 Ma are uncommon and associated with D2 contraction, M2 metamorphism and the peak of early high-Ca granites (e.g. Robert et al., 2005; Blewett and Czarnota, 2007a). These post-date the cessation of volcanism at ~2670 Ma and pre-date post-volcanic late basins and are restricted to the Kalgoorlie Terrane. The first major mineralization event (Au2) occurred between 2655-2665 Ma and is associated with D3 lithospheric extension, prograde M3 metamorphism, peak of late high-Ca granite and mafic granite, syenite and carbonic intrusions (e.g. Campbell et al., 1998; Mueller and Nemchin, 2003). Au2-mineralization is temporally and spatially closely associated with post-volcanic late basins and mantle-derived magmatism (Hall, 1997), and involved mixing of fluids from both of these sources (Mikucki, 1997; Neumayr et al., 2007; Walshe et al., 2008a,b). All Au2-mineralization is restricted to the Kalgoorlie and Kurnalpi terranes, the largest deposits coincide with the cessation of late basin deposition, and most formed at depths ~1 kbar shallower than the peak of M2 metamorphism (Hall, 1997). The second major mineralization event (Au3) occurred between 26402655 Ma and is widespread across the Kalgoorlie, Kurnalpi and Burtville terranes. This period of mineralization is associated with the stress switch to D4 sinistral transpression, and overlaps with the peak of M3 metamorphism and start of low-Ca granite magmatism (e.g. Mueller and Nemchin, 2003; Bucci et al., 2004; Salier et al., 2005). Au4-mineralization at 2610-2640 Ma is widely scattered across the Kalgoorlie, Youanmi and Southwest terranes, and overlaps with the near cratonwide M4 thermal anomaly and low-Ca granites (e.g. Barnicoat et al., 1991; Wang et al., 1993; Kent and Hagemann, 1996; Campbell et al., 1998; Bucci et al., 2004; McNaughton et al., 2005). These deposits are associated with D5 dextral transtension and exhumation to ~1.0-1.8 kbar depths in the EYC, and low-Ca granites were the local heat source for fluid flow and gold remobilization. 9.8. Isotopic Resetting and Cooling Ages

A significant number of anomalously young metamorphic age determinations <2615 Ma are scattered throughout the Yilgarn craton. Most are Pb-Pb titanite isochron dates of 2607±19, 2594±30, 2571±13, 2600±16, 2598±36 and 2589±34 Ma in Southwest, Youanmi, Kalgoorlie and Kurnalpi terranes, and Sm-Nd garnet-whole rock isochron ages of 2610±16, 2593±13, 2561±11 and 2547±17 Ma in the Kalgoorlie and Kurnalpi terranes (Table 1). Garnets in the analysed samples are very small (<2 mm diameter) and conducive to resetting at relatively low temperatures (Goscombe et al., 2005). SHRIMP apatite age of ~2575 Ma, Pb-Pb whole rock isochron age of ~2530 Ma and U-Pb zircon age of 2581±2 Ma are reported from the Southwest Terrane (Nieuwland and Compston, 1981; Rennie, 1998; Nemchin and Pidgeon, 1999). These age determinations cannot be related to any known orogenic event and are best explained by ongoing resetting of isotopic systems during protracted dissipation of the M4 thermal anomaly (Kent et al., 1996; Goscombe et al., 2009). Ar-Ar muscovite and biotite cooling ages range 2563-2604 Ma in the Kalgoorlie Terrane and 2540-2585 Ma in the Southern Cross domain (Kent and McDougall, 1995; Kent et al., 1996; Kent and Hagemann, 1996; Napier et al., 1998; Campbell et al., 1998; Phillips and Miller, 2006), suggesting very slow cooling rates of ~3.4-7.2 ºC/Ma, which is conducive to ongoing resetting of isotopic systems. 10. DISCUSSION AND CONCLUSIONS: NEOARCHAEAN ACCRETIONARY OROGENIC CYCLE

Containing world-class gold and nickel fields, the Yilgarn craton has received considerable attention over the last 100 years, and a concerted effort by University, Government and Industry researchers in the last 25 years have resulted in craton-wide relational geological datasets (e.g. Blewett et al.,

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2008; Korsch et al., 2013). Coverage has been detailed and comprehensive and includes but is not limited to: geological mapping, mafic and granite magmatism, stratigraphy, deformation history, metamorphism, geochronology, Nd-isotopes and geophysical datasets pertaining to crustal structure. All of these available datasets are consistent with and best explained in the context of a protracted middle Neoarchaean accretionary orogenic cycle, entirely within the plate-tectonic paradigm (e.g. Belt and Standing, 2007; Mary Gee and Swager, 2007; Czarnota et al., 2010a; Blewett et al., 2008; Korsch et al., 2013). The metamorphic datasets have been particularly significant in defining the tectonic setting, and indeed the duality of high-P, low-T/depth M1 parageneses and high-T, highT/depth M2-M3-M4 parageneses is one of the hallmarks of plate tectonics (Brown, 2006). Metamorphic parameters such as T, P and T/depth are first-order variables that vary smoothly with time and encapsulate the average salient crustal response. As a result, the available metamorphic data allows an almost continuous record of crustal response throughout the middle Neoarchaean orogenic cycle (Figure 33). Furthermore, metamorphic data are unfungible and P-T paths for example, allow unambiguous constraints on particle paths, vertical crustal movement and crustal thickening versus thinning. Similarly, average thermal gradients document advection- versus conduction-dominated metamorphic response, constraining probable tectonic settings (Appendix 7). All known metamorphic events (i.e. reaching a distinct thermal peak) have been characterised for; distribution and spatial pattern of metamorphism, relationship to tectono-stratigraphic history and magmatism, metamorphic conditions including average thermal gradient, P-T evolution paths and metamorphic age. Metamorphic parageneses from across the Yilgarn craton preserve a near continuous record of a protracted middle Neoarchaean orogenic cycle spanning ~2610-2750 Ma that shows a complex thermo-barometric evolution which is outlined below (Figure 33). On the basis of overprinting criteria, tectono-stratigraphic setting, geochronology, metamorphic conditions and P-T evolutions, eight temporally distinct metamorphic events (each developing peak metamorphic matrix parageneses) have been identified at: ~2730-2810 Ma, 2748±19 Ma, 2727±8 Ma, ~2675-2715 Ma, 2706±10 Ma, 2671±6 Ma, 2656±5 Ma, 2644±4 Ma and 2629±7 Ma (Figure 32). These conform to a framework of five distinct periods: Ma, M1, M2, M3 and M4 each showing characteristic metamorphic response. These events are not discontinuous and metamorphic conditions vary smoothly through the orogenic cycle, with preceding metamorphic conditions and tectonic setting greatly influencing subsequent tectono-metamorphic response (Figure 33). The middle Neoarchaean metamorphic record shows a progression from subduction related accretionary growth (Sections 10.1, 10.2 and 10.3), terminated by a stress switch to lithospheric extension (Section 10.4), leading inturn to delamination of the eclogitic lower-crust and mantle lithosphere with associated crustal-scale thermal anomalies accompanying reactivation events (Section 10.5), followed by slow heat dissipation during cratonization. 10.1. Magmatic Arcs [Ma]

The earliest metamorphic parageneses (Ma) are rare, relict upper-amphibolite to granulite facies assemblages formed in low to static strain conditions within the Kurnalpi and Burtville terranes. These formed in low-P/high-T thermal regimes of 51-55 ºC/km and low-pressures of 3.7-5.0 kbar, consistent with high heat flow conditions in the upper-crust of magmatic arcs (Ernst, 1976; Kemp et al. 2007). Those in the Kurnalpi Terrane are associated with magmatic arc volcano-stratigraphy and HFSE granites of ~2675-2715 Ma age within the Gindalbie arc (e.g. Champion et al., 2001; Brown et al., 2002; Cassidy, et al., 2006; Barley et al., 2002, 2006; Groenewald et al., 2006; Mary Gee and Swager, 2007). Higher-T granulite parageneses in the western margin of the Burtville Terrane are associated with the Duketon magmatic arc within a rifted fragment of older lithosphere indicated by volcano-stratigraphy of ~2730-2810 Ma age (e.g. Champion and Cassidy, 2010) and Sm-Nd model age signatures (Figure 5; Cassidy and Champion, 2004; Champion et al., 2006; Champion and Cassidy, 2007). These different age magmatic arcs young westward in concert with the sequence of arc accretion events documented by M1 parageneses, discussed below.

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10.2. Arc Accretion [M1]

M1 metamorphism produced high-P (7.0-8.7 kbar), upper-amphibolite facies parageneses in highP/moderate-T thermal regimes, constituting the only Barrovian-series rocks in the Yilgarn craton. The rocks evolved through hairpin clockwise P-T paths of steep burial to peak-PT conditions followed by isothermal decompression, indicating fast burial-exhumation cycles. Low T/depth ratios of 18-25 ºC/km also indicate rapid burial-exhumation settings where advection out-competes heat diffusion. These average thermal gradients are higher than those developed in subduction settings which typically experience T/depth ratios between 13-18 ºC/km (e.g. Brown, 2006; Beyssac et al., 2008), indicating some thermal response to burial. Distribution is restricted to km-scale, low-strain shear lenses preserved within major crustal-scale shear zones, primarily at terrane and magmatic arc boundaries such as the Ida, Ockerburry, Laverton and Hootanui shear zones (Figure 5). Structural settings, depth of burial and rapid exhumation are consistent with partial burial of attenuated margins of buoyant magmatic arcs or terranes during docking events in “subductive”-like settings (e.g. Belt and Standing, 2007; Beyssac et al., 2008; Korsch et al., 2013). M1 parageneses are overprinted and reworked in all the major deformation events that reactivated these shear zones (D2, D3, D4 and D5), indicating they are relict expressions of early accretionary growth. Al-in-hornblende geobarometry of primary magmatic and secondary metamorphic hornblende from Manyutup tonalite calc-alkaline magmatic arc in the Ravensthorpe belt (Witt, 1998, 1999), show 3.0-5.9 kbar conditions during magmatic crystallization, followed by 5.5-8.0 kbar depths during a later burial event (Goscombe et al., 2009). This barometric history is consistent with partial burial of a magmatic arc within a thrust stack (Witt, 1998, 1999), and is shared in common with the Ma and M1 metamorphic histories of accreted arcs in the EYC. Radiometric dates confirm formation of M1 assemblages early in the geological history (~27002750 Ma), over-lapping in age with mafic volcano-stratigraphy at ~2690-2720 Ma, magmatic arcs at ~2675-2715 Ma and older mafic stratigraphy at ~2730-2810 Ma (e.g. Champion and Cassidy, 2010; Champion et al., 2001; Cassidy, et al., 2006; Groenewald et al., 2006). M1 metamorphism was diachronous and age determinations indicate three periods of terrane accretion that young westward (Figure 32). Earliest events at 2748±19 Ma and 2727±8 Ma are restricted to the Hootanui and Laverton shear zones at the boundary between older lithosphere of the Burtville Terrane and relatively juvenile Kurnalpi Terrane (Figures 5 and 34). The youngest M1 events averaging 2706±10 Ma are restricted to shear zones from within the Kalgoorlie and Kurnalpi terranes: in particular the Ockerburry shear zone on the west margin of the Gindalbie magmatic arc, and Ida shear zone at the hinterland margin. Earliest accretionary events assembled the Burtville – Kurnalpi fore-arc, and youngest events eventually closed the Kalgoorlie back-arc basin, prior to shallow subduction, felsic volcanism and associated high-Ca granite magmatism at ~2690-2695 Ma (Figure 34; e.g. Czarnota et al., 2010a; Korsch et al., 2013). 10.3. Subduction and Granite Bloom [M2]

M2 metamorphism produced widespread regional-contact type metamorphic patterns associated with voluminous early high-Ca granite magmatism. Metamorphism is characterized by; low-pressures (3.5-5.0 kbar), low-P/moderate-T thermal regimes (~35-45 ºC/km), and tight low-P/T clockwise P-T paths showing only minor burial and crustal thickening accompanying contraction. The juvenile mafic-ultramafic volcano-stratigraphy of the EYC formed with plume-related processes active in the back arc of a retreating subduction system to the east, (Figure 34; e.g. Rey et al., 2003; Cassidy and Champion, 2004; Champion and Cassidy, 2007; Blewett and Czarnota, 2005; Czarnota et al., 2008, 2010a; Korsch et al., 2013). Similar plume and back-arc settings of the same age have been proposed for the Abitibi Subprovince (Wyman, 1999; Wyman et al., 2002). From ~2690-2695 Ma onwards, shallow-dipping subduction generated bimodal volcanism and large volumes of early high-Ca granite by partial melting of the down-going slab leaving eclogite resite (Figure 35; Champion et al., 2001; Cassidy, 2006; Champion and Cassidy, 2007). Neutral buoyancy pooling of high-Ca granite immediately below the mafic-ultramafic volcano-stratigraphy,

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contributed ~40-45 % volume of the upper-crust (e.g. Champion et al., 2001; Czarnota et al., 2010a). The regional-contact type metamorphic pattern and high average thermal gradients of M2 metamorphism (Figure 7) are consistent with advection of heat into upper-crust by repeated ascent of large volumes of felsic magma (e.g. Binns et al., 1976; Ridley, 1993; Stüwe et al., 1993; Mikucki, 1997; Goscombe et al., 2007, 2009). The magmatic heat budget was enhanced by radiogenic heat from anomalously high Thorium contents in the Yilgarn granites (Champion and Cassidy, 2007). Peaks in early high-Ca granite magmatism were diachronous between 2670-2675 Ma and overlap exactly with the peak of M2 metamorphism dated at 2671±6 Ma (e.g. Table 1; Champion et al., 2001). M2 metamorphism contributed the first significant and widespread thermal anomaly, substantially increasing the thermal budget of the upper-crust, reducing rheological rock strength and priming the crust for reworking during subsequent stress-switches (Goscombe et al., 2007, 2009). M2 metamorphism overlaps in age with crustal shortening at ~2665-2670 Ma (D2) that produced widespread but heterogeneous foliation development at moderate strains (Figure 6). Low-P/T clockwise P-T paths during M2 metamorphism show only minor amounts of burial (~1.0-1.5 kbar) through maximum pressures no greater than 6.0 kbar. Consequently, D2 deformation was insufficient to generate appreciable crustal thickening or support gravitational instability (Goscombe et al., 2007, 2009). Deformation coincides with the termination of volcanism at ~2670-2673 Ma, which is interpreted due to either choking of the subduction zone by an oceanic plateau or advancing subduction system; resulting in the cessation of shallow subduction (Blewett and Czarnota, 2007c,b; Czarnota et al., 2010a). Cessation of subduction typically leads to down sagging of the subducted slab, inducing asthenospheric inflow and melting of the mantle wedge, further contributing to highCa granite magmatism (e.g. Champion and Cassidy, 2007; Czarnota et al., 2010a). Termination of subduction in this way can induce roll back of the subduction hinge and extension of the upper-plate lithosphere (Rawling and Lister, 1999). M2 crustal thickening was insufficient to induce internal, gravitationally driven extension. Consequently, lithospheric extension during M3 (Section 10.4) is interpreted to be due to an external stress field such as slab rollback. Similar stress switches from horizontal 1 to vertical 1 are a common feature of Phanerozoic accretionary orogens due to changes in outboard subduction zone dynamics (e.g. Foster and Gray, 2000). 10.4. Lithospheric Extension [M3]

Termination of shallow subduction generated magmatism at ~2670-2673 Ma was followed only ~58 m.y. later by a stress-switch to lithospheric extension at ~2665-2668 Ma. It is interpreted that the two events are related, with sag of the floundered subducted plate leading to slab rollback and extension of the upper-plate (Figure 36; Blewett and Czarnota, 2005; Goscombe et al., 2007, 2009; Czarnota et al., 2010a). Extension produced linear and arcuate turbiditic and fluvial rift basins floored by crustal-scale extensional detachments and granite-gneiss domal core complexes in the footwall (e.g. Blewett and Czarnota, 2005). Deep seismic reflection profiles show metamorphic core complex geometries, including east dipping low-angle and listric extensional shear zones that merge at depth and penetrate the full ~40 km crustal thickness (Goleby et al., 1993, 2002a; Blewett et al., 2004a; Blewett and Czarnota, 2007c). The lower-plate, granite-gneiss domes are wrapped by arcuate extensional detachments and turbiditic late basins, and these footwalls exhumed older volcanostratigraphy and high-P M1 parageneses (Figure 5; Blewett and Czarnota, 2005; Goscombe et al., 2007, 2009). This pattern is consistent with extensional telescoping of the crust with transport away from lower-plate domal cores. The cores of lower-plate domes are commonly intruded by late-stage low-Ca granite plutons (Blewett and Czarnota, 2005). Extension was mostly focused in the EYC, though late clastic sequences have been documented across the entire craton, indicating that extension was craton-wide. Late clastic sequences were deposited from ~2665-2669 Ma in the EYC, ~2663-2675 Ma in the Youanmi and Southwest terranes and from ~2665 Ma in the Narryer Terrane (e.g. Table 1; Bagas, 1999; Chen et al., 2003; Krapez et al., 2000; Squire 2006, 2007; Sircombe et al., 2007; Rasmussen et al., 2010). Turbiditic basins were followed by fluvial sequences in linear rifts from ~2657-2662 Ma, and deposition continued to ~2652 Ma in the EYC, ~2646 Ma in the Southwest Terrane and ~2660 Ma in the Narryer Terrane (Rasmussen et al., 2010; Sircombe et al.,

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2007). This age range overlaps with late high-Ca granites that involved a greater component of crustal melt, as well as mafic granite, syenite and carbonatite derived by partial melting of metasomatized mantle (Champion et al., 2001; Champion and Cassidy, 2007). M3 metamorphism occurred in widely separated arcuate zones around 10-50 km wide by 40-200 km long, closely associated with extensional detachments and rift basins. These zones of metamorphism coincide with the locus of maximum extension and high heat flow in the upper-plate of asymmetric extensional settings (e.g. Buck et al., 1988; Issler et al., 1989; Ruppel, 1995). These thermal anomalies were superimposed on the waning regional-contact M2 metamorphic pattern, producing a second period of low-P/moderate-T metamorphism with 31-45 ºC/km average thermal gradients. M3 metamorphic parageneses within extensional shear zones and post-volcanic clastic sequences document anticlockwise P-T paths involving steep-P/T burial to ~2.5-4.5 kbar at the peak of metamorphism, followed by near isobaric cooling. These P-T paths are diagnostic of extensional settings and the high T/depth ratios are typical of rift settings (Appendix 7; Buck et al., 1988; Issler et al., 1989; Sandiford et al., 1998; Webb, 2001; Alias et al., 2002). Metamorphic field gradients across extensional shear zones show steep gradients to lower-PT in the hanging wall, consistent with extensional telescoping (Williams and Currie, 1993; Williams and Whitaker, 1993). Beddingparallel main foliations and small-scale recumbent isoclinal folding is developed in the turbiditic Belches Formation (Groenewald et al., 2000). These are indicative of vertical 1 during rifting, and common in other metamorphosed extensional basins (Mawby et al., 1999; Webb, 2001; Alias et al., 2002; Foden et al., 2006; Buick et al., 2005). M3 metamorphic parageneses are primarily preserved in the EYC where extension was greatest and associated with widespread post-volcanic late basins, crustal-scale shear zones and core complexes (Blewett and Czarnota, 2007c). However, relict, early low-pressure parageneses showing anticlockwise P-T paths, metamorphosed late clastic sequences and M3 metamorphic age determinations have been documented at localities across the Youanmi, Southwest and Narryer terranes. Across all terranes, the peak of M3 metamorphism was attained at 2656±5 Ma, approximately 10-15 m.y. after the start of extension and deposition of the late clastic sequences. The turbiditic Belches basin was inverted by minor shortening and crenulation development, immediately after the peak of metamorphism and during D4 sinistral transpression between ~2645-2655 Ma (Blewett and Czarnota, 2007c). These metamorphic conditions, T/depth ratios and anticlockwise P-T evolutions are typical of turbiditic rift basins showing fast deposition rates (Mawby et al., 1999; Webb, 2001; Alias et al., 2002; Buick et al., 2005; Foden et al., 2006). The ~10-15 m.y. delay between the start of extension at ~2665-2669 Ma and the peak of M3 metamorphism at ~2656 Ma, is consistent with fast deposition rates documented in turbiditic rift basins and heat diffusion rates in rift settings, including the following examples: 1. Deposition rates of 2.5 km/m.y. have been documented for the Neoarchaean Porcupine Group turbiditic basin in the Superior Province and turbidite basins of 2660-2680 Ma age in the Slave Craton (Bleeker, 2003). Such a deposition rate would give the 14 km of burial documented in the Belches turbiditic basin within 5.6 m.y. 2. Thermal modelling of asymmetric extension has shown that basin depths of 16 km can be attained within <2 m.y. producing temperatures >500 ºC at the base of basinal sequences within <10 m.y. (Issler et al., 1989). 3. The Cambrian Kanmantoo basin in South Australia, contains an 8 km thick turbiditic sequence that was deposited within <10 m.y., and is associated with the earliest phase of slab rollback induced extension in the Tasmanide Orogen (Foden et al., 2006). Metamorphism in the Kanmantoo rift basin followed anticlockwise P-T paths and peak metamorphism was attained only ~2 m.y. later, immediately before inversion by shortening (Webb, 2001; Alias et al., 2002; Foden et al., 2006). 4. The Cambrian Irindina Group in the Arunta Block represents a >20 km deep turbiditic rift basin floored by mafic igneous rocks that similarly experienced fast deposition and peak metamorphism within <10 m.y. of the start of extension (Mawby et al., 1999; Buick et al., 2005). In all of these cases the peak of metamorphism was attained soon after maximum depths of burial. Both modelling and empirical examples indicate burial-deposition rates and thermal diffusion keep pace with each other, resulting in fast deposition-metamorphism cycles occurring within 10-12 m.y.; and consistent with the 10-15 m.y. delay documented for the Belches Formation and M3 metamorphism in general.

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The stress-switch from horizontal 1 during D2-M2 contraction at 2670-2675 Ma, to vertical 1 at 2665-2669 Ma during D3-M3 extension, profoundly changed the subsequent evolution of the Yilgarn craton. The key event that induced the switch to M3 lithospheric extension within 5-8 m.y., was cessation of volcanism at ~2670-2675 Ma and presumably termination of shallow subduction (Figure 35; Blewett and Czarnota, 2007c). Onset of D3-M3 extension has been interpreted to follow from gravitational sag of the floundered subducted plate in a chocked subduction zone, leading to slab roll back and hinge migration inducing extension of the upper-plate lithosphere (Figure 36; e.g. Blewett and Czarnota, 2007c; Goscombe et al., 2007, 2009; Blewett et al., 2008, 2010; Czarnota et al., 2010a). An alternative first-order cause could be global reorganization of plate trajectories following attainment of a metastable plate configuration, such as assembly of supercontinents. Similar stressswitches leading to extensional reactivation have been documented following: culmination of Gondwana assembly at ~510-520 Ma (e.g. Goodge et al., 1993; Goodge, 1997; Bierlein et al., 2002, 2006; Meert, 2003; Foster et al., 2005, 2009; Foden et al., 2006; Goscombe et al., 2017b), and culmination of accretionary growth leading to slab roll back at the east Gondwana margin at ~440 Ma (Squire and Miller, 2003). If the Yilgarn craton stress-switch was due to a global plate reconfiguration event, similar fundamental switches in tectonics should be recorded in other Neoarchaean terranes world wide at ~2665-2669 Ma, with the possibility of a self-similar global gold mineralization period between ~2610-2665 Ma. Extension elsewhere of similar age is indicated by turbidite late basins of 2660-2680 Ma age in the Slave Province that are broadly contemporaneous with Yilgarn craton late basins (Bleeker, 2003; Bleeker and Hall, 2007). The Abitibi Subprovince experienced lithospheric extension and post-volcanic turbiditic and fluvial rift basins with ages estimated between 2670-2680 Ma (Daigneault et al., 2002) and 2675-2696 Ma (Ayer et al., 2002). The Superior Province experienced a middle Neoarchaean orogenic cycle with almost identical history to the Yilgarn craton (Beakhouse et al., 1999; Beakhouse, 2007). A similar stress-switch to lithospheric extension and orogen collapse late in the orogenic cycle has been documented for the Palaeoarchaean evolution of the Barberton craton (Kisters et al., 2003). Alternative models for lithospheric extension by gravitational spreading in response to lithospheric over-thickening, such as Basin and Range-type settings (Tirel et al., 2006), is not probable for the Yilgarn craton because D2 contraction merely thickened previously thin lithosphere and did not result in appreciable crustal thickening (Section 10.3). It could be argued that the Yilgarn Craton would not be as endowed in world-class gold deposits without the switch to D3-M3 lithospheric extension. Many of the mineralization processes, structural architectures and spatial-temporal juxtapositions required for large-scale gold deposits, formed during the M3 period. The M3 thermal anomaly overlaps in age with the second gold-mineralization event and immediately pre-dates the main gold-mineralization events (Figure 4; Section 9.7). Lithospheric extension and crustal thinning resulted in higher heat flow that drove many of the key processes leading to mineralization; in particular formation of a range of different fluid types (Figure 37; Section 8.6; e.g. Mikucki, 1997; Neumayr et al., 2007; Walshe et al., 2008a,b). Lithospheric extension was heterogeneous, resulting in basinal rifts and linear to arcuate zones of high heat flow, resulting in steep thermal gradients that drove fluid flow (e.g. Czarnota et al., 2010b). The extensional stress field modified fluid flow trajectories, giving both lateral and in part downward flow paths (Sheldon et al., 2007, 2008), allowing the mixing of different fluid reservoirs. Fluid circulation cells associated with magmatic intrusions persist longer in extensional stress regimes, increasing the temporal window for mineralization (Sheldon et al., 2008). Extensional rifting dropped saline basinal fluid reservoirs into the upper-crust, significantly increasing the ambient fluid volume and introduced a new fluid type (Section 8.6; e.g. Mikucki, 1997; Neumayr et al., 2007). Lithospheric extension and associated deep penetrating structures facilitated the transport of mafic granite, syenite, lamprophyre and carbonatite melts and dry strongly reduced fluids from the mantle up into the upper-crust (Walshe et al., 2008a,b). These mantle-derived melts and fluids are the most probable primary source of the gold introduced into the upper-crust (Walshe et al., 2008a). The mixing of this distinct fluid type with different hydrous fluid reservoirs in the upper-crust ultimately

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resulted in precipitation and mineralization (Neumayr et al., 2007; Sheldon et al., 2008; Walshe et al., 2008a,b). Lithospheric extension created and reactivated crustal shear zones, generating suitable structural architectures such as extensional detachments and footwall domes focusing fluid flow, facilitating fluid mixing and generating precipitation sites (Blewett et al., 2010; Czarnota et al., 2010b). High heat flow and decompression of lower-plate domes facilitated crustal melting and emplacement of low-Ca granites into the upper-crust, extending the thermal anomaly and facilitating further fluid flow and mineralization. 10.5. Delamination [M4]

M4 metamorphism occurred during a protracted period of regional-scale high-T pulses spanning ~2610-2650 Ma, with peaks at 2644±4 Ma and 2629±7 Ma (Section 9.6). These thermal anomalies produced low-P/high-T parageneses up to granulite facies grade across a broad region of the Narryer and Southwest terranes, lower- to middle-amphibolite facies in the western Youanmi Terrane, and a broad lower-grade margin extending east across the Kalgoorlie Terrane. Metamorphic grade varies along smooth gradients between these regions except for a sharp increase in grade across the Yalgar shear zone into the Narryer Terrane, and relatively steep gradient across the Westonia shear zone into the Southwest Terrane. Average thermal gradients were high at all metamorphic grades (38-48 ºC/km), and metamorphism followed tight, low-P/T clockwise P-T paths, indicating only minor burial and crustal thickening accompanying the low to moderate strains. This period of metamorphism coincides with large volumes of ~2615-2650 Ma low-Ca granite intruded cratonwide, and amounting to ~30-35% of all granite in the Yilgarn craton and 20-24% of the upper-crust (e.g. Champion et al., 2001). Low-Ca granite formed by partial melting of the middle- to lowercrust, and together with charnockite and HFSE granite in the Southwest Terrane, indicates that the M4 thermal anomaly was responsible for high temperatures in the lower-crust and extended cratonwide (e.g. Champion and Sheraton, 1997; Champion and Cassidy, 2007). The large volume of lowCa granite involved a significant transfer of heat into the upper-crust by advection of granite magma (e.g. Champion and Sheraton, 1997; Champion and Cassidy, 2007). Regionally extensive and relatively smooth patterns of M4 metamorphism indicate diffusion of heat on a broad-scale front, and with thermal anomalies significant enough to generate whole-scale granulite facies metamorphism in the middle- to upper-crust. The high intensity of M4 thermal anomalies and high T/depth ratios of ~38-48 ºC/km, indicates an asthenospheric heat source at the base of the crust (Figure 38). These features are consistent with delamination of the eclogitic lowercrust and mantle lithosphere, and upwelling of the asthenosphere (e.g. Smithies and Champion, 1999; Blewett et al., 2004a, 2008; Czarnota et al., 2010a; Korsch et al., 2013). Density driven inversion caused by delamination generates relatively shallow upwelling of the asthenosphere and high geothermal gradients in the crust above, resulting in relatively fast metamorphism compared to other tectonic settings (Rey, 2006; Percival and Pysklywec, 2007). The Yilgarn craton experienced multiple thermal anomalies over a protracted period of time and was accompanied by relative low strain states late in the orogenic cycle; all of which are typical features of delamination (e.g. Houseman et al., 1981; Arndt and Goldstein, 1989; Houseman and Molnar, 1997; Lustrino, 2005; Scharf et al., 2013; Goscombe et al., 2017a). Upper mantle tomography mapped by broadband seismology and magnetotelluric data, indicate a flat Moho and fast shear-wave velocity body interpreted to be delaminated eclogitic lower-crust beneath the craton (Blewett et al., 2008, 2010; Czarnota et al., 2010a). Magnetotelluric data indicate that the base of the lithosphere is at 130 km, shallower than typical Archaean cratons and indicating lithospheric delamination probably occurred (Blewett et al., 2008, 2009). Craton-scale variation in intensity of M4 thermal anomalies, are anticipated to result from variation in size, shape and thickness of detached and partially detached slabs, and the geometry of slab edges and asthenospheric flow (Blewett et al., 2004a). Upper mantle tomography shows a very large detached slab below the Southwest Terrane, consistent with the broad thermal anomaly and the relatively smooth pattern of metamorphism (Figure 9; Blewett et al, 2008). The Southwest Terrane has a broad

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low-gravity anomaly with gradient decreasing towards the southwest, indicating a relatively thicker section of lower-crust was detached, consistent with the high-metamorphic temperatures attained. Delamination of eclogitic lower-crust reduces density of the crustal section as well as bringing hot material into shallower levels, resulting in rapid isostatic rebound of the remaining crustal section (Arndt and Goldstein, 1989; Lustrino, 2005; Rey, 2006). Variation in average crustal depth of the current exposure level across the craton is consistent with greater exhumation of the Southwest Terrane compared to elsewhere in the craton (Figure 22). Average pressures in different parts of the Southwest Terrane range between ~4.9-6.2 kbar, indicating exhumation of P~1.2-2.5 kbar relative to the Youanmi Terrane and P~0.3-1.9 kbar relative to the remainder of the craton in general (Table 2). Low-Ca granite intrusion was diachronous and systematically youngs westward, from 2630-2660 Ma in Burtville, 2620-2655 Ma in Kurnalpi, 2600-2655 Ma in Kalgoorlie, 2600-2650 Ma in Youanmi and 2570-2645 Ma in the Southwest Terrane (Figure 4; Champion et al., 2001). This pattern has been interpreted to be due to westward propagation of delamination in the EYC (Czarnota et al., 2010a). However, timing of arrival of the M4 thermal anomaly in the upper-crust shows almost no diachronaity across the craton (Figure 32). M4 thermal anomalies in the Kalgoorlie Terrane are represented by latest-stage gold mineralization and alteration events spanning 2619-2644 Ma, and averaging 2630±8 Ma (Figure 32). Whereas, near absence of gold mineralization younger than 2650 Ma in the Kurnalpi and Burtville terranes (Figure 4), indicate a diminished thermal anomaly and the eastern limit of appreciable M4 effects. Timing of the M4 thermal peak varies little across the WYC and overlaps with alteration in the western EYC, with pooled averages of 2630±8, 2629±5, 2629±10, 2630±7 and 2628±8 Ma in the Kalgoorlie, Southern Cross, Murchison, Southwest and Narryer terranes respectively (Figure 32; Section 9.6). At the boundary between the EYC and WYC, there is an apparent decoupling of low-Ca granite peaks from arrival times of the M4 thermal anomaly in the upper-crust, as represented by metamorphism, alteration and mineralization ages (Figure 32). This is due to low-Ca granite melts being produced earlier in the EYC in response to extreme lithospheric extension and the M3 thermal anomaly, in addition to the later M4 thermal anomaly. Whereas, in the WYC M3 lithospheric extension was much diminished and low-Ca granite melt was generated almost exclusively in response to the M4 thermal anomaly. Three probable triggers for delamination in the Yilgarn craton have been proposed. Internal triggers such as, [1] thermal softening and gravitational instability of the eclogitic underplate that was generated by extensive partial melt extraction during high-Ca granite magmatism (e.g. Arndt and Goldstein, 1989; Ducea, 2002; Lee et al., 2006; Champion and Cassidy, 2007). External triggers such as, [2] D2-M2 contraction at ~2670-2675 Ma during termination of subduction (e.g. Smithies and Champion, 1999; Blewett et al., 2004a), or [3] D3-M3 lithospheric extension at ~2665-2668 Ma (e.g. This paper; Blewett et al., 2008; Czarnota et al., 2010a; Korsch et al., 2013). Delamination either occurred as a natural consequence of internal processes such as reaching a density threshold in the underplate, or is more typically triggered by an external stress field at some stage after establishing the dense underplate (Houseman et al., 1981; Houseman and Molnar, 1997). Irrespective of the proximal trigger, delamination is premised on the existence of a widespread eclogitic underplate to generate the gravitational instability necessary (Arndt and Goldstein, 1989). As a result, a maximum possible age limit for delamination can be established by the peak of early high-Ca granite magmatism at ~2670-2675 Ma (Champion et al., 2001), and associated regionalcontact M2 metamorphism at 2671±6 Ma (Section 5.3). M3 anticlockwise metamorphic parageneses and late clastic sequences of <2665-2675 Ma age are preserved in all terranes, indicating lithospheric extension propagated across the whole craton and was the most probable proximal trigger for delamination. D3 extensional detachments are listric and shallow with depth, merging into low-angle structures that penetrate the lower-crust, potentially decoupling the eclogitic underplate as broad slabs (Goleby et al., 1993, 2002a,b; Drummond et al., 2000; Blewett and Czarnota, 2007a,c; Korsch et al., 2013). In contrast, un-triggered delamination due to softening and gravitational instability of the eclogitic underplate would result in local

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“drips”, and a heterogeneous pattern of delamination, metamorphic anomalies and age determinations. Similarly, the pattern of strain distribution during D2-M2 contraction was highly distributed (Figure 6) and unlikely to trigger craton-wide decoupling of the eclogitic underplate. Consequently, the craton-wide extent and seemingly instantaneous delamination event indicated by craton-wide thermal peak in a narrow age band of 2629±7 Ma, indicates that the most probable trigger for delamination was D3-M3 lithospheric extension at ~2665-2668 Ma. The first thermal peak was attained at 2644±4 Ma in the Southwest Terrane and Murchison domain, and followed by craton-wide peak metamorphism at 2629±7 Ma across the Narryer, Southwest, Youanmi and Kalgoorlie terranes (Figure 32). These thermal pulses arrived in the upper-crust ~21-24 m.y. and ~36-39 m.y. after M3 lithospheric extension at ~2665-2668 Ma, and are consistent with the time scales for heat diffusion through ~40-47 km of crust (e.g. Houseman et al., 1981; Campbell and Hill, 1988; Houseman and Molnar, 1997; Ducea, 2011; Rey, 2006; Percival and Pysklywec, 2007; Czarnota et al., 2010a). 10.6. Comparison with Phanerozoic Accretionary Orogens

The progressive evolution of middle Neoarchaean metamorphic response in the Yilgarn craton is shared in common with Phanerozoic accretionary orogenic systems, such as the Tasmanides (e.g. Foster and Gray, 2000; Bierlein et al., 2002, 2006; Foster et al., 2005; Gray and Foster, 2004; Gray et al., 2002). Both evolved from early magmatic arcs and arc docking events, widespread low-P regional-contact type metamorphism and magmatism related to long-lived subduction, stressswitches to lithospheric extension and core complex development after the termination of subduction, and late-stage broad-scale thermal anomalies and post-kinematic granite magmatism. These two accretionary orogens separated in time by more than 2 billion years share remarkably similar deformation and magmatic histories, evolution in metamorphic response and gold mineralization. The Neoarchaean period is responsible for significant and rapid crustal growth and may have involved a greater proportion of accretionary orogenic settings than is currently recognised (e.g. Mueller et al., 2010; Condie and Aster, 2010; Gray et al., 2007). Neoarchaean accretionary crustal growth in the Yilgarn craton built out from a stretched and rifted Mesoarchaean cratonic margin (e.g. Drummond et al., 2000; Champion and Cassidy, 2007; Blewett et al., 2010), similar to the Slave Province (Bleeker and Hall, 2007). The preferred tectonic model for the middle Neoarchaean evolution of the Yilgarn craton is an accretionary orogen with broad and complex back-arc and fore-arc regions containing; back-arc basins, fore-arc basins, multiple magmatic arcs and subduction systems, rifted island arcs and crustal slivers, all in the upper-plate of a long-lived shallow subduction system (e.g. Brown et al., 2002; Belt and Standing, 2007; Czarnota et al., 2010a; Korsch et al., 2013). Similar shallow subduction zone settings, protracted orogenic histories and timing are shared in common with the Neoarchaean Slave and Superior Provinces (Bleeker, 2003; Ketchum et al., 2004; Bleeker and Hall, 2007). Both the Slave Province and Abitibi Subprovince show similar orogenic histories to the Yilgarn, with early magmatic arc accretion and subduction related deformation and granite bloom, followed by extension and post-volcanic turbiditic and fluvial rift basins, craton-wide low-Ca magmatism and late-stage extension and high-grade metamorphism accompanying exhumation (Daigneault et al., 2002; Wyman et al., 2002; Ayer et al., 2002; Bedard et al., 2003; Ketchum et al., 2004; Bleeker and Hall, 2007). The accretionary setting and similar histories have long been recognised in common with other Neoarchaean and younger gold provinces worldwide, and with current west Pacific settings (e.g. Kerrich and Wyman, 1990; Foster and Gray, 2000; Bierlein et al., 2002, 2006; Cawood et al., 2011). Ma-M1 parageneses suggest east to west propagating closure of subduction systems and accretion of at least two magmatic arcs within the back-arc region of the outboard master subduction system. These were contemporaneous with back-arc extensional basins and plume-related mafic-ultramafic volcanic sequences (e.g. Czarnota et al., 2010a; Korsch et al., 2013). A similarly complex assemblage of upper-plate tectonic settings and accretionary closure of sub-basins is in common with the Lachlan Orogen and modern accretionary systems in general (e.g. Foster and Gray, 2000).

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Furthermore, the greater part of the Yilgarn Neoarchaean deformation history was extensional (D1, D3, D5 and D6) and punctuated by relatively short-lived contractional episodes (D2 and D4), a pattern also shared in common with accretionary orogens in general (e.g. Foster and Gray, 2000; Cawood et al., 2011; Collins, 2002a,b; Collins and Richards, 2008). The Cambrian-Ordovician Lachlan Orogen and broader Tasmanide Orogen in general, experienced multiple slab roll back episodes producing the earliest back-arc settings, later extensional sub-basins, high-heat flows and post-kinematic granites, and was punctuated by at least five short-lived contractional episodes (Figure 39; e.g. Foster and Gray, 2000; Gray et al., 2002; Collins, 2002a,b; Gray and Foster, 2004; Foster et al., 2005; Collins and Richards, 2008). The orogenic cycles in both were long-lived and spanned ~300 m.y. in the Tasmanide Orogen and ~140 m.y. in the Yilgarn craton (Figure 39). This contrasts with collisional orogens, where main phase orogenesis and metamorphism typically last no longer than ~50-60 m.y. (e.g. England and Thompson, 1984; Jamieson et al., 2002; Goscombe et al., 2017a, 2018). Both the Yilgarn and Lachlan are major gold provinces, probably reflecting crustalscale extensional structures tapping the mantle, multiple basin and fluid reservoirs and multiple stress-switches reactivating and remobilizing previous structures and mineralization (e.g. Kerrich and Wyman, 1990; Hronsky, 1993; Hall, 1997). In both orogens the far field stress switches between extension and contraction modes are crucial to gold mineralization models (Foster and Gray, 2000; Gray et al., 2002; Gray and Foster, 2004; Blewett et al., 2008, 2010; Czarnota et al., 2010b). Metamorphic conditions in both the Tasmanides and Yilgarn are characterized by low-P/high-T thermal regimes in general throughout most of their histories, with T/depth ratios spanning ~28 to 50 ºC/km (Figure 40). High-P/low-T metamorphic parageneses (i.e. <18 ºC/km) associated with obducted subduction settings are well preserved within the Tasmanides, from the Tyennan, central Lachlan and New England Orogens, yet completely absent from the Yilgarn craton (Figure 40). Similarly, the Yilgarn craton is devoid of widespread regional metamorphic high-P/moderate-T (Barrovian-series) parageneses with T/depth ratios in the range ~18-28 ºC/km, which are only preserved as spatially restricted relicts (i.e. M1). Barrovian conditions are characteristic of collisional orogens that supported over-thickened crust (Appendix 7). For example, collisional orogens spanning a range of different crustal architectures, such as shallow-dipping fold-thrust belts (e.g. Himalayas) and those with steep orogenic margins (e.g. Zambezi Belt), both develop similar metamorphic response with thermal regimes skewed to T/depth ratios in the range ~18-28 ºC/km (Figure 40). Metamorphic response in collisional belts contrast strongly with accretionary orogens, dominated as they are by much higher thermal regimes and T/depth ratios typically >25 ºC/km (Figure 40). When compared with the full range of plate-tectonic settings represented by Phanerozoic orogens globally, the Yilgarn and Archaean in general differ primarily in; [1] the predominance of low-P/high-T thermal regimes, [2] lower proportions of high-P/moderate-T metamorphism such as in collisional orogens, and [3] near absence of high-P/low-T thermal regimes developed in subduction systems (Figure 41; Brown, 2006). The pattern of metamorphic response in the Yilgarn craton is shared in common with most Neoarchaean cratons and representative of Archaean metamorphism in general (Figure 41); indicating accretionary-type systems probably dominated the Neoarchaean. Despite a Neoarchaean magmatic history showing a long-lived shallow subduction system outboard to the east (Cassidy et al., 2002; Champion and Cassidy, 2007; Czarnota et al., 2010a), high-P/low-T blueschist-eclogite subduction parageneses with T/depth ratios <18 ºC/km are absent from the Yilgarn craton. This is shared in common with both ancient and modern orogens worldwide, where subduction parageneses and obducted ophiolite are only very rarely preserved in the rock record. This is typical of orogenic belts because; [1] subduction parageneses are transported into the mantle and lost, [2] the deeply buried rocks need to be accreted into the hanging wall by ad hoc structural histories, and [3] these parageneses are formed early in the orogenic cycle and thus typically recrystallized and obliterated during ongoing orogenesis and metamorphism (e.g. Yardley, 1982; Mortimer, 2000). In contrast, the absence of widespread high-P/moderate-T (Barrovian-series) metamorphism in the Yilgarn and Archaean in general is unlikely to be a preservation issue given that Barrovian parageneses are well preserved in Proterozoic and Phanerozoic collisional orogens.

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Applying Ockham’s razor, the simplest explanation is a lower proportion of collisional orogens producing over-thickened crust in the Archaean. This would be the natural result of low proportions and smaller fragments of continental crust in this early period of plate tectonics and generally shallower dipping subduction. It is probable then, that accretionary orogens such as the Yilgarn craton predominated by default and the Neoarchaean was a period of rapid and significant continental-growth and recycling necessary before continent-continent collisional systems could become common (e.g. Condie and Aster, 2010). ACKNOWLEDGEMENTS

This research was a largely self-funded ITAR project that was also generously supported with funding for fieldwork and analyses by AMIRA, GA, GSWA, ARC grants A00103456 and DP0210178 to Prof. David Gray and the University of Florida, making the large-scale metamorphic research program possible. GA and GSWA directors are acknowledged for facilitating the research support and allowing access to thin sections, databases, field vehicles and geochronology. Stephen Wyche, A.H. Hickman, Martin Van Kranendonk, Tim Ivanic, Sarah Jones, Sandra Romano and Ivan Zibra supplied useful metamorphic samples. Anthony Crawford, John Walsh, Martin Van Kranendonk, John Miller, Heather Sheldon, Kevin Cassidy, Dave Champion, Paul Hanson, Roland Maas, Mike Sandiford and Nicolas Thebaud are acknowledged for helpful discussions and insights. Angus Netting (Adelaide Microscopy) helped with electron microprobe analyses and garnet maps. Bulk rock analyses and GIS metamorphic map production were undertaken at GA. Dr Roland Maas (Melbourne University) and GSWA staff (Curtin University) assigned with some of the geochronology. Editorial advice of Prof. Zhao, Prof. M. Santosh, Prof. Tim Horscroft and anonymous reviewer significantly improved the manuscript. Mike and Rida Jacob, Mr Gift, Margaret and Roddy MacAskill, Finlay and Rachel McDonald and Alex Urquhart are thanked for their kind hospitality at bonnie camps, farms and crofts during analysis and write up.

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FIGURE CAPTIONS:

Figure 1: Simplified geological map of the Yilgarn Craton, compiled from published GSWA and GA interpretive solid geology maps. Nomenclature and boundaries of domains and terranes are after Cassidy et al. (2006). Interpreted ages of lithosphere based on Nd model ages are after Champion and Cassidy (2010) and Champion et al. (2006). Figure 2: Framework of Neoarchaean metamorphic events based on schematic evolution of metamorphic variables such as T, P and T/depth ratio, with time progressing to the right. Juxtaposed on this framework are crust formation events such as volcanism, plutonism and sedimentary basins, and the deformation events recognised in the EYC. The pre-2685 Ma period is a schematic representation of multiple magmatic arc docking events. The metamorphic framework is based primarily on the EYC, though M2, M3 and M4 periods are applicable to the whole craton. Sedimentary basins: SVB – syn-volcanic clastic basin, PVTB – post-volcanic turbiditic basin, and PVFB – post-volcanic fluvial basin. Figure 3: Summary of peak metamorphic PT fields and representative P-T paths for the different metamorphic events recognized in the Yilgarn craton. Figure 4: Time-space diagram for the Yilgarn Craton, including summaries of stratigraphy, gold mineralization, granite ages and interpretive metamorphic events, based on currently accepted geological history (Blewett and Czarnota, 2007c; Blewett et al., 2010; Korsh et al., 2013). Stratigraphic nomenclature based on that currently accepted by GSWA. Age range of rock units are constrained by volcanic age determinations and detrital zircon maximum deposition ages in literature and from this study (Table 1; Appendix 1; e.g. Groenewald et al., 2006; Kositicin et al., 2008; Van Kranendonk and Ivanic, 2009; etc). Metamorphic events are constrained by stratigraphy, granite peaks and metamorphic age determinations (Figure 32). Deformation mode is indicated by different coloured metamorphic events: either contraction (blues) or extension (red). A near comprehensive compilation of granitoid age determinations from the literature (Appendix 1) is presented as frequency curves (greys), with the number of determinations indicated. Approximate ages of all known mafic-ultramafic intrusives are presented as dark green bands. A time-space diagram containing metamorphic age data is available in Figure (32). Figure 5: Interpretive map of early metamorphic events (>2700 Ma) in the Yilgarn craton. All probable Ma granulite and M1 high-P metamorphic parageneses, PT determinations and age determinations have been plotted. All Ma and M1 data has been sourced from this study (Goscombe et al., 2009), except (#) in the Marymia Inlier (Vielreicher et al., 2002). M1 parageneses have been identified by: [1] anomalously high-P >7.0 kbar, [2] Barrovian series high-P/moderate-T average thermal gradients <27 ºC/km, such as epidote amphibolite and garnet±clinopyroxene±grunerite amphibolite facies, and [3] old metamorphic age determinations between 2697-2748 Ma. M1 parageneses are spatially restricted, typically within shear zones and reworked by lower-pressure M2 parageneses. M1 parageneses pre-date the high-Ca granite bloom and associated M2 regionalcontact metamorphism. M1 parageneses are often associated with older stratigraphic units indicated by mafic volcano-stratigraphy dated at >2730 Ma (yellow hexagons). Ma parageneses are rare and identified as low-P/high-T granulite assemblages that are associated with the Gindalbie magmatic arc and HFSE granites, or within older lithosphere in the Burtville Terrane. The distribution of old lithosphere is defined by Nd depleted mantle model ages (Champion et al., 2006; Champion and Cassidy, 2010), where (+) indicates old lithosphere >3000 Ma and (-) indicates juvenile lithosphere <2850 Ma. Interpretation of the Gindalbie magmatic arc is after Barley et al. (2008). Note that only Neoarchaean metamorphic events have been mapped in the Yilgarn craton (Figures 5, 7, 8 and 9). Older metamorphic events are poorly documented and based only on age determinations from isolated mineral grains with no associated metamorphic parageneses reported. All Mesoarchaean and Palaeoarchaean metamorphic age determinations available in the literature (#) have been plotted and all are restricted to the WYC (Nieuwland and Compston, 1981; Williams and Myers, 1987; Kinny et

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al., 1988; Wang et al., 1998; Mueller and McNaughton, 2000; Wilde, 2001; Wilde and Spaggiari, 2007; Iizuka et al., 2008). Contact metamorphic hornfels (green) around large, early (>2710 Ma) mafic-ultramafic intrusive complexes such as the Windimurra complex have also been plotted (Wang, et al., 1998; Nelson, 2001; Ivanic et al., 2010). At least six different Proterozoic events have variably reworked and reactivated the margins of the craton; these are ignored in this study. Figure 6: Interpretive bulk strain map for the EYC based on qualitative foliation intensity (Goscombe and Gray, 2008) interpreted from rock descriptions in GSWA and GA field databases. Approximately n=27,000 sites. Map represents accumulative strain over multiple deformation events, though largely represents strain pattern during D2 accompanying termination of volcanism at ~2670-2675 Ma. High strain zones were repeatedly reactivated during D3, D4 and D5 events. Foliation intensity fields modified after Goscombe and Gray (2008): blue – essentially unstrained, mostly in granites, green – low strain cleavage or weak grain shape fabric, yellow – weak spaced schistosity or moderate grain shape fabric, orange – continuous schistosity or foliated gneiss, red – strongly schistose, and purple – high strain foliation and grain refinement. Figure 7: Interpretive metamorphic map of M2 regional-contact, low-P/moderate-T Buchan series metamorphic parageneses formed in associated with the high-Ca granite bloom. M2 metamorphism occurred between 2665-2685 Ma, after M1 arc accretion and old mafic volcano-stratigraphy, and before M3 extensional rift basins. M2 metamorphism and high-Ca granite bloom is of broadly the same age in all parts of the craton except the Narryer Terrane (Figure 4). Metamorphic map pattern represents peak metamorphic conditions attained at that locality, which may be due to one of a number of different metamorphic events. Nevertheless, the metamorphic pattern is dominated by M2 regional-contact metamorphism with M3 thermal anomalies superimposed in the vicinity of extensional late basins and extensional shear zones (Figure 8). M2 parageneses have been overprinted and obscured by later, higher-grade M4 parageneses in the Narryer and Southwest Terranes and western Youanmi Terrane (Figure 9). M2, M3 and M4 metamorphic map patterns (Figures 7, 8 and 9) have been constrained by the following data sources as outlined in Goscombe et al. (2009) and Goscombe and Blewett (2009). [1] Metamorphic grade based on ~1,200 quantitative PT determinations and detailed petrology undertaken for this study (Goscombe et al., 2007, 2009). [2] Reinterpretation of metamorphic facies on the basis of ~53,000 field sites and ~3,000 thin section descriptions in GSWA and GA databases (Goscombe et al., 2007, 2009). [3] Metamorphic grade based on published metamorphic maps and petrology and PT determinations in literature (e.g. Binns et al., 1976; Archibald et al., 1978; Purvis, 1978, 1984; McQueen, 1981; Phillips and Groves, 1982; Bickle and Archibald, 1984; Hallberg, 1985; Spray, 1985; Ahmat 1986; Ho, 1987; Phillips, 1986; Neall and Phillips, 1987; Williams and Myers, 1987; Clark et al., 1989; Ho et al., 1990; Watkins and Hickman, 1990; Swager et al., 1990; Witt, 1993, 1998; Ridley, 1992, 1993; Williams and Whitaker, 1993; Williams and Currie, 1993; Knight et al., 1993, 1996, 2000; Wang et al., 1993; Swager, 1994a,c, 1997; Bloem et al., 1994, 1997; Bloem, 1994; Dalstra, 1995; Witt and Davy, 1997; Mikucki, 1997; Wilkins, 1997; Hall, 1997; Swager and Nelson, 1997; Ridley et al., 1997; Cassidy et al., 1998; Dalstra et al., 1998, 1999; Rennie, 1998; Groenewald et al., 2000; Painter and Groenewald, 2001; Dugdale and Hagemann, 2001; Wilde, 2001; Mikucki and Roberts, 2003; Hagemann and Luders, 2003; White et al., 2003; Roberts, 2004; Groenewald et al., 2006; Evans et al., 2006; Elmer et al., 2006; Hodge, 2007; Wilde and Spaggiari, 2007; Spaggiari, 2007; Spaggiari et al., 2008; Thebaud and Miller, 2009; Goscombe et al., 2007, 2009; and many others listed in Appendix (6)). Annotated M2 metamorphic PT determinations are by THERMOCALC and geothermobarometry calculations sourced from this study (Table 7), except (#) indicates data for Marymia Inlier from Vielreicher et al. (2002). M2 metamorphic age determinations are sourced from this study (Table 1) and literature (unpublished CET data; Bagas, 1999; Van Kranendonk, 2007; Rennie, 1998). Figure 8: Interpretive metamorphic map for M3 low-P/moderate-T Buchan series metamorphism with anticlockwise P-T paths, associated with extensional shear zones and late-stage extensional clastic basins. M3 metamorphism occurred between 2650-2665 Ma, after regional-contact M2

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metamorphism, early high-Ca granites and main phase contractional deformation, and immediately after deposition of post-volcanic clastic basins. M3 metamorphic parageneses are confidently identified in post-volcanic clastic basin sequences with detrital zircon maximum deposition ages of; <2670 Ma in White Flag Formation, <2667 Ma in turbiditic Belches Formation, and probable <2675 Ma in turbiditic Diemal Formation. M3 metamorphic localities are also based on: [1] parageneses indicating probable anticlockwise P-T paths, [2] samples showing evidence for a second lowP/moderate-T metamorphic event, and [3] direct metamorphic age determinations in the age range 2665-2650 Ma. Red lines outline the probable extent of M3 thermal anomalies and extended upperplate domains associated with rifting. Flanks of lower-plate domes coincide with uplifted high-P M1 parageneses (Figure 5) and domal cores commonly coincide with low-Ca granites. Low-Ca granite crustal melt and mantle-derived syenite and mafic granite are also associated with M3 lithospheric extension. All M3 metamorphic PT determinations are sourced from this study (Table 7). M3 metamorphic age determinations are sourced from this study (Table 1), and (#) indicate data from literature (GSWA geochronology database; unpublished UWA data; Sircombe et al., 2007; McMillan, 1996; Schiotte and Campbell, 1996; Rennie, 1998; Wilde, 2001; Rasmussen et al., 2010). See Figure (7) for all other details and data sources. Figure 9: Interpretive metamorphic map for M4 low-P/high-T metamorphism with clockwise P-T paths, associated with a regionally extensive late-stage thermal pulse accompanying crustal shortening. M4 metamorphism occurred in multiple events through a protracted period between 2610-2650 Ma. Metamorphism is interpreted as a delayed thermal anomaly resulting from lowercrust and mantle lithosphere delamination triggered by M3 lithospheric extension. The metamorphic anomaly grades from lower amphibolite in central Youanmi Terrane to granulite grades in Narryer and Southwest Terranes. M4 metamorphic parageneses are identified by: [1] direct metamorphic age determinations in the age range 2610-2650 Ma, [2] samples showing evidence for a second lowP/high-T metamorphic event, [3] late-stage metamorphism with clockwise P-T path, and [4] metamorphic parageneses developed in late-stage clastic sequences with detrital zircon maximum deposition ages <2665 Ma. Low-Ca granite plutons derived from crustal melting are associated with the M3-M4 metamorphic anomalies (Figure 4). All M4 metamorphic PT determinations are by THERMOCALC and geothermobarometry calculations sourced from this study (Table 7), except (•) indicates PT estimates based on the stability field of matrix assemblages where PT calculations have been re-equilibrated during cooling. M4 metamorphic age determinations are sourced from this study (Table 1), and (#) indicate data from literature (GSWA geochronology database; Fletcher and McNaughton, 2001; Sircombe et al., 2007; Mueller and McNaughton, 2000; Mueller et al., 2004; Dalstra et al., 1998; Thebaud and Miller, 2009; VanKranendonk, 2008; Nemchin et al., 1994; Pigeon et al., 1990; Nemchin and Pigeon, 1999; Campbell et al., 1998; Yumin et al., 1998; Wilde, 2001; Muhling et al., 2008). Metamorphic age determinations <2610 Ma are interpreted as isotopically disturbed during protracted cooling and labelled (reset); these are sourced from this study (Table 1) and literature (Rennie, 1998; Nemchin et al., 1994; Pigeon et al., 1990; Nemchin and Pigeon, 1999). See Figure (7) for all other details and data sources. Figure 10: Microphotos of high-P M1 and low-P/high-T Ma parageneses in mafic and metapelite samples from the Kalgoorlie, Kurnalpi and Burtville Terranes: (a) Sample BG06-178 with gneissic banding defined by variation in clinopyroxene-hornblende proportions and plagioclase-quartz, and discordant low-angle granite vein. (b) Sample BG06-159k with aligned grainshape fabric in coarsegrained polygonal granoblastic matrix. (c) Sample BG06-171b high-P garnet-clinopyroxene amphibolite sample with aligned polygonal granoblastic matrix. (d) Sample BG06-171i with matrix assemblage of garnet, hornblende and grunerite laths. (e) Sample BG06-146a polygonal granoblastic mafic granulite with orthopyroxene-clinopyroxene-hornblende-plagioclase Ma assemblage. Note the ghosted ophitic texture indicating static metamorphism of ophitic gabbro protolith. (f) Sample BG06-183a high-P metapelite with early-formed staurolite grains over grown by post-kinematic garnet porphyroblast. All images are in plain polarized light. Field of view is 1.8 mm, except 0.9 mm in (c). All mineral abbreviations are after Kretz (1983).

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Figure 11: Microphotos of medium-grade M2 and M3 parageneses in metapelite samples from the Kalgoorlie, Kurnalpi and Youanmi Terranes: (a) Sample 98957 from pre-late basin stratigraphy in Kalgoorlie Terrane, with M2 parageneses of small post-kinematic staurolite overprinted by garnet porphyroblast, and with late-stage cordierite present in the rock. (b) Sample Y530b from Belches Formation, with M3 parageneses of small garnet idioblast and secondary biotite laths overprinting the main foliation. (c) Sample Y242a from aluminous schists at Kyanite Hill; M3 parageneses of early boudinaged and enveloped andalusite porphyroblast. (d) Sample Y242b from aluminous schists at Kyanite Hill; M3 parageneses of late skeletal andalusite growth on the margin of postkinematic kyanite lath. (e) Sample BG06-197c from Ravensthorpe belt with M3 parageneses of early cordierite porphyroblast enveloped by biotite-sillimanite main foliation. (f) Sample BG06-198c from Ravensthorpe belt with M3 parageneses of early, resorbed chloritoid aligned in chlorite-biotite foliation and late-stage post-kinematic staurolite porphyroblast. All images are in plain polarized light. Field of view is 4.2 mm in (b, c, f), 1.8 mm in (e) and 1.5 mm in (a, d). All mineral abbreviations are after Kretz (1983). Figure 12: Microphotos of medium-grade M4 parageneses in metapelite samples from the Youanmi Terrane: (a) Sample Y400a from Southern Cross belt, skeletal staurolite aligned within the main foliation that also contains biotite, quartz and graphite. Matrix foliation is over-grown by postkinematic garnet, and with late-stage andalusite in sample. (b) Sample Y400c from Southern Cross belt, partial view of a very large, late-stage andalusite porphyroblast that over-grew smaller postkinematic garnet idioblasts. (c) Sample LXB106782 from Cheritons belt, metapelite with quartz, biotite, staurolite and tourmaline matrix over-grown by garnet porphyroblast. (d) Sample BG09133a from Cheritons belt, metapelite with main foliation assemblage of quartz, biotite, graphite and relict staurolite, overprinted by post-kinematic garnet porphyroblasts and late-stage andalusite moats replacing staurolite. (e) Sample 198197a from Murchison domain, staurolite inclusion within latestage andalusite porphyroblast, and early foliation of ilmenite, biotite, quartz and muscovite within both the staurolite and andalusite. (f) Sample BG09-165a from Ravensthorpe belt, early small garnet idioblasts occluded within cordierite (now pseudomorphed) and andalusite porphyroblasts that overgrew the quartz, biotite, ilmenite, graphite and plagioclase matrix foliation. All images are in plain polarized light except cross-polarized light in (e). Field of view is 4.2 mm in (b, c, d, f), 1.5 mm in (a) and 0.9 mm in (e). All mineral abbreviations are after Kretz (1983). Figure 13: Microphotos of high-grade M4 parageneses in metapelite granulites from the Narryer and Southwest Terranes: (a) Sample BG10-39d from Lake Grace belt, orthopyroxene-garnet-quartz granulite with garnet porphyroblasts enveloped by weak orthopyroxene-biotite foliation. (b) Sample BG09-158c from eastern Southwest Terrane, meta-psammopelite granulite with quartz, plagioclase, biotite, partial melt and garnet matrix assemblage and secondary plagioclase-biotite coronas on garnet. (c) Sample BG10-8b from Jimperding belt, relict garnet resorbed by cordierite and biotite moat. (d) Sample BG10-10d from Jimperding belt, metapelite granulite with matrix garnet, orthopyroxene, phlogopite and plagioclase. (e) Sample BG10-69i from Narryer Terrane, relict garnet resorbed by matrix cordierite containing sillimanite inclusions. (f) Sample BG10-70i from Narryer Terrane, garnet porphyroblast with cordierite moat. All images are in plain polarized light. Field of view is 4.2 mm in (a, b, d, e, f) and 1.8 mm in (c). All mineral abbreviations are after Kretz (1983). Figure 14: Representative compositional maps of garnet porphyroblasts in typical metapelite samples in different terranes from west to east across the Yilgarn craton. All scale bars are 1 mm. Figure 15: Bulk composition of metamorphic samples from across the Yilgarn craton: (a) metapelites plotted in AFM space with excess muscovite, and (b) mafic samples plotted in ACF space. Labelled samples are XRF bulk compositions undertaken for this study. Colour outlines indicate the compositional field of typical metapelite samples from different regions. Open squares are bulk compositions of published PT pseudosections utilized in this study. Typical mineral compositions and tielines are plotted as visual guides.

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Figure 16: Generic, partial petrogenetic grid used for representing M1 parageneses in deeply buried mafic samples and Ma mafic granulites from the EYC. The limits of M1 garnet±clinopyroxene amphibolite assemblages have been drawn from a number of pseudosections covering a range of tholeiite bulk compositions that overlap with EYC samples (Figure 15), as listed: [1] Limit of orthopyroxene is from a pseudosection for moderately aluminous tholeiite in the NCKFMASHTO system with excess quartz (Groppo et al., 2007, figure 9). [2] Limits of clinopyroxene and epidote are from a pseudosection for sub-aluminous, low-Fe MORB tholeiite in the NCFMASH system with excess quartz and water (Diener et al., 2007, figure 18a). Limits of garnet-bearing assemblages are from a range of pseudosections and experimental studies for typical sub-aluminous tholeiite with moderate XMg range of 0.35-0.65 in the CFMASH and NCFMASH±T systems with excess quartz and water: [3] Pattison, 2003, [4] Green and Ringwood, 1967, [5] Mahan et al., 2008 (figure 11d), and [6] Clarke et al., 2006 (figure 4a). [7] Limit of garnet-hornblende-grunerite and garnet-gruneritehornblende-clinopyroxene assemblages are from a petrogenetic grid for Fe-rich amphibolite in the CFASH system with excess quartz and water (Zeh et al., 2005). Because phase relationships have not been calculated specifically for EYC samples, these generic stability fields are only utilized as guides. Bulk composition of the utilized pseudosections and EYC mafic samples are listed in the insets on the right. Sequence of metamorphic mineral growth in a select number of EYC mafic samples is also listed. Garnet compositional maps from grunerite-amphibolite illustrate compositional zoning during M1 garnet growth; showing decreasing Mn, increasing Ca and near flat Fe and Mg, indicating steep loading trajectories along the compositional isopleths. Scale bar is 1 mm. Black dots – peak metamorphic PT determinations based on garnet rims and matrix mineral cores, with error ellipses for average uncertainty of ±35 ºC and ±0.8 kbar. Red dots – prograde PT determination based on garnet core and inclusion parageneses. Blue dots – post-peak PT determinations based on garnet outer rims and rims of matrix minerals. Circles – indicate documented mineral parageneses constraining the P-T path and plotted with respect to mineral fields in the petrogenetic grid. Black box – results from Si-in-chlorite and calcite geothermometers, with pressures estimated. All mineral abbreviations are after Kretz (1983). Figure 17: Summary of peak metamorphic conditions and P-T evolutions documented for Ma granulite facies static metamorphosed magmatic arc gabbros and M1 high-P parageneses from major shear zones and terrane boundaries. (a) Ma and M1 parageneses from Burtville, Kurnalpi and Kalgoorlie terranes in the EYC. (b) Early-formed high-P parageneses of indeterminate age from the Southern Cross and Ravensthorpe belts in the WYC. These sample populations are interpreted to characterize the general P-T evolution of the early metamorphic events Ma and M1. Black dots – are peak metamorphic PT determinations with error ellipses based on the average error for the sample population. Red dots – are prograde PT determinations from garnet core and inclusion assemblages. Blue dots – are post-peak PT determinations from mineral outer rims. The semi-quantitative P-T paths interpreted for different samples are defined by prograde, peak and post-peak PT determinations, in conjunction with sequence of mineral growth compared with generic petrogenetic grids (see text; Figure 16). Figure 18: Summary of peak metamorphic conditions and P-T evolutions documented for M2 regional-contact metamorphic parageneses with clockwise P-T evolutions. M2 parageneses from: (a) Kurnalpi and Kalgoorlie Terranes, and (b) northeast Youanmi Terrane. These sample populations characterize the generalized P-T evolution of M2 metamorphism. Black dots are peak metamorphic PT determinations with error ellipses based on the average error for the sample population. Red dots are prograde PT determinations from garnet core isopleths. The semi-quantitative P-T paths interpreted for different samples are defined by prograde, peak and post-peak PT determinations, in conjunction with sequence of mineral growth compared with published PT pseudosections (see text; Figure 19). Red circles – are granitoid crystallization depths based on Al-in-hornblende geobarometers, typical errors are 0.5-1.0 kbar and crystallization temperatures of ~750 ºC are assumed.

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Figure 19: Generic petrogenetic grid used for the interpretation of aluminous low-Fe metapelite samples with M2 parageneses showing clockwise P-T paths from the EYC. PT pseudosection in MnNCKFMASH space with excess quartz and H2O, and composition above the garnet-chlorite tie line and low-Fe side of the biotite-staurolite tie line is sourced from Martinez et al. (2004, figure 7). Because the pseudosection has not been calculated for specific EYC samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). Sequence of metamorphic mineral growth in a select number of samples is listed at the top. Bulk composition of the PT pseudosection and EYC sample Y667a are listed. Whole-rock analyses do not exist for a number of EYC aluminous metapelites, which share similar modal compositions to samples with know bulk composition. Garnet compositional maps illustrate typical compositional zoning during M2 garnet growth of decreasing Mn, and increasing Ca, Fe and Mg towards rims, indicating loading with heating prograde trajectories. Scale bar is 1 mm. Black dots – peak metamorphic PT determinations based on garnet rims and matrix mineral cores, with error ellipses for average uncertainty of ±35 ºC and ±0.8 kbar. Red dots – prograde PT estimates based on garnet core compositional isopleths. Blue dots – post-peak PT determinations based on garnet outer rims and rims of matrix minerals. Circles – indicate documented mineral parageneses constraining the P-T path and plotted with respect to mineral fields in the pseudosection. Black box – Si-inchlorite geothermometer results from retrograde chlorite, with pressures estimated. All mineral abbreviations are after Kretz (1983). Figure 20: Generic petrogenetic grid used for the interpretation of typical metapelite samples with M3 parageneses showing anticlockwise P-T paths from the late turbiditic basin, Belches Formation. PT pseudosection in MnNCKFMASH space with excess quartz and H2O, and composition below the garnet-chlorite tie line and low-Fe side of the biotite-staurolite tie line is sourced from Martinez et al. (2004, figure 6). Because the pseudosection has not been calculated for specific EYC samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). Garnet compositional maps illustrate typical compositional zoning during M3 garnet growth showing typically flat compositional profiles and small idioblastic garnets with weak decrease in Mn, and increase in Mg towards rims. Scale bar is 1 mm. See caption of Figure (19) for all other details. Figure 21: Summary of peak metamorphic conditions and P-T evolutions documented for M3 extension-related metamorphic parageneses with anticlockwise P-T evolutions. M3 parageneses from: (a) post-volcanic turbiditic rift basin, Belches Formation in the Kurnalpi Terrane, (b) samples with anticlockwise P-T paths from the hanging wall and within the extensional Ockerburry shear zone between Kalgoorlie and Kurnalpi Terranes, and (c) parageneses showing anticlockwise P-T paths in the Youanmi Terrane, such as the Sandstone, Ravensthorpe, Southern Cross and Murchison belts, and Diemal Formation post-volcanic late clastic sequences. Black dots – are peak metamorphic PT determinations with error ellipses based on the average error for the sample population. Red dots – are prograde PT determinations from garnet core isopleths. Typical aluminosilicate phase relationships in aluminous schists are indicated in blue. The semi-quantitative P-T paths interpreted for different samples are defined by prograde, peak and post-peak PT determinations, in conjunction with sequence of mineral growth compared with published PT pseudosections (see text; Figures 20 and 23). Figure 22: Schematic metamorphic field gradients illustrating spatial variation in metamorphic response (T, P and T/depth) across the Yilgarn craton. Two generalized profiles are plotted; across the north (dashed lines) and south (solid lines), and organized around distinct tectono-metamorphic zones bounded by major shear zones. Field gradient curves are plotted through maximum peak metamorphic conditions irrespective of the age of metamorphism, and thus diachronous. Most metamorphic parageneses in the EYC are associated with M2 regional-contact metamorphism. Otherwise, localized Ma, M1 and M3 parageneses are indicated where known. Most metamorphic parageneses in the WYC are associated with the M4 metamorphic anomaly. All P, T and T/depth

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data are sourced from this project as listed in Table (7) and discussed in text. Errors are ignored for clarity, with average errors being ±35 ºC, ±0.8 kbar and ±6.0 ºC/km. Figure 23: Generic petrogenetic grid used for the interpretation of Fe-rich sub-aluminous metapelite samples with M3 parageneses showing early anticlockwise P-T paths from the Ravensthorpe belt. PT pseudosection in MnNCKFMASHTO space with excess quartz and H2O, and composition below the garnet-chlorite tie line and high-Fe side of the biotite-staurolite tie line is sourced from Tinkham and Ghent (2005). Because the pseudosection has not been calculated for specific Ravensthorpe samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). See caption of Figure (19) for all other details. Figure 24: Summary of peak metamorphic conditions and P-T evolutions documented for matrix forming regional metamorphic M4 parageneses with clockwise P-T evolutions from: (a) Ravensthorpe belt and (b) Southern Cross Belt in the western Youanmi Terrane. Most parageneses from the western Youanmi Terrane have formed during the M4 metamorphic event, which experienced similar PT conditions and clockwise P-T evolutions to M2. The semi-quantitative P-T paths interpreted for different samples are defined by prograde, peak and post-peak PT determinations, in conjunction with sequence of mineral growth compared with published PT pseudosections (see text; Figures 25 and 26). See Figure (18) for all other details. Figure 25: Generic petrogenetic grid used for the interpretation of Fe-rich sub-aluminous metapelite samples with M4 parageneses showing late clockwise P-T paths from the Ravensthorpe belt. PT pseudosection in MnNCKFMASHTO space with excess quartz and H2O, and composition below the garnet-chlorite tie line and high-Fe side of the biotite-staurolite tie line is sourced from Tinkham and Ghent (2005). Because the pseudosection has not been calculated for specific Ravensthorpe samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). Note sample BG09-164b records parageneses from two metamorphic events; early low-P cordierite growth presumably during M3 anticlockwise metamorphism and late clockwise metamorphism during M4 thermal overprint. Garnet compositional maps illustrate typical growth zoning during M4, of decreasing Ca and Mn, and increasing Fe and Mg towards rims. Scale bar is 1 mm. See caption of Figure (19) for all other details. Figure 26: Generic petrogenetic grid used for the interpretation of Fe-rich sub-aluminous metapelite samples with M4 parageneses showing late clockwise P-T paths from the Southern Cross belt. PT pseudosection in MnNCKFMASHTO space with excess quartz and H2O, and composition below the garnet-chlorite tie line and high-Fe side of the biotite-staurolite tie line is sourced from Tinkham and Ghent (2005). Because the pseudosection has not been calculated for specific Southern Cross samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). Garnet compositional maps illustrate typical growth zoning during M4, of decreasing Ca and Mn towards rims, indicating decompressive heating during garnet growth. Scale bar is 1 mm. See caption of Figure (19) for all other details. Figure 27: Summary of peak metamorphic conditions and P-T evolutions documented for highgrade M4 regional metamorphic parageneses with clockwise P-T evolutions. M4 parageneses from: (a) Narryer Terrane, (b) Jimperding belt, and (c) Lake Grace belt and southeast Southwest Terrane. Black dots are peak metamorphic PT determinations with error ellipses based on the average error for the sample population. Red dots are prograde PT determinations from garnet core isopleths. Blue dots are post-peak PT determinations. The semi-quantitative P-T paths interpreted for different samples are defined by prograde, peak and post-peak PT determinations, in conjunction with sequence of mineral growth compared with published PT pseudosections (see text; Figures 28 to 30).

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Figure 28: Generic petrogenetic grid used for the interpretation of Mg-rich sub-aluminous metapelite samples with M4 parageneses showing clockwise P-T paths from the Jimperding belt. PT pseudosection in NCKFMASHTO space with excess ilmenite, and composition below the garnetchlorite tie line and low-Fe side of the biotite-staurolite tie line is sourced from Longridge et al. (2017, figure 6). Bulk composition of the pseudosection is expressed as mol%. Because the pseudosection has not been calculated for specific Jimperding samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). Ca compositional map illustrates thin plagioclase corona on the margin of resorbed peak metamorphic garnet porphyroblast. Scale bar is 1 mm. See caption of Figure (19) for all other details. Figure 29: Generic petrogenetic grid used for the interpretation of Mg-rich sub-aluminous metapelite samples with M4 parageneses showing clockwise P-T paths from the Lake Grace belt. PT pseudosection in NCKFMASHTO space with excess ilmenite, and composition below the garnetchlorite tie line and low-Fe side of the biotite-staurolite tie line is sourced from Longridge et al. (2017, figure 6). Bulk composition of the pseudosection is expressed as mol%. Because the pseudosection has not been calculated for specific Lake Grace samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). See caption of Figure (19) for all other details. Figure 30: Generic petrogenetic grid used for the interpretation of Mg-rich sub-aluminous metapelite samples with M4 parageneses showing clockwise P-T paths from the Narryer Terrane. PT pseudosection in MnNCKFMASH space with excess quartz, plagioclase and water, and composition below the garnet-chlorite tie line and low-Fe side of the biotite-staurolite tie line is sourced from Tinkham et al. (2001, figure 5a). Because the pseudosection has not been calculated for specific Narryer Terrane samples, these phase relations are utilized only as a generic petrogenetic grid for comparison of samples with similar compositions (Figure 15). See caption of Figure (19) for all other details. Figure 31: Interpretive alteration map for the EYC based on ~2655-2610 Ma alteration parageneses documented in the literature (see Figure 7 and text for sources) and rock descriptions in GSWA and GA field databases. Alteration fields as defined by Witt (1993): brown – upper amphibolite >560 ºC, red – middle amphibolite 500-560 ºC, dark orange – lower amphibolite 420-500 ºC, and pale orange – greenschist 220-420 ºC. Green hatch – carbonate alteration and veins, red outline – bulk compositions with significant loss on ignition, and purple outline – hydrous alteration and veins. Black – shear zones and low-Ca granites. Dark green polygons indicate contact metamorphic overprint near 2410 Ma mafic-ultramafic dykes, and other green polygons are Mesoproterozoic thermal overprint and reworking associated with the Albany-Fraser Orogen. Figure 32: Time-space diagram including all available metamorphic age determinations sourced from literature (n=62; Appendix 1) and this paper (n=52; Table 1). Only robust U-Pb-Th, Lu-Hf and Sm-Nd age determinations undertaken directly on metamorphic mineral phases such as monazite, zircon, xenotime, allanite, titanite or garnet have been presented. Pooled average age and standard deviation for different metamorphic events are indicated by thick red line and grey shading. Key stratigraphic and magmatic constraints on the different metamorphic events are indicated. Individual age determinations are listed in Appendix (1). The source of metamorphic age determinations are labelled as follows: [1] this study, [2] GSWA geochronology database, [3] unpublished CET data, [4] unpublished UWA data, [5] Fletcher and McNaughton, 2001, [6] Sircombe et al., 2007, [7] Mueller and McNaughton, 2000, [8] Wang et al., 1998, [9] Bagas, 1999, [10] McMillan, 1996, [11] Mueller et al., 2004, [12] Dalstra et al., 1998, [13] Thebaud and Miller, 2009, [14] Schiotte and Campbell, 1996, [15] Van Kranendonk, 2007, [16] Rennie, 1998, [17] Nemchin et al., 1994, [18] Pigeon et al., 1990, [19] Nemchin and Pigeon, 1999, [20] Campbell et al., 1998, [21] Yumin et al., 1998, [22] Wilde, 2001, [23] Muhling et al., 2008, [24] Rasmussen et al., 2010, [25] Kinny et al.,

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1988, [26] Iizuka et al., 2008, [28] McNaughton et al., 2005, [29] Bucci et al., 2004, [30] Roberts, 2004, [31] Salier et al., 2005, [32] Mueller and Nemchin, 2003. Figure 33: Metamorphic evolution curves for the Yilgarn craton, tracking variation in metamorphic parameters with time: temperature (T), pressure or crustal depth (P), and T/depth ratio or average thermal gradient (G). Only Neoarchaean evolution at the current exposure level is documented. There is no metamorphic data for the cryptic Mesoarchaean and Palaeoarchaean metamorphic events, which are ignored. Curves are hypothetical composites of the peak metamorphic conditions attained during all of the different metamorphic events documented in the terrane. These evolution curves do not track any single real sample or location, which may not have recorded all metamorphic events experienced by the wider terrane. These curves are gross simplifications of the full range of metamorphic conditions experienced in each terrane. For simplicity of presentation the curves seek only to trace through the peak metamorphic conditions documented in each terrane (Table 7). Different domains experienced different peak metamorphic conditions for the same event, and where possible these have been represented by different dashed curves to illustrate the range in metamorphic response during a single metamorphic event. For clarity, lower-grade conditions have not always been plotted, but would merely fill left of the curves, to the minimum metamorphic conditions documented in the domain. Note that seemingly localized metamorphic anomalies such as those recorded by M1 shear zones, M3 late basin sequences or M4 gold mineralization in the EYC, may be sampling more extensive metamorphic anomalies that are otherwise unrecorded by new mineral growth or remain unrecognised outside of these tectono-stratigraphic settings. Deformation history has been confidently established for the EYC by Blewett and Czarnota (2007a,b,c), and is correlated with the metamorphic history. Age placement and width of the metamorphic anomalies are constrained by the available peak metamorphic age determinations (yellow histogram) based on direct dating of metamorphic minerals such as garnet, monazite, zircon and titanite (Figure 32). Only peak metamorphic conditions are known (Table 7), and ambient metamorphic conditions before and after metamorphic anomalies are generally unknown. Ambient conditions of 200 ºC and 2 kbar have been chosen that represent typical depths and temperatures at the base of moderately deep sedimentary basin cover. An ambient average thermal gradient of 19 ºC/km has been chosen that is a tectonically significant boundary between advection and conduction dominated systems (Appendix 7). Though these assumed ambient conditions are only probable, the flat inter-metamorphic trace is to be understood as diagrammatic representation for the absence of data. Figure 34: Interpretive model for the crustal growth period in the EYC, with metamorphic parageneses indicated (blue). Probable tectonic settings are; Ma granulite within magmatic arcs, M1 high-P parageneses in partially subducted margins of accreted magmatic arcs, and seafloor alteration and volcanism in a back arc setting (see text). Figure 35: Tectonic model for the M2 period during regional-contact metamorphism and high-Ca granite bloom, interpreted to be associated with termination of subduction and volcanism, and overlapping with a period of crustal shortening (see text). Figure 36: Tectonic model for the M3 period during lithospheric extension, interpreted to have been induced by roll back of the subduction hinge after termination of subduction (see text). Figure 37: Diagrammatic model for M3 lithospheric extension, not to scale. Extension led to: postvolcanic late rift basins, listric growth fault and lower-plate domal architecture, input of mantlederived mafic granite melt and reduced-CO2-metal-rich fluids into the upper-crust, post-kinematic M3 thermal anomaly, crustal melting and low-Ca granites, dewatering and dehydration H2O-rich fluid released from late basins, horizontal fluid flow driven by thermal perturbation and vertical compressive stress. All of these outcomes together resulted in fluid mixing, precipitation and goldmineralization, which peaked from ~2665 Ma onward (Figure 4). Later gold-mineralization events are associated with remobilization during successive M4 thermal perturbations, emplacement of low-Ca granites and development of suitable mineralization sites by strike-slip and transpressive

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deformation events. Note that M3 lithospheric extension can result in exhumation of high-P M1 parageneses by domal uplift in the footwall, juxtaposing them at the same crustal level as M2 and M3 parageneses in the hanging wall, as currently exposed. Figure 38: Tectonic model for late-stage M4 metamorphism; interpreted as a delayed thermal anomaly resulting from lower-crust and mantle lithosphere delamination triggered by M3 lithospheric extension (see text). Figure 39: Metamorphic evolution curves for the Cambrian-Permian Tasmanides in Eastern Australia, tracking variation in T, P and average thermal gradient with time. Curves are hypothetical composites of the peak metamorphic conditions attained during the different metamorphic events. Refer to Figure (33) for further explanation and legend. Data sourced from literature (Goscombe unpublished data). Figure 40: Frequency distribution of average thermal gradient (T/depth ratio) of metamorphic response in Yilgarn craton compared with typical continent-continent collisional orogens such as the Himalayan and Pan-African orogens, and typical accretionary orogen such as the Tasmanides of Eastern Australia. Data for the Yilgarn craton from this study, Himalayan data from Goscombe et al. (2018), Pan-African and Tasmanide data from Goscombe (unpublished data). Note the thermal regime in the Yilgarn craton has similarities to modern accretionary orogens represented by the Lachlan Orogen, and contrast strongly to subduction or collisional orogens. Figure 41: Comparison of metamorphic response between [a] Yilgarn craton, [b] Archaean cratons globally, and [c] Phanerozoic metamorphic belts globally. Data sourced from this study, Goscombe et al. (2018) and Goacombe (unpublished data). Note J-curve metamorphic array for the EYC, indicating a thermally stratified crust due to neutral buoyancy pooling of voluminous high-Ca granite in the upper-crust; and lower thermal gradients in the mid-crust sampled by M1 parageneses exhumed in shear zones. Also note the absence of metamorphic conditions indicative of subduction zones or continent-continent collision in the Yilgarn craton. TABLE CAPTIONS:

Table 1: Table summarizing all new metamorphic age determinations in the Yilgarn craton undertaken for this study (Goscombe et al., 2007, 2009). Metamorphic age determinations were derived by four different methods: [1] In situ U-Pb analysis of metamorphic monazite grains in polished thin-sections of metapelite, using LA-ICPMS at Adelaide University and calculated as weighted mean Pb207-Pb206 ages undertaken by Dr Ben Wade. [2] In situ U-Pb analysis of metamorphic titanite grains in polished thin-sections of amphibolite, using MC-ICPMS at Florida University and calculated as Pb207-Pb206 isochron ages undertaken by Prof. David Foster. [3] Sm-Nd and Lu-Hf analysis of garnet, hornblende and clinopyroxene fractions from amphibolite, using ICPMS at Melbourne University and calculated as isochron ages undertaken by Prof. Roland Maas at Melbourne University. [4] U-Pb analysis of metamorphic and detrital zircon separates using the SHRIMP at Curtin University and calculated as weighted mean U-Pb ages undertaken by GSWA staff, with results available in the GSWA geochronology database. This metamorphic data has been synthesized with metamorphic, magmatic and maximum deposition age data in the literature (Appendix 1), to constrain the time-space diagram of stratigraphic, metamorphic and magmatic evolution of the Yilgarn craton (Figures 4 and 32). Geochronology methods, sample details and full results are presented in the text, Appendices (1 and 2) and Goscombe et al. (2009). ¥ - Interpretive correlation of metamorphic age determinations with the M1, M2, M3 and M4 Neoarchaean metamorphic events outlined in the text. Table 2: Summary of all Archaean metamorphic events recognised, at different levels of confidence, in the Yilgarn craton. This diagnostics table summarizes; (1) the salient characteristics and criteria that define the different metamorphic events, and (2) variation in conditions from east to west across

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the craton for each metamorphic event. Proterozoic events in marginal metamorphic belts such as the Albany Fraser, Pinjarra and Capricorn orogens, have variably reworked or reactivated margins of the Yilgarn craton; these are ignored in this study. Ø - Average and standard deviation of peak metamorphic conditions based on the best-constrained PT determinations using THERMOCALC (Table 7). n – Number of samples used to constrain average peak metamorphic conditions. § - Range of metamorphic age determinations with known mineral parageneses; sourced from this study and literature (Table 1; Appendix 1). ~ - Estimated or poorly constrained age determination from literature; mostly single age determinations from an isolated zircon or monazite grain, with associated mineral parageneses unknown. ◊ - Deformation state based on correlation with EYC tectonic framework (Blewett and Czarnota, 2005, 2007a,b; Blewett et al., 2010; Czarnota et al., 2010a), and inferred from burial versus exhumation trajectories constrained by P-T paths. # Diagnostic mineral parageneses and sequence of mineral growth; predominantly from metapelite rocks (Tables 3 to 6). Table 3: Petrology of representative Ma granulite and M1 high-P assemblages in predominantly mafic and one metapelite rock from different parts of the Yilgarn craton. Relative timing and textural setting of mineral phases are indicated: I – inclusion setting with host phase; E – exsolution; S – within main foliation; XS - aligned matrix phase; X – unaligned matrix phase or porphyroblast with number indicating relative timing; C – corona or reaction texture; L – late and discordant to the main foliation; F – late shear band or foliation seam; R – retrograde; P – pseudomorphed. All mineral abbreviations are after Kretz (1983). Additional abbreviations used in petrology tables (Tables 3 to 6), microphotographs (Figures 10 to 13) and P-T plots are: Op – opaque, Sill – sillimanite, L – partial melt, Feox – Fe oxides. Predominant mineral end-member is in brackets. Peak metamorphic conditions are based on THERMOCALC PT determinations undertaken in this study (Table 7; Appendix 5). [ ] – Indicates PT determinations from other samples at the same locality. • – Indicates PT determinations based on the phase stability field, where THERMOCALC result is inconsistent with the stability field of the matrix assemblage due to re-equilibration during cooling from granulite facies conditions. ◊ – Indicates PT determinations based on conventional geothermobarometry or phase stability constraints, where sample is significant and no other constraint is available. P-T paths: CW – clockwise, ACW – anticlockwise, ITL – isothermal loading, ITD – isothermal decompression, DC – decompressive cooling, IBC – isobaric cooling, IBH – isobaric heating, LH – loading heating. Table 4: Petrology of representative middle amphibolite facies metamorphic parageneses in metapelite rocks showing clockwise P-T paths. These parageneses formed during M2 regionalcontact metamorphism in the Kalgoorlie-Kurnalpi-Burtville Terranes, and in the Youanmi Terrane either formed during M2 regional-contact metamorphism or later during M4 metamorphism as indicated. # - Petrology interpreted from descriptions and sketches in literature (Bickle and Archibald, 1984; Mikucki and Roberts, 2003). See Table (3) for all other details. Table 5: Petrology of representative M3 metamorphic parageneses in metapelite rocks showing anticlockwise P-T paths, from different parts of the Yilgarn craton. See Table (3) for all other details. Table 6: Petrology of representative high-grade M4 metamorphic parageneses in metapelite and orthopyroxene granulites showing clockwise P-T paths, from the Narryer and Southwest Terranes. See Table (3) for all other details. Table 7: Table summarizing the best-constrained peak metamorphic PT determinations (n=287), limited to those calculated in this study using THERMOCALC v3.25 (Powell and Holland, 1994). Only robust PT determinations are listed; all data by conventional geothermobarometry, assumed P or T, phase stability constraints, estimates or otherwise poorly constrained results are ignored. All PT determinations from literature (n=422) are by conventional geothermobarometry and most assume a pressure of 4 kbar, and consequently ignored. Similarly, 53% of PT determinations

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undertaken by this study (n=330) used conventional geothermobarometry and have been ignored and not included in Tables (2 and 7). Full details of all calculations, by all methods and from all available samples including from literature, are contained in Appendix (5). Average thermal gradient (G) is calculated simply as temperature/depth (ºC/km), assuming a density of 2.8 g/cm3. Ø - Peak metamorphic PT determination that is the single best constrained THERMOCALC result calculated by the method listed. n – Indicates the number of additional independent PT determinations using either multiple mineral analyses, alternative methods in THERMOCALC, or conventional geothermobarometers; that overlap with and confirm the peak metamorphic result. • - Indicates an alternative PT determination has been used, where THERMOCALC average-PT result is inconsistent with the phase stability field of the matrix assemblage due to re-equilibration during cooling from granulite facies conditions. ◊ - Indicates a small number of PT determinations based on conventional geothermobarometry or phase stability constraints, where sample is significant and no other constraint is available. ¥ - Most PT determinations were undertaken using mineral analyses from this study, with a small number of PT calculations undertaken using mineral analyses available in the literature. Mineral analyses sources are numbered: (1) This study, (2) Gole and Klein, 1981; (3) Bickle and Archibald, 1984; (4) Mikucki, 1997; (5) Williams and Currie, 1993; (6) Mueller ans McNaughton, 2000; (7) Baga L. unpublished; (8) Dalstra, 1995; (9) Witt, 1998; (10) Dalstra et al., 1999; (11) Bloem et al., 1994; (12) Purvis, 1978; (13) Spray, 1985. § - Mineral analysis sites are organised by location within garnet porphyroblast followed by matrix minerals. Prograde PT determinations based on calculations from mineral cores undertaken in THERMOCALC, or estimates based on garnet core isopleths in Spear (1993) or Gaidies et al. (2008). Estimated fluid composition for aH2O, otherwise XH2O indicated by square brackets. Interpretive P-T paths are based on the sequence of mineral growth documented in thin section (Tables 3 to 6), and prograde, peak and post-peak PT calculations as discussed in text. P-T path abbreviations are listed in Table (3) caption. Samples are located and organised by: (1) easting and northings, (2) stratigraphic unit for post-volcanic late-basin samples, (3) metamorphic belt, and (4) terrane. Further location details and sample location maps are available in Appendices (3 and 5). Correlated metamorphic event is interpreted on the basis of criteria such as; (1) tectono-stratigraphic setting, (2) P-T path, (3) metamorphic conditions, (4) overprinting relationships, (5) assemblages, and (6) age determinations, as listed in the metamorphic maps (Figures 5, 7, 8 and 9) and discussed in text. Samples with a metamorphic age determination confirming the correlation are indicated by bold text (Table 1). BIOGRAPHICALS:

Ben Goscombe is currently Courtesy Professor with Department of Geology at University of Florida. Research interests are structural and metamorphic processes in a range of orogenic systems utilizing integrated structural-metamorphic-chronologic datasets. PhD at Melbourne University 1989; government geological surveys contracts in Zimbabwe, Namibia and Australia; Honorary Research Associate at Adelaide University 2000-2014; Research Associate at Melbourne University 2001-2003. Freelance researcher from 2006 and established Integrated Terrane Analysis Research (ITAR) to undertake self-funded research and collaborations with Geoscience Australia and University of Florida. Received Phaup Award in 1998 and 2000 for contributions to Zimbabwe Geology and Gondwana Research best paper award in 2007.

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David A. Foster is Professor and Chair of the Department of Geological Sciences, University of Florida. He studies the tectonic evolution of continents throughout earth history using high and low temperature thermochronology integrated with other geological and geophysical data. He is a Fellow of the Geological Society of America and was twice the recipient of the Stillwell award from the Geological Society of Australia

Richard Blewett is the General Manager of the Minerals Systems Branch at Geoscience Australia. He has responsibility for leading GA’s minerals science programme and the promotion of Australia as an attractive investment destination for minerals exploration. Richard graduated 1st class Hons in Geology from Swansea University (Wales) in 1985. Following a year in the seismic industry in South Africa, he completed a PhD in structural geology from Leicester University in the UK (1989). During this time he worked as a geologist in the French Alps, Canadian Appalachians, British Caledonides and Nepalese Himalaya. Richard joined Geoscience Australia in 1990 as a research scientist and for the past 28 years has worked in a number of minerals-related mapping projects across many of Australia’s mineral provinces. He combined this with work on joint projects in the Sultanate of Oman, China and is currently leading a project of institutional strengthening in India. Richard was the Chief Editor and leader of a major book (2012) on Australia’s geology – Shaping a Nation: a Geology of Australia. Richard has been involved in the development of the UNCOVER initiative of the Australian Academy of Science. Since 2012 he has been the leader of the Mineral Systems Branch in the Resources Division at Geoscience Australia, which has carriage of the minerals component of the new Exploring for the Future programme. Richard is interested in the management and leadership of science and in geoscience education. He has an MBA in Technology Management from Deakin University (2001).

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Karol Czarnota is currently Senior Geoscientist at Geoscience Australia and team leader of Australian lithosphere studies. Doctor of Philosophy (2010 to 2013) in Geology and Geophysics at University of Cambridge. Master of Science (2008 to 2009) in Petroleum Geoscience at Royal Holloway, University of London. Bachelor of Science (1999 to 2002) in Applied Geology at University of NSW, Australia, with first Class Honors and University Medal.

Benjamin Wade is a senior microscopist at Adelaide Microscopy in the University of Adelaide. He received his PhD (2006) from the University of Adelaide for research on the Musgrave Provence, Arunta Provence and Gawler Craton. He has worked in industry and held undertaken post-doctoral research from 2006 to 2008, and since has been employed at Adelaide Microscopy. His primary interests and research include the application of micro-analytical techniques such as SEM, TEM, EPMA, and LA-ICP-MS to geological problems.

Bruce Groenewald has published research on Proterozoic and Archaean crustal evolution, stratigraphy, and orogeny in Africa, Australia, and Antarctica. After lecturing for 7 years at the University of KwaZulu Natal, he worked at the Geological Survey of Namibia, and then the

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Geological Survey of Western Australia, where he compiled a seamless geological map and information system of the Eastern Goldfields Superterrane. Since moving to industry in 2008, he has explored for base metals, gold, and bauxite across Australia and Papua New Guinea.

David Gray is currently Associate Professor at the University of Tasmania (2012- ). Research interests are in structural geology and tectonics with particular focus on evolution of eastern Australia, Otago Schist belt in New Zealand, Damara Orogen in Namibia, and ophiolite obduction in Oman. Was an academic for 35 years at University of New South Wales (1975-1977), Virginia Tech in the USA (1977-1983) and Monash University (1983-2001), and Australian Professorial Fellowship at the University of Melbourne (2002-2006). Structural geology consultant to the minerals industry through his company Geostructures Pty Ltd (2006-2016). Awarded the Geological Society of Australia Selwyn Medal for his contributions to Victorian Geology (1997), Carey Medal for contributions to Structural Geology and Tectonics (1998), and Stillwell Award (1997 and 2004).

Sample

Domain Constraint

Narryer Terrane: BG10-70g Narryer Granulites weighted mean BG10-63a Narryer Granulites weighted mean BG10-63a Narryer Granulites concordia age BG10-61b Narryer Granulites weighted mean BG10-70a Narryer Granulites intercept zircon Palaeoarchaean meta, BG10-70a Narryer Granulites weighted mean BG10-70a Narryer Granulites weighted mean

Rock Type Mineral

Laboratory Genesis

Isotopic SystemMethod Age (Ma) 2 sd

metapelite granulite monazite metapelite granulite monazite metapelite granulite zircon metapelite granulite monazite metapelite granulite metamorphic

Adelaide University metamorphic Adelaide University metamorphic Florida University detrital Adelaide University metamorphic Florida University 3195

Pb207-Pb206 2626.4 Pb207-Pb206 2623.6 Pb207-Pb206 ~3322 Pb207-Pb206 2715 Pb207-Pb206 81

LA-ICPMS 8.1 LA-ICPMS 7.4 MC-ICPMS

metapelite granulite zircon metapelite granulite zircon

Curtin University detrital Curtin University metamorphic

Pb207-Pb206 3343 Pb207-Pb206 2659

SHRIMP 17 SHRIMP 3

89

LA-ICPMS 12 MC-ICPMS 15

n

Interpretation

17

Peak M4 metam

19

Peak M4 metam

1

Maximum depo

9 upper

Early Neoarcha

1

Maximum depo

60

M3 metamorph

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BG10-67 BG10-67 BG10-63b

Narryer Granulites weighted mean Narryer Granulites weighted mean Narryer Granulites weighted mean

Southwest Terrane: BG10-4c Jimperding Belt weighted mean BG10-4c Jimperding Belt concordia age BG10-4b Jimperding Belt weighted mean BG10-4b Jimperding Belt weighted mean BG10-35c Lake Grace Terrane weighted mean BG10-39d Lake Grace Terrane weighted mean BG10-39c Lake Grace Terrane weighted mean BG10-39c Lake Grace Terrane weighted mean BG10-39c Lake Grace Terrane weighted mean BG10-39a Lake Grace Terrane weighted mean BG09-130 East Transition age titanite disturbance BG10-41 North Transition age titanite disturbance BG10-45 North Transition age titanite disturbance Murchison Domain: 194524 Cue Belt age titanite Neoarchaean meta. BG10-59h Dalgaranga Belt weighted mean GSD073091 Dalgaranga Belt age titanite M4 metamorphism GSD083343 Koonmarra Belt age titanite metamorphism GSD081675 Tieraco Belt age titanite M4 metamorphism GSD081675 Tieraco Belt age titanite metamorphism GSD083357 Yalgoo Belt age titanite Neoarchaean meta. Southern Cross Domain: BG09-128 Cheritons Belt age titanite metamorphism

quartzite granulite zircon quartzite granulite zircon metapelite granulite zircon

Curtin University detrital Curtin University metamorphic Curtin University detrital

Pb207-Pb206 3340 Pb207-Pb206 3195 Pb207-Pb206 3241

SHRIMP 20 SHRIMP 18 SHRIMP 5

metapelite granulite zircon metapelite granulite zircon metapelite granulite zircon metapelite granulite zircon metapelite granulite monazite opx-grt granulite zircon opx-grt granulite zircon opx-grt granulite zircon opx-grt granulite zircon opx-grt granulite zircon amphibolite metamorphic

Florida University metamorphic Florida University detrital Curtin University detrital Curtin University metamorphic Adelaide University metamorphic Florida University metamorphic Curtin University detrital Curtin University metamorphic? Curtin University metamorphic Curtin University metamorphic Florida University 2607

Pb207-Pb206 ~2624 Pb207-Pb206 ~2642 Pb207-Pb206 ~2663 Pb207-Pb206 2641 Pb207-Pb206 2645 Pb207-Pb206 2638 Pb207-Pb206 2664-2687 Pb207-Pb206 2655 Pb207-Pb206 2643 Pb207-Pb206 2638 Pb207-Pb206 19

MC-ICPMS

amphibolite metamorphic

Florida University 2594

Pb207-Pb206 30

MC-ICPMS 21

isochron Isotopic

amphibolite metamorphic

Florida University 2571

Pb207-Pb206 13

MC-ICPMS 14

isochron Isotopic

meta-ultramafic metamorphic

Florida University 2728

Pb207-Pb206 32

MC-ICPMS 19

isochron Early

metapelite monazite calcsilicate rock meta-overgrowth

Adelaide University metamorphic Florida University 2631

Pb207-Pb206 2613.1 Pb207-Pb206 27

LA-ICPMS 9.3 MC-ICPMS 15

19 isochron Peak

amphibolite metamorphic

Florida University 2681

Pb207-Pb206 16

MC-ICPMS 19

isochron M2

amphibolite meta-overgrowth

Florida University 2644

Pb207-Pb206 16

MC-ICPMS 17

isochron Early

amphibolite metamorphic

Florida University 2667

Pb207-Pb206 6

MC-ICPMS 11

isochron M2

amphibolite metamorphic

Florida University 2724

Pb207-Pb206 13

MC-ICPMS 15

isochron Early

amphibolite metamorphic

Florida University 2664

Pb207-Pb206 10

MC-ICPMS 15

isochron M2

90

3

Maximum depo

1

Palaeoarchaea

1

Maximum depo

1

Peak M4 metam

1

Maximum depo

3

Maximum depo

3

Early M4 metam

28

Early M4 metam

4

Peak M4 metam

2

Maximum depo

5

M3 metamorph

1

Early M4 metam

27 isochron Isotopic

Peak M4 metam

MC-ICPMS SHRIMP SHRIMP 8 LA-ICPMS 8 MC-ICPMS 9.2 SHRIMP SHRIMP 11 SHRIMP 13 SHRIMP 2 MC-ICPMS 19

Peak M4 metam

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BG09-129 Cheritons Belt age titanite disturbance BG09-37b Diemals Belt age titanite Neoarchaean meta. BG10-46b Southern Cross Belt weighted mean BG09-71 Wiluna Belt age titanite M4 metamorphism Kalgoorlie Terrane: 98221 Nth Boorara Domain age titanite disturbance BG06-182b Nth Coolgardie Domain age titanite Mesoarchaean meta? BG06-190b Nth Coolgardie Domain age titanite deep burial in EYC BG06-190b Nth Coolgardie Domain age titanite deep burial in EYC, 1 BG06-190b Nth Coolgardie Domain age titanite deep burial in EYC Y667a Depot Domain age garnet disturbance Y667a Depot Domain age garnet deep burial in EYC BG06-162 Ora Banda Domain age titanite deep burial in EYC, 1 BG06-162 Ora Banda Domain age titanite deep burial in EYC Kurnalpi Terrane: Y534b Belches Basin age garnet disturbance Y534b Belches Basin age garnet age: zircon inclusions Y530a Belches Basin age garnet disturbance Y530a Belches Basin age garnet age: zircon inclusions Y455f Belches Basin weighted mean Y459d Belches Basin age titanite Maximum deposition age Y690a Edjudina Domain age garnet disturbance

amphibolite metamorphic

Florida University 2600

Pb207-Pb206 16

MC-ICPMS 12

isochron Isotopic

amphibolite metamorphic

Florida University 2782

Pb207-Pb206 37

MC-ICPMS 18

isochron Early

meta-pelite monazite amphibolite metamorphic

Adelaide University metamorphic Florida University 2632

Pb207-Pb206 2624 Pb207-Pb206 30

LA-ICPMS 11 MC-ICPMS 14

22 isochron Peak

amphibolite metamorphic

Florida University 2598

Pb207-Pb206 36

MC-ICPMS 17

isochron Isotopic

amphibolite metamorphic?

Florida University 2823

Pb207-Pb206 31

MC-ICPMS 12

isochron

amphibolite metamorphic

Florida University 2688

Pb207-Pb206 25

MC-ICPMS 30

isochron M1

amphibolite metamorphic

Florida University 2706

Pb207-Pb206 25

MC-ICPMS 30

isochron M1

amphibolite metamorphic

Florida University 2711

Pb207-Pb206 26

MC-ICPMS 29

isochron M1

metapelite metamorphic

Melbourne University Sm-Nd 2593 13

dissolution 3

isochron Isotopic

metapelite metamorphic

Melbourne University Lu-Hf 2712 13

dissolution 3

isochron M1

amphibolite metamorphic

Florida University 2702

Pb207-Pb206 30

MC-ICPMS 11

isochron M1

amphibolite metamorphic

Florida University 2702

Pb207-Pb206 60

MC-ICPMS 11

isochron M1

metapelite metamorphic

Melbourne University Sm-Nd 2547 17

dissolution 5

isochron Isotopic

metapelite metamorphic

Melbourne University Lu-Hf 2688 8

dissolution 5

isochron Mixed

metapelite metamorphic

Melbourne University Sm-Nd 2561 11

dissolution 5

isochron Isotopic

metapelite metamorphic

Melbourne University Lu-Hf 2695 9

dissolution 5

isochron Mixed

metapelite monazite calcsilicate rock detrital

Adelaide University metamorphic Florida University 2682

LA-ICPMS 8 MC-ICPMS 23

22 isochron

amphibolite metamorphic

Melbourne University Sm-Nd 2610 16

dissolution 3

isochron Isotopic

91

Pb207-Pb206 2649 Pb207-Pb206 25

Peak M4 metam

M3 rift basin me

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

Y690a Edjudina Domain age garnet deep burial Hootanui SZ BG06-174a Laverton Domain age titanite deep burial Hootanui SZ Y0360 Laverton Domain age titanite deep burial in EYC Y0360 Laverton Domain age titanite Mesoarchaean meta? BG06-121 Sth Linden Domain age titanite disturbance Burtville Terrane: BG6-151b Merolia Domain age titanite deep burial in Burtville

Thermal Event

Age § (Ma)

amphibolite metamorphic

Melbourne University Lu-Hf 2726 13

dissolution 3

isochron M1

amphibolite metamorphic

Florida University 2727

Pb207-Pb206 10

MC-ICPMS 29

isochron M1

amphibolite metamorphic

Florida University 2697

Pb207-Pb206 23

MC-ICPMS 17

isochron M1

amphibolite metamorphic?

Florida University 2824

Pb207-Pb206 18

MC-ICPMS 5

isochron

amphibolite metamorphic

Florida University 2589

Pb207-Pb206 34

MC-ICPMS 17

isochron Isotopic

calcsilicate rock metamorphic

Florida University 2748

Pb207-Pb206 19

MC-ICPMS 17

isochron M1

Terrane-Domain

Setting

Deformation ◊

Magmatism

Petrology #

Early metamorphic events: poorly constrained, identified by age determinations, parageneses typically not preserved P-Archaean ~3284-330 Narryer [unknown] [unknown] P-Archaean ~3180 Southwest [unknown] [unknown] M-Archaean 2853-2856 Youanmi-Southern Cross [unknown] [unknown] M-Archaean 2800-2815 Youanmi-Murchison Windimurra IC aureole static metamorphism mafic-ultramafic IC early N-Arch. ~2800 Southwest [unknown] [unknown] early N-Arch. ~2775 Narryer [unknown] [unknown] early N-Arch. 2772-2782 Youanmi-Southern Cross regional-contact meta? [unknown] TTG association N-Archaean 2724-2728 Youanmi-Murchison regional-contact meta? [unknown] high-Ca granite N-Archaean ~2700-2715 Narryer [unknown] [unknown]

[none preserved] [none preserved] [none preserved] Opx-Crd-Qtz hornfels [none preserved] [none preserved] Cpx amphibolite Cpx amphibolite [none preserved]

Ma metamorphism: early high-grade parageneses, probably related to magmatic arc metamorphism Ma [unknown] Burtville-Duketon arc small granulite domains static metamorphism Ma [unknown] Kurnalpi-Gindalbie arc high-grade domains static metamorphism

ophitic mafic granulite Cpx amphibolite

HFSE granite HFSE granite

M1 metamorphism: early high-P parageneses in shear zones, probably related to partial subduction of accreted arcs M1 2748 Burtville-Merolia high-P lenses in SZ's contraction & burial M1 2697, 2727 Kurnalpi high-P lenses in SZ's contraction & burial M1 2702-2712 Kalgoorlie high-P lenses in SZ's contraction & burial M1 [unknown] Youanmi-Southern Cross isolated high-P rocks contraction & burial M1 [unknown] Youanmi-Ravensthorpe isolated high-P rocks contraction & burial

Cpx amphibolite Grt-Cpx±Gru amph. Grt±Cpx amphibolite Cpx amphibolite Grt amphibolite

M2 metamorphism: Buchan series regional-contact metamorphism, related to high-Ca granite bloom in back-arc of a subduction system M2 [unknown] Burtville & Kurnalpi regional-contact meta. contraction high-Ca granite M2 2670-2675 Kalgoorlie regional-contact meta. contraction high-Ca granite M2 2664-2680 Youanmi-Southern Cross regional-contact meta. contraction high-Ca granite M2 2667-2681 Youanmi-Murchison regional-contact meta. contraction high-Ca granite M2 2670 Southwest [unknown] [unknown] high-Ca granite M2 [unknown] Narryer [unknown] [unknown] high-Ca granite

Cpx amphibolite Ctd > St-Grt > And > Crd Grt > And > Crd St > And > Crd [none preserved] [none preserved]

M3 metamorphism: localized high-T/low-P thermal anomalies, related to lithospheric extension and rifting in the back-arc on cessation of subduction M3 2649 Burtville [late basin sequences] extension > transpress high-Ca & mafic G. M3 2649-2656 Kurnalpi Belches late basin rift extension > transpress high-Ca & mafic G. M3 2663 Kalgoorlie Ockerburry extension SZ extension > transpress high-Ca & mafic G. M3 [unknown] Youanmi-Southern Cross ACW parageneses probable extension high-Ca granite M3 2660 Youanmi-Southern Cross Diemal late basin rift probable extension high-Ca granite M3 [unknown] Youanmi-Ravensthorpe ACW parageneses probable extension high-Ca granite M3 2653 Youanmi-Murchison ACW parageneses probable extension high-Ca granite M3 2644-2663 Southwest [late basin sequences] probable extension high-Ca granite

[none indentified] Crd > And > Grt > St Crd > Grt > St Crd-Bt-Ms > Grt Crd > And > Chl Ctd > And > St > Grt And±Crd > Grt > Ctd [none preserved]

92

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

M3

2653-2660

Narryer

[late basin sequences]

probable extension

high-Ca granite

[none preserved]

M4 metamorphism: high-T/low-P regional metamorphism accompanying contraction, probably related to delayed thermal pulse after lithospheric delamination M4 2610-2655 Kalgoorlie (literature) alteration, mineralization transpress>transtension low-Ca granites M4 2631-2638 Kalgoorlie late thermal overprint transtension low-Ca granites M4 2620-2636 Youanmi-Southern Cross regional metamorphism contraction low-Ca granites M4 [unknown] Youanmi-Ravensthorpe regional metamorphism contraction low-Ca granites M4 2613-2644 Youanmi-Murchison regional metamorphism contraction low-Ca granites M4 [unknown] Southwest-northern regional metamorphism contraction low-Ca granites M4 [unknown] Southwest-NE margin regional metamorphism contraction low-Ca granites M4 2645-2646 Southwest-eastern regional metamorphism contraction low-Ca granites M4 2615-2643 Southwest-Lake Grace regional metamorphism contraction low-Ca granites M4 2641-2646 Southwest-Jimperding regional metamorphism contraction low-Ca granites M4 2625-2645 Southwest-Boddington regional metamorphism contraction low-Ca granites M4 2620-2640 Narryer regional metamorphism contraction none

alteration alteration Ctd>St>Grt±Sill>Crd-An Ctd > Grt > Crd-And Grt > Crd pelite mafic granulite Cpx±Ep amphibolite mafic granulite Sill > Grt+Opx > Crd > P Sill > Grt±Opx > Crd > P [unknown] Sill > Grt+Crd > Crd > P

Late-stage thermal effects during cratonization: poorly constrained, identified by cooling ages and disturbed isotopic systems, parageneses typically not developed Late 2547-2610 Kurnalpi disturbed isotopics none low-Ca granites Late 2593-2598 Kalgoorlie disturbed isotopics none low-Ca granites Late 2600 Southern Cross disturbed isotopics none low-Ca granites Late 2571-2607 Southwest disturbed isotopics none low-Ca granites Late 2410 Kurnalpi Widgemooltha aureole extension mafic dykes

[no parageneses] [no parageneses] [no parageneses] [no parageneses] late Crd growth

Sample

Domain Grt Ttn

Rock type Rt Accessory

P-T path Ilm

Peak T, P G Mag

Qtz Hbl

Pl Gru

Bt Opx

Ms Chl Cpx Ep-Cz

Ma early-formed mafic granulite associated with magmatic arcs - in Burtville and Kurnalpi Terranes: BG6-146c Duketon mafic granulite 715, 3.4, 60.1 X? X BG6-146g Duketon mafic granulite IBC 710, 4.0, 50.7 X? X BG6-147 Duketon mafic granulite 810, 5.2, 44.5 X? X BG6-148 Duketon mafic granulite 761, 4.8, 45.3 X? X BG6-146b Duketon mafic granulite IBC 712, 3.7, 55 X BG6-135a Gindalbie mafic granulite 706, 3.3, 61.1 X? X M1 early-formed high-P parageneses in shear zones - Burtville Terrane: BG6-172b Duketon amphibolite CW>ITD 572, 7.9, 20.7 179693B Merolia amphibolite DC 600, 8.0, 21.4 M1 early-formed high-P parageneses in shear zones - Kurnalpi Terrane: BG6-120k Edjudina amphibolite CW>ITD 607, 7.0, 24.8 Y280 Laverton amphibolite 642, 8.5, 21.6 Y0360 Laverton amphibolite CW>DC 575, 7.3, 22.5 Y0361 Laverton amphibolite LH 580, 6.8, 24.4 X1 X2 X? BG6-171d Linden amphibolite ◊ 593, 8.9, 19 BG6-170a Linden amphibolite CW>ITD 575, 7.5, 21.9 BG6-171b Linden amphibolite DC 625, 8.2, 21.8 BG6-171j Linden amphibolite CW>ITD 600, 6.8, 25.2 BG6-171l Linden ultramafic CW>ITD 615, 7.7, 22.8 BG6-106c Murrin amphibolite CW>ITD 623, 7.2, 24.7 X1 XS M1 early-formed high-P parageneses in shear zone - Kalgoorlie Terrane: BG6-183a Coolgardie meta-pelite CW>ITD 527, 8.5, 17.7 99966141A Boorara amphibolite CW>ITD 580, 8.5, 19.5 BG6-190b Coolgardie amphibolite CW>ITD 580, 8.5, 19.5 BG6-184b Coolgardie amphibolite CW>ITD 611, 8.4, 20.8 BG6-208.1 Depot amphibolite CW>ITD 606, 7.2, 24 CEHC168 Norseman amphibolite CW>ITD 618, 6.8, 26 BG6-178b Ora Banda amphibolite CW>ITD 640, 7.0, 26.1 M1 early-formed high-P parageneses - Southern Cross Domain:

93

X? X?

X X

XS X? X X X X? S XS XS

XS? X

X

R

X

S X S? X? X XC X

XS

XS

Igrt X

Igrt X

X? X XC

X S? XS XS C

X X X

Igrt X Igrt X

XS XS X1 S2 S X S X X? Igrt X

Iopx X Iopx X X

S1 X2 S S R R XS

S

S

S

X

Py X S X XS S

X X X X Ap X

S

S

S? X X X

X X X

X

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BG6-204.1 BG9-163b BG6-206.1 BG10-45

Sample

Forrestania XS Ravensthorpe Cheritons West margin

Domain

amphibolite amphibolite amphibolite amphibolite

Rock type

CW Igrt S CW>ITD CW>ITD

619, 7.1, 24.9 522, 7.9, 18.9 600, 7.6, 22.6 627, 8.0, 22.4

P-T path

Peak T, P G

S S S X? X

Qtz

S

S

R

S X X

S X

S L

Pl

Bt

M2 middle amphibolite facies regional metamorphism with clockwise P-T paths - in Kalgoorlie Terrane and northeast Youanmi Terrane: BG6-207 Boorara meta-pelite CW 537, 5.1, 30.1 S S? 98957 Boorara meta-pelite CW 572, 3.2, 51.1 XS XS C S? BG6-181b Coolgardie meta-pelite CW 556, 5.8, 27.4 Igrt XS S? #78694 Coolgardie meta-pelite CW? 550, 3.8, 41.4 S S? X2 Y667bv Depot meta-pelite CW [544, 6.4, 24.3] Iand X X? Iand S X2 Y528.2i Depot meta-pelite CW>ITD 544, 3.1, 50.1 Iand S S Iand,crd S Y667bii Depot meta-pelite CW 550, 6.2, 25.3 Igrt XS XS S S2 Y666i.2 Depot meta-pelite CW>ITD [545, 5.7, 27.3] Igrt S X X Igrt S X2 #63555-111 Depot meta-pelite CW 569, 3.5, 46.4 X X #63555-104 Depot meta-pelite CW 560, 4.0, 40 X X #NA20 Depot psammo-pelite CW? 540, 3.5, 44.1 X X XC 83001 Depot meta-pelite CW 582, 3.9, 42.6 X X Icrd BG6-213b Moilers meta-pelite CW>ITD 580, 5.8, 28.6 Ist,grt X1 X1 XS1 BG6-213a Moilers meta-pelite CW 540, 6.0, 25.7 X X S1 AXR418e Lake Mason meta-pelite CW? 560, 2.4, 66.7 Igrt X C X XC M4 (±M2) middle amphibolite facies regional metamorphism with clockwise P-T paths - in southwest Youanmi Terrane: BG10-46b Sth Cross meta-pelite CW 577, 3.6, 45.8 Icrd S S1 Icrd S1 Y400a Sth Cross meta-pelite CW>ITD ◊ 588, 4.0, 42.0 S S Y400b,c Sth Cross meta-pelite CW ◊ 535, 4.0, 38.2 S ±S S Y400e Sth Cross meta-pelite CW>ITD 581, 4.0, 41.5 S S BG9-113 Sth Cross calc-psammite CW 550, 4.6, 34.2 X X Igrt X1 BG9-133a Cheritons meta-pelite CW>ITD 602, 3.6, 47.8 S S? S BG9-163f Ravensthorpe meta-pelite CW 562, 3.8, 42.3 S S X1 S BG9-165a Ravensthorpe meta-pelite CW>DC 581, 5.0, 33.2 S S S X3 L4 BG9-172a,b Ravensthorpe meta-pelite CW? 550, 4.2, 37.4 Iand,grt S Iand S Iand S 198197 Mt Magnet meta-pelite CW 555, 3.2, 49.6 S1 S1 X2 GSD073096 Dalgaranga meta-pelite 562, 3.0, 53.5 Icrd X X? X GSD073178 Dalgaranga meta-pelite CW>ITD 577, 2.3, 71.7 Icrd X X S

Sample

Domain

Rock type

P-T path

Peak T, P G

Qtz

Pl

Bt

X X

Ms

Chl

St

Grt

X3 S

X3 S XS S

X1

X2 X2 X

S1 R CS Icrd S X? R R XC

R R R

X Iand X3 X1 X2

X2 X1 X2 Iand X2

X2 X X2 X1

R R

S1 X3 R L X

S1 ±X1 S1 Igrt X1 X1

R

L4 R S C4 S1 X3 C4

X2 Iand X2 X2 X2 X1 X X1 X1

R L4 L5 X3 Iand R

Ms

M3 middle amphibolite parageneses with anti-clockwise P-T paths - in post-volcanic late basin Belches Formation, Kurnalpi Terrane: Y455b Belches Fm. meta-pelite ACW 603, 4.1, 42 Iand X C X XS L Y534a.1 Belches Fm. meta-pelite ITL>ACW 564, 4.1, 39.3 Igrt S S S X2 SLC Y534b Belches Fm. meta-pelite ITL>ACW 570, 4.0, 40.7 Igrt S S S X2 CL Y455f Belches Fm. meta-pelite ACW ◊ 567, 4.0, 40.5 Ist X X X X Y533f Belches Fm. meta-pelite ACW 576, 3.3, 49.9 Ist,and S C S Iand S X2 C C L Y530b Belches Fm. meta-pelite ACW>LC 567, 3.5, 46.3 Ist,crd S C S Icrd S X2 R Y533a Belches Fm. meta-pelite ACW 581, 4, 41.5 Ist,grt S Ist S S X2 C L Y530a Belches Fm. meta-pelite ACW [577, 4.5, 36.6] S S S X2 Y530bii Belches Fm. meta-pelite ACW 577, 4.5, 36.6 S S S X2 L Y455c Belches Fm. meta-pelite ACW ◊ 550, 4.2, 37.4 X X X X 153471 Belches Fm. meta-pelite ACW? ◊ 580, 4.0, 41.4 S S X1 X3 X1 M3 middle amphibolite parageneses with anti-clockwise P-T paths - in extensional Ockerburry shear zone, KalgoorlieKurnalpi Terrane boundary: Y507a,b,e,f Mount Martin aluminous ACW>LH ◊ 500, 4.0, 35.7 Iand S S Iand S F4 S? X3 Iand X2 L4 Iky S1 F4 Gr?

94

Py, Cal S? S X

X1

Chl

St

R LR CR R C SR

X X X2 X3 X2 X3

L R

Grt

X X2 X X2 X

X X3

X3

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BG6-209 Y242a,b

Mount Martin Kyanite Hill

aluminous aluminous

ACW ACW

S S

S C4 S C4

M3 early-formed middle amphibolite parageneses with anti-clockwise P-T paths - in Southern Cross domain, Youanmi Terrane: BG9-27a Diemals Fm. meta-pelite ACW ◊ 550, 2.4, 65.5 S S? S BG9-58c.1 Gum Creek meta-pelite LH 612, 5.6, 31.2 S S? S S BG9-169b Ravensthorpe meta-pelite ACW>IBC 578, 4.6, 35.9 S1 S3 S1? S1 S3 S1 S3 L5 X4 X? X2 BG9-164b Ravensthorpe meta-pelite ACW…M4 CW 584, 3.5, 47.7 S Igrt S S? Igrt S R BG6-197a,b,c Ravensthorpe meta-pelite ACW 575, 3.6, 45.6 Iand XS XS1 X2 Iand XS1 F3 L BG6-198a,b Ravensthorpe meta-pelite ITL>ACW 575, 5.1, 32.2 Igrt,st XS XS XS L XS L BG6-198c Ravensthorpe meta-pelite LH 560, 5.0, 32 XS XS S1 X4 BG6-198d Ravensthorpe meta-pelite LH>DC 570, 4.7, 34.7 Igrt XS XS? S BG6-198f Ravensthorpe meta-pelite ACW 560, 5.2, 30.8 XS XS? XS1 X2 C L6 M3 middle amphibolite parageneses with anti-clockwise P-T paths - in Murchison domain, Youanmi Terrane: BG10-59b,f,h Dalgaranga meta-pelite ACW 514, 1.9, 77.3 S SC Icrd S BG10-59e,k,l Dalgaranga meta-pelite ACW? ◊ 500, 2.0, 71.0 S S Icrd S1 X2 BG10-59q Dalgaranga meta-pelite ACW [500, 2.0, 71.0] S2 S2 190962 Murchison meta-pelite ITL>ACW 619, 3.2, 55.3 XS XS TJIMDG175 Murchison meta-pelite ACW 520, 1.6, 92.9 Igrt XS XS?

Sample

Domain

Rock type

P-T path Peak T, P G

Qtz

Pl

Kfs

M4 late-stage granulite facies regional metamorphism with clockwise P-T paths - in Narryer Terrane: BG10-62 Mt Narryer opx-crd granulite 734, 4.7, 44.6 Iopx X C BG10-63a Mt Narryer metapelite granulite CW 808, 4.8, 48.1 Igrt X BG10-63b Mt Narryer metapelite granulite CW 681, 4.0, 48.6 Igrt X BG10-69i Mt Narryer metapelite granulite CW 770, 5.0, 44.0 XS XC X BG10-69k Mt Narryer metapelite granulite CW • 737, 4.4, 47.9 Igrt X X BG10-69l Mt Narryer metapelite granulite CW • 726, 4.4, 47.2 Icrd,grt X X BG10-70a Mt Narryer metapelite granulite CW • 725, 4.2, 49.3 Igrt X S X1 C2 C2 CR R X XS BG10-70c Mt Narryer metapelite granulite CW • 752, 4.8, 44.8 Igrt X S XS BG10-70g Mt Narryer metapelite granulite CW • 784, 4.6, 48.7 Icrd,grt X S X S R R X1 X2 S2 Icrd BG10-70i Mt Narryer metapelite granulite CW 770, 5.4, 40.7 Igrt X S X S1 C2 CR R X XSC Icrd,grt C M4 late-stage granulite facies regional metamorphism with clockwise P-T paths - in Southwest Terrane: BG10-10b Jimperding opx-grt granulite CW 814, 6.0, 38.8 X X BG10-10d Jimperding opx-grt-crd granulite CW 814, 6.1, 38.1 Igrt,opx X Iopx X X BG10-5a Jimperding metapelite granulite CW 790, 4.4, 51.3 Igrt X XC S R R X X Igrt X Igrt X BG10-5c Jimperding metapelite granulite CW 868, 7.5, 33.1 Igrt X X X BG10-8b Jimperding metapelite granulite CW • 798, 6.0, 38.0 X X BG10-35b Lake Grace metapelite granulite CW 750, 6.9, 31.1 Igrt X X E BG10-35d Lake Grace metapelite granulite CW 758, 6.8, 31.8 X X BG10-39d Lake Grace opx-grt granulite CW 852, 6.4, 38 Igrt X S XS E BG10-39h Lake Grace opx-grt-crd granulite CW 833, 5.2, 45.8 Igrt,crd X X Igrt,crd X S X X1 X2 X1 BG10-39j Lake Grace opx-grt granulite 810, 5.9, 39.2 X X BG10-39k Lake Grace opx-grt granulite CW 809, 5.2, 44.5 Iopx,grt X X S X Iopx X1 BG10-39p Lake Grace opx-grt-crd granulite CW 831, 5.2, 45.7 Icrd X C X1 C2 X1 C2 X X X BG09-158e Eastern psammopelite granulite LH 776, 4.0, 55.4 Igrt X X

Sample

Easting Northing

Domain

Rock Type

P-T Path Interp.

95

Prograde: T, P, G

method

S1 X2 C S X4 C S2 C4 S XS

Bt X Igrt X X Icrd S S1 X2 Icrd X X1 S1 Igrt XS Icrd X

R R

X2

Gr L C L X2 L X4 S XS L5 X3

X2 X2 X3 X2 X4

C CR C X3 X2 X3

XS

Ms

C CR

CR

Chl

Melt

C C R R R

C?

Zrn R Mnz

Igrt X1 Zrn Igrt X Iopx X Igrt X X X1 C2 Igrt X XC X S2

LR LR

R R

R

R

X X

X1 C2Py, Gr X

X Iopx X

X1 X2 Py, Gr Icrd X Igrt X

Peak Metamorphism: T (ºC) P (kb)

R

ºC/km Site §

X

Method ø

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BURTVILLE TERRANE: BG6-146G 461999 6794269 BG6-146B 461999 6794269 BG6-146C 461999 6794269 BG6-148 462580 6787934 BG6-147 462515 6785372 179693B 532291 6918067 SAJMBX13 547378 6901116 BG6-151B 482246 6828172 BG6-152C 482213 6827337 SSRDHS230 583290 6883824 SSRDHS8 563433 6879547

Duketon Duketon Duketon Duketon Duketon Merolia Merolia Merolia Merolia Yamarna Yamarna

mafic granulite mafic granulite mafic granulite mafic granulite mafic granulite Cpx-amphibolite felsic schist calcsilicate Cpx-amphibolite Cpx-amphibolite amphibolite

IBC IBC

710 ± 27 712 ± 25 715 ± 44 761 ± 29 810 ± 31 600 ± 10 • 608 ± 17 650 ± 15 609 ± 16 548 ± 10 585 ± 59

4.0 ± 1.0 3.7 ± 1.5 3.4 ± 1.7 4.8 ± 0.9 5.2 ± 1.3 8.0 ± 0.4 8.6 ± 1.1 6.5 ± 1.0 6.4 ± 1.0 4.8 ± 0.6 5.9 ± 0.5

50.7 55.0 60.1 45.3 44.5 21.4 20.2 28.6 27.2 32.6 28.3

cores rims cores rims rims rims rims edge-rims cores cores cores

KALGOORLIE TERRANE: 98221 305000 6718900 99966118 326493 6817608 BG6-207 337500 6799700 98965829H 349000 6748300 98904 318400 6726700 98957 311000 6707500 99966141A 309649 6824532 Y691f 318279 6722515 MM2/2 375437 6567856 BG6-190B 269098 6735397 BG6-183A 259190 6785065 BG6-189A 268981 6735471 BG6-185B 258245 6785072 BG6-184B 259190 6785065 14 298300 6515000 Y527 317500 6550300 Y667a 346767 6566076 Y667a(2) 346767 6566076 NA24 372137 6500156 Y666e 342875 6568590 Y528(2) 369131 6514812 Y666i(1) 342875 6568590 83001 371437 6470155 BG6-208(1) 359220 6478820 BG6-208(2) 359220 6478820 Y666b 342875 6568590 PR1523 380700 6425200 PR1515 379000 6435700 BG6-213A 297637 6963157 BG6-213B 297637 6963157 Y535d 387522 6431501 SC16 386437 6417155 CEHC137 461835 6393747 CEHC168 447495 6358089 BG6-162 258583 6848571 BG6-159K 259092 6848620 BG6-178B 270317 6807167 120031 258137 6891757 120046 258137 6891757 106726 258137 6891757

Boorara Boorara Boorara Boorara Boorara Boorara Boorara Boorara Boorara Coolgardie Coolgardie Coolgardie Coolgardie Coolgardie Coolgardie Coolgardie Depot Depot Depot Depot Depot Depot Depot Depot Depot Depot Kambalda Kambalda Moilers Moilers Norseman Norseman Norseman Norseman Ora Banda Ora Banda Ora Banda Ora Banda Ora Banda Ora Banda

Cpx-Ep-amphibolite Ep-amphibolite Grt-Ctd-metapelite Ep-amphibolite Grt-Cpx-amphibolite Grt-St-Crd-metapelite Cpx-amphibolite Grt-amphibolite amphibolite Cpx-amphibolite Grt-St-metapelite Grt-metapelite Grt-amphibolite Grt-Gru-amphibolite Grt-Cum-amphibolite Grt-amphibolite Grt-metapelite Grt-metapelite metapelite amphibolite St-Crd-metapelite Grt-And-metapelite St-And-Crd-metapelite Grt-amphibolite Grt-amphibolite Cpx-amphibolite Grt-amphibolite Grt-amphibolite Grt-Ged-Crd-pelite Grt-St-metapelite Grt-Gru-amphibolite amphibolite Cpx-amphibolite Grt-Cpx-amphibolite Cpx-Ep-amphibolite Cpx-Ep-amphibolite amphibolite Grt-amphibolite calcsilicate alteration calcsilicate alteration

553, 3.7, 43 THERMO 580 ± 16 CW > ITD 494, 4.3, 33 THERMO 535 ± 9 CW 487, 6.4, 22 THERMO 537 ± 10 CW > ITD 539 ± 13 569 ± 18 CW 572 ± 20 ◊ 580 ± 25 CW > ITD 625 ± 25 DC 542 ± 11 580 ± 25 CW > ITD 490, 8.2, 17 Gaidies 527 ± 14 548 ± 50 CW > DC 600 ± 25 611 ± 60 607 ± 82 502 ± 29 CW > ITD 544 ± 21 CW > ITD 544 ± 21 CW > ITD 513 ± 19 ◊ 532 ± 25 CW > ITD 544 ± 12 CW > ITD 545 ± 38 CW 550, 4.2, 37 phase 582 ± 14 ◊ 606 ± 60 ◊ 630 ± 25 640 ± 25 ACW >IBC 503 ± 40 LH 527 ± 42 CW 540 ± 20 CW > ITD 480, 5.5, 25 THERMO 580 ± 21 CW > DC 595 ± 26 ◊ 705 ± 25 682 ± 25 CW > ITD 618 ± 43 607 ± 16 607 ± 15 IBC 640 ± 18 511 ± 35 535 ± 35 540 ± 10

8.5 ± 0.6 7.8 ± 0.4 5.1 ± 0.7 7.0 ± 0.8 5.5 ± 0.5 3.2 ± 0.6 8.5 ± 0.8 8.0 ± 0.8 4.2 ± 0.8 8.5 ± 0.8 8.5 ± 1.8 4.1 ± 1.0 6.7 ± 0.8 8.4 ± 2.1 8.3 ± 0.5 3.7 ± 0.6 6.4 ± 0.4 6.4 ± 0.4 2.7 ± 1.1 7.0 ± 0.8 3.1 ± 0.5 5.7 ± 1.0 3.9 ± 0.2 7.2 ± 0.8 7.0 ± 0.8 6.0 ± 1.3 2.9 ± 1.0 3.0 ± 0.7 6.0 ± 0.7 5.8 ± 1.4 4.2 ± 2.8 7.6 ± 0.8 7.5 ± 0.5 6.8 ± 0.9 6.0 ± 1.1 6.2 ± 1.6 7.0 ± 0.8 4.2 ± 0.6 3.1 ± 1.0 3.3 ± 1.0

19.5 19.6 30.1 22.0 29.6 51.1 19.5 22.3 36.9 19.5 17.7 38.2 25.6 20.8 20.9 38.8 24.3 24.3 54.3 21.7 50.1 27.3 42.6 24.0 25.7 30.5 49.6 50.2 25.7 28.6 40.5 26.5 26.0 26.0 28.9 28.0 26.1 34.8 49.3 46.8

cores rims rims cores cores edge-rims cores rim-cores rims cores rims rims cores rims cores cores cores cores cores cores cores cores cores rims

KURNALPI TERRANE: Y534b 442442 6542635 Y534b(2) 442442 6542635 Y455f 439348 6520675 Y459a 416649 6521859 Y533e 442586 6542832 Y458c 443146 6539425 Y534a(1) 442442 6542635 Y534a(2) 442442 6542635 YD091c 418934 6514851 Y530b 441044 6516355 Y531a 443974 6539519 Y533f 442586 6542832 Y530b(ii) 441044 6516355 Y531g 443974 6539519 Y531j 443974 6539519 Y533a 442586 6542832

Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin Belches Basin

Grt-And-Crd-pelite Grt-Crd-And-pelite St-And-metapelite Grt-metapelite St-And-metapelite Grt-metapelite Grt-St-metapelite Grt-St-metapelite Grt-metapelite St-Crd-metapelite Grt-metapelite St-And-Crd-pelite Grt-metapelite Grt-And-metapelite calcsilicate Grt-St-And-metapelite

ITL >ACW ITL >ACW ACW ITL >ACW ITL >ACW ACW >IBC ITL >ACW

4.0 ± 0.7 4.0 ± 0.7 4.0 ± 0.3 4.5 ± 0.7 3.5 ± 0.4 2.8 ± 2.6 4.1 ± 1.0 4.0 ± 1.0 3.6 ± 1.5 3.5 ± 0.3 3.1 ± 0.8 3.3 ± 0.6 4.5 ± 0.9 4.9 ± 0.8 4.3 ± 0.6 4.0 ± 0.17

40.7 40.7 40.5 32.4 45.1 56.6 39.3 40.4 44.8 46.3 52.4 49.9 36.6 33.8 38.5 41.5

rims rims rims rim-cores rims rime-cores rims rims rim-cores rims cores rims rim-cores rims rims rims

DC 2ME DC

550, 4.0, 39 Gaidies

DC CW > ITD

96

ACW > LC ACW >IBC ITL >ACW ACW ITL >ACW ITL >ACW

525, 3.5, 43 Spear 525, 3.5, 43 Spear

570 ± 7 570 ± 7 567 ± 10 511 ± 21 553 ± 7 530, 3.5, 43 Gaidies 555 ± 14 525, 3.5, 43 Spear 564 ± 15 525, 3.5, 43 Spear 565 ± 15 565 ± 60 530, 3.4, 45 Spear 567 ± 6 568 ± 18 576 ± 10 565, 3.4, 48 THERMO 577 ± 12 560, 2.4, 67 phase 580 ± 17 580 ± 33 513, 1.7, 86 THERMO 581 ± 11

THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO

THERMO THERMO THERMO THERMO THERMO THERMO conventio THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO conventio THERMO THERMO THERMO conventio phase sta cores THERMO cores THERMO cores THERMO rim-cores THERMO edge-rims THERMO cores THERMO cores conventio cores THERMO edge-rims THERMO cores THERMO cores THERMO cores THERMO cores THERMO cores THERMO cores THERMO

THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

YD246 Y455b Y531j(ii) Y690a Y690a(2) BG6-120C BG6-120E BG6-120K BG6-122 AB85 99966130L BG6-136B BG6-135A BG6-138A Y0360 Y0361 BG6-143B BG6-174A Y280 BG6-173B BG6-170B BG6-170A BG6-171D BG6-170D BG6-171J BG6-171I BG6-171B BG6-121 BG6-192A BG6-192B BG6-104 BG6-108 BG6-109 BG6-106C ER1363

439929 439348 443974 473916 473916 473953 473953 473953 464698 339337 339115 333673 348120 331271 429226 429226 426100 359215 391455 364575 395070 395070 395065 395070 395065 395065 395065 474380 436381 436381 419443 420951 420493 419979 485620

6514305 6520675 6539519 6682101 6682101 6682134 6682134 6682134 6659177 6798357 6799058 6736993 6753988 6738551 6833819 6833819 6826305 6937917 6839778 6941081 6957641 6957641 6958405 6957641 6958405 6958405 6958405 6682118 6625767 6625767 6755759 6743234 6743256 6756980 6539950

Belches Basin Belches Basin Bulong Edjudina Edjudina Edjudina Edjudina Edjudina Edjudina Gindalbie Gindalbie Gindalbie Gindalbie Gindalbie Laverton Laverton Laverton Laverton Laverton Laverton Linden Linden Linden Linden Linden Linden Linden Linden Menangina Menangina Murrin Murrin Murrin Murrin Murrin

Grt-St-And-metapelite St-And-metapelite amphibolite Grt-mafic mylonite Grt-mafic mylonite Grt-amphibolite Grt-amphibolite Grt-amphibolite amphibolite amphibolite calcpelite schist amphibolite mafic granulite Cpx-amphibolite Grt-amphibolite Grt-amphibolite amphibolite Cpx-amphibolite amphibolite amphibolite Grt-amphibolite Grt-Gru-amphibolite Cpx-Gru-amphibolite Grt-Gru-amphibolite Grt-Cum-amphibolite Grt-Gru-amphibolite Grt-Cpx-amphibolite amphibolite Grt-amphibolite Grt-amphibolite Ep-amphibolite amphibolite Ep-amphibolite Grt-Ep-amphibolite Grt-metapelite

2ME ACW DC

588 ± 41 603 ± 12 ◊ 609 ± 25 510 ± 15 510 ± 15 ◊ 570 ± 25 578 ± 55 607 ± 10 482 ± 83 533 ± 18 545 ± 11 626 ± 25 706 ± 25 686 ± 29 575 ± 30 580 ± 36 610 ± 18 ◊ 630 ± 25 642 ± 44 ◊ 670 ± 25 531 ± 13 575 ± 25 ◊ 593 ± 25 ◊ 600 ± 25 600 ± 72 615 ± 53 625 ± 41 580 ± 15 550 ± 49 594 ± 20 520 ± 32 545 ± 10 ◊ 580 ± 25 623 ± 61 570 ± 14

4.4 ± 1.3 4.1 ± 0.4 7.1 ± 0.75 6.6 ± 0.5 6.6 ± 0.5 8.3 ± 0.8 8.3 ± 0.8 7.0 ± 1.0 4.4 ± 1.2 6.9 ± 0.9 5.5 ± 1.4 4.5 ± 0.8 3.3 ± 1.1 4.0 ± 0.8 7.3 ± 2.8 6.8 ± 0.8 6.0 ± 0.8 7.0 ± 0.8 8.5 ± 1.4 7.5 ± 0.8 7.9 ± 1.1 7.5 ± 0.8 8.9 ± 0.8 7.4 ± 0.8 6.8 ± 2.1 7.7 ± 0.8 8.2 ± 1.2 5.5 ± 0.7 8.0 ± 2.0 6.8 ± 2.8 7.0 ± 0.8 4.7 ± 0.4 8.5 ± 0.8 7.2 ± 1.6 5.1 ± 0.8

38.2 42.0 24.5 22.1 22.1 19.6 19.9 24.8 31.3 22.1 28.3 39.7 61.1 49.0 22.5 24.4 29.0 25.7 21.6 25.5 19.2 21.9 19.0 23.2 25.2 22.8 21.8 30.1 19.6 25.0 21.2 33.1 19.5 24.7 31.9

rims cores cores rim-cores rim-cores cores rim-cores cores cores cores cores cores cores rim-cores rim-cores cores cores cores cores cores cores cores cores rim-cores cores rims cores cores cores cores cores cores rim-cores rim-cores

THERMO THERMO conventio THERMO THERMO phase sta THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO conventio THERMO conventio THERMO THERMO conventio conventio THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO phase sta THERMO THERMO

491 ± 8 520 ± 9 575 ± 20 490 ± 36 562 ± 85 565 ± 25 560, 4.6, 35 Gaidies 577 ± 72 600 ± 28 555 ± 9 470 ± 50 520 ± 7 566 ± 50 598 ± 74 603, 2.8, 62 THERMO 619 ± 57 635 ± 104 575 ± 35 560 ± 50 625 ± 25 709 ± 25 556 ± 25 568 ± 25

4.3 ± 0.5 4.9 ± 0.4 6.4 ± 1.0 2.2 ± 1.1 3.0 ± 2.5 5.4 ± 0.7 2.3 ± 0.9 3.2 ± 0.6 3.2 ± 0.5 1.5 ± 0.5 1.6 ± 0.4 3.0 ± 0.5 4.4 ± 1.3 3.2 ± 0.7 5.1 ± 1.1 4.8 ± 0.5 4.6 ± 0.5 5.1 ± 1.0 0.24 ± 0.3 4.0 ± 1.0 3.2 ± 0.7

32.6 30.3 25.7 63.6 53.5 29.9 71.7 53.6 49.6 89.5 92.9 53.9 38.8 55.3 35.6 34.2 34.8 35.0 844.0 39.7 50.7

cores cores cores rims cores cores cores cores cores cores cores cores rims rim-cores rims cores cores cores cores cores cores

THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO

500, 4.2, 34 THERMO 552 ± 70 580 ± 30 600 ± 14 542 ± 20 530, 5.6, 27 Gaidies 556 ± 75 510, 5.0, 29 THERMO 557 ± 25 540, 4.9, 32 Gaidies 579 ± 50 600 ± 35 607 ± 50 610 ± 29 611 ± 30 612 ± 35

4.3 ± 0.9 4.2 ± 0.7 7.6 ± 0.4 4.4 ± 0.5 4.6 ± 1.3 4.3 ± 0.7 4.3 ± 1.3 5.1 ± 1.0 6.3 ± 1.0 6.2 ± 1.0 6.3 ± 1.4 6.2 ± 0.7

36.7 39.5 22.6 35.2 34.5 37.0 38.5 33.6 27.5 28.1 27.7 28.2

rim-cores cores rim-cores cores cores rims rims cores cores cores cores cores

THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO

CW > ITD CW CW > ITD DC DC

CW > DC LH CW > ITD

CW > ITD CW > ITD CW > ITD CW > ITD DC CW > ITD CW > ITD DC IBC DC CW > ITD 543, 5.4, 29 Gaidies

YOUANMI TERRANE - Murchison Domain: 190546 565921 6979856 Cue 190542 562748 6975148 Cue 194524 629811 6993968 Cue GSD073091 521300 6934300 Dalgaranga GSD073096 521700 6934900 Dalgaranga GSD073054 500600 6956900 Dalgaranga GSD073178 521000 6934800 Dalgaranga GSD083343 563300 7077400 Koonmarra 198197 590796 6925066 Mt Magnet 190964 589278 7020064 Murchison TJIMDG175 589334 7020233 Murchison AHH073631 610800 6997900 Murchison 199678 611150 6908165 Murchison 190962 589341 7020101 Murchison 199687 626826 6956472 Murchison BG10-60 450304 6989946 Murgoo AHH087657 418100 6911300 Tardie GSD081674 640900 7053000 Tieraco 193916X 632511 6805722 Windimurra GSD083357 458700 6854400 Yalgoo GSD083366 467800 6826100 Yalgoo

Grt-amphibolite Ep-amphibolite Cpx-amphibolite Cpx-Ep-amphibolite Crd-metapelite Cpx-amphibolite Grt-Crd-metapelite Cpx-amphibolite St-And-Crd-metapelite Grt-And-metapelite Grt-And-Ctd-pelite Ep-amphibolite Grt-amphibolite Grt-Crd-aluminous rock Grt-amphibolite amphibolite calcsilicate Cpx-amphibolite Crd-Opx hornfels Cpx-amphibolite Cpx-amphibolite

YOUANMI TERRANE - Southern Cross Domain: BG09-95g 696191 6587623 Bullfinch BG09-96b 683090 6621664 Bullfinch BG6-206(1) 744381 6507218 Cheritons NUG12-57 746000 6500000 Cheritons LXB106781 742100 6499600 Cheritons LXB106778 742100 6499600 Cheritons LXB106782 742100 6494300 Cheritons NUG12-53 746000 6500000 Cheritons 5 755000 6518200 Cheritons 1B 756500 6527800 Cheritons LXB106758 744300 6507100 Cheritons 4C 756500 6527800 Cheritons

Grt-Cum-amphibolite CW amphibolite Grt-amphibolite Grt-Bt-Cum Fe-pelite IBH Grt-metapelite CW Grt-And-metapelite CW Grt-St-metapelite CW Grt-Hbl skarn Grt-Hbl skarn ultramafic rock Grt-Cpx-amphibolite CW > ITD Opx-Cpx-Gru-ultramafic

97

510, 6.0, 24 Spear

CW > ITD CW > ITD DC CW ITL >ACW ITL >ACW DH DC ACW >IBC DC DC

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BG09-121c BG09-128 BG09-133a BG09-27a BG09-37b BG09-89a BG09-84f BG09-87b BG09-87a BG09-80a 121540 121542 BG09-23 5202 5405 DR5 5398/1 BG6-204(2) 18912 BG6-205 18994 BG6-204(1) WR120 DR6 5220 BG09-135 SXWYA460 BG09-58c1 AXR1037 AXR411b AXR502 BG6-196A 105941 BG6-198C BG6-198F BG6-198D BG6-198A 113915 118839 BG09-164c BG09-169b BG09-165a BG09-167 BG09-164b BG09-175a BG09-174 BG09-173a BG09-163e BG09-163b BG09-172b BG09-163f BG09-163a 38037 121426 121600 121595 121598 121427 Y400a,e 121597 121428 CD17 CD10 BG10-46c BG10-46d FR06 BG10-46a PS01 PS05 HH37 HH17

735480 743954 746790 705479 709791 764840 763864 764223 764223 760733 740500 740500 749008 763000 763000 763000 763000 759000 763000 759000 763000 759000 763000 763000 763000 749936 699767 755841 764223 766001 764840 775130 234500 234475 234475 234475 234475 225100 765900 230114 232677 230521 231971 230114 764453 763591 769898 224959 224959 233100 224959 224959 721591 682450 721591 721591 721591 717750 721591 721591 717750 712000 712000 736056 736056 722000 736056 723500 723500 718000 718000

6515686 6491536 6473487 6693740 6759972 6877709 6878479 6878298 6878298 6892602 6706000 6706000 6724862 6376000 6376000 6376000 6376000 6395000 6376000 6395000 6376000 6395000 6376000 6376000 6376000 6439333 7014851 6967910 6878298 6871243 6877709 6280069 6281000 6280769 6280769 6280769 6280769 6280100 6273900 6285587 6285010 6285559 6284852 6285587 6261651 6265403 6278666 6286075 6286075 6279700 6286075 6286075 6536781 6621750 6536781 6536781 6536781 6546050 6536781 6536781 6546050 6553000 6553000 6515395 6515395 6540000 6515395 6539200 6539200 6547000 6547000

Cheritons Cheritons Cheritons Diemal Basin Diemal Edale SZ Edale SZ Edale SZ Edale SZ Edale SZ Evanston Evanston Evanston Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Forrestania Gum Creek Gum Creek Lake Mason Lake Mason Lake Mason Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Ravensthorpe Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross

Grt-amphibolite Cpx-amphibolite St-Grt-And-metapelite Crd-And-metapelite Cpx-amphibolite Grt-Gru-amphibolite amphibolite Grt-Cum-amphibolite Grt-Cum-amphibolite Grt-Ged-Crd mylonite Ep-amphibolite Cpx-amphibolite Grt-metapelite amphibolite amphibolite amphibolite amphibolite Grt-Bt schist amphibolite Grt-amphibolite amphibolite Grt-amphibolite amphibolite amphibolite amphibolite amphibolite Cpx-amphibolite Grt-Crd-metapelite Grt-Crd-And-pelite Grt-amphibolite Grt-Cum-amphibolite Grt-metapelite Grt-St-And-metapelite Grt-St-Ctd-metapelite Grt-St-And-metapelite Grt-Ctd-metapelite Grt-St-And-metapelite Grt-metapelite Grt-Crd-metapelite Grt-amphibolite Grt-And-metapelite Grt-Crd-And-metapelite Grt-metapelite Grt-Crd-metapelite Grt-Gru-amphibolite Grt-Cum-amphibolite Grt-Cum-amphibolite Grt-metapelite Grt-amphibolite Grt-metapelite Grt-Crd-metapelite Grt-And-metapelite Grt-And-Crd-metapelite Grt-Crd-Anth-schist Grt-And-Crd-metapelite Bt-Crd-mineralization Cpx-amphibolite Grt-Crd-metapelite Grt-St-And-metapelite Grt-Crd-metapelite Grt-And-Crd-metapelite Grt-Hbl-BIF felsic schist Grt-Cum-amphibolite Grt-amphibolite amphibolite Grt-Cum-Bt Fe-pelite Cpx-amphibolite Grt-amphibolite felsic schist felsic schist

98

DH

606 ± 38 613 ± 50 CW > ITD 564, 3.6, 45 Gaidies 602 ± 4 ACW ◊ 550 615 ± 50 CW 476, 3.4, 40 THERMO 502 ± 46 527 ± 20 565 ± 46 CW > ITD 604 ± 48 CW 545, 3.6, 43 Gaidies 612 ± 72 DC 557 ± 15 DC 626 ± 14 CW > ITD 574, 5.4, 30 Gaidies 568 ± 7 454 ± 43 552 ± 64 578 ± 23 600 ± 19 DC ◊ 600 ± 25 601 ± 33 CW > ITD 524, 3.1, 48 THERMO 610 ± 25 612 ± 66 CW 619 ± 17 626 ± 14 DC 660 ± 50 676 ± 17 575 ± 20 629 ± 51 LH 530, 3.6, 42 Gaidies 612 ± 14 CW 537 ± 8 CW > ITD 576 ± 64 2ME 602 ± 50 544 ± 94 CW > ITD 555, 4.3, 37 Spear 548 ± 16 LH 550, 4.3, 37 Gaidies 560 ± 13 ITL >ACW 540, 2.3, 67 Gaidies 560 ± 11 LH > DC 530, 3.5, 43 Gaidies 570 ± 10 ITL >ACW 539, 2.3, 61 Gaidies 575 ± 11 ACW > LC 590, 2.7, 62 Spear 620 ± 108 CW > ITD 583, 6.8, 25 Spear 620 ± 86 LH 525 ± 25 ACW >IBC 528, 3.2, 47 Gaidies 578 ± 117 CW > DC 523, 3.0, 50 Gaidies 581 ± 11 CW 530, 3.2, 47 Gaidies 582 ± 50 ACW…CW 492, 3.0, 47 Gaidies 584 ± 12 558 ± 55 559 ± 23 584 ± 20 CW > ITD 450, 3.2, 40 Gaidies 528 ± 8 CW > ITD 522 ± 7 538, 4.0, 38 Gaidies 550 ± 50 CW 482, 3.1, 44 Gaidies 562 ± 16 CW > IBC 540, 4.2, 37 Gaidies 595 ± 21 CW > ITD 505, 4.9, 29 Spear 511 ± 26 520 ± 46 CW > ITD 516, 4.7, 31 Spear 539 ± 8 555 ± 9 561 ± 25 ACW? 538, 2.6, 59 Gaidies 580 ± 25 CW > ITD 520, 5.0, 30 Gaidies 581 ± 12 CW 510, 4.6, 32 Spear 592 ± 9 593 ± 88 480 ± 35 LH 480, 5.0, 27 Gaidies 495 ± 50 CW 501 ± 67 505 ± 70 DC 510 ± 35 DC 519 ± 70 DC 525 ± 25 DC 540 ± 25 573 ± 17 575 ± 50

5.0 ± 0.5 3.1 ± 0.4 3.6 ± 0.1 2.4 3.6 ± 0.7 3.6 ± 0.5 4.3 ± 0.4 2.6 ± 0.7 3.2 ± 0.6 5.9 ± 0.9 6.9 ± 0.6 7.1 ± 0.8 5.2 ± 0.7 5.4 ± 1.0 4.9 ± 1.7 6.6 ± 2.0 8.3 ± 2.0 7.0 ± 0.8 4.9 ± 1.1 6.5 ± 0.8 4.4 ± 1.4 7.1 ± 1.1 6.7 ± 1.0 7.0 ± 2.0 6.3 ± 2.1 3.2 ± 0.5 7.7 ± 1.6 5.6 ± 1.4 2.0 ± 0.3 3.3 ± 1.0 4.9 ± 1.1 5 ± 1.6 2.2 ± 0.6 5.0 ± 1.0 5.2 ± 0.6 4.7 ± 1.0 5.1 ± 0.5 3.9 ± 0.9 5.9 ± 1.4 7.1 ± 0.5 4.6 ± 1.1 5.0 ± 0.6 4.3 ± 1.0 3.5 ± 1.1 6.0 ± 1.7 3.2 ± 1.0 3.2 ± 1.2 3.6 ± 0.5 7.9 ± 0.4 4.2 ± 0.8 3.8 ± 1.4 4.6 ± 0.6 2.6 ± 0.5 4.1 ± 1.7 2.5 ± 0.5 3.9 ± 0.5 7.0 ± 1.2 4.9 ± 0.8 4.0 ± 1.5 3.5 ± 0.4 3.3 ± 0.5 2.8 ± 0.7 3.7 ± 1.0 3.1 ± 1.3 3.6 ± 0.8 2.9 ± 0.6 3.6 ± 0.9 3.1 ± 0.5 6.2 ± 1.1 3.4 ± 0.7 4.4 ± 1.0

34.6 56.5 47.8 65.5 48.8 39.8 35.0 62.1 53.9 29.6 23.1 25.2 31.2 24.0 32.2 25.0 20.7 24.5 35.0 26.8 39.7 24.9 26.7 26.9 30.7 51.3 23.3 31.2 76.7 49.9 35.1 31.1 71.2 32.0 30.8 34.7 32.2 45.4 30.0 21.1 35.9 33.2 38.7 47.7 26.6 49.9 52.1 41.9 18.9 37.4 42.3 37.0 56.2 36.2 61.6 40.7 22.9 33.8 41.5 48.3 51.3 49.0 38.2 46.2 40.1 50.2 41.2 48.4 24.9 48.2 37.3

rims cores rim-cores rims rims cores rim-cores rims rim-cores cores cores rim-cores cores cores cores cores cores edge-rims cores cores cores cores cores cores cores rim-cores rim-cores cores rims cores rim-cores edge-rims rim-cores rim-cores edge-rims cores rims rim-cores cores rims cores rim-cores rim-cores cores edge-rims rims cores cores rim-cores rim-cores rims rim-cores rims rims rims rim-cores edge-rims rims cores cores rims cores rim-cores cores rim-cores cores cores cores cores

THERMO THERMO THERMO phase sta THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO phase sta THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BG10-46b PS04 FR16 FR01 HD12 BG09-113 BG09-114c BG09-111b BG09-123 BG09-68 BG09-69b

736056 723500 722000 722000 718000 720539 722322 716687 733393 202200 228343

6515395 6539200 6540000 6540000 6547000 6542924 6541393 6548263 6521541 7033300 7032090

Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Southern Cross Wiluna Wiluna

And-Sill-Crd-metapelite Cpx-amphibolite Grt-amphibolite Ol-Opx-Cum-ultramafic Cpx-amphibolite Grt-St-Sill-metapelite Cpx-Ep-amphibolite Grt-Cum-Crd rock Cum-amphibolite Cpx-amphibolite felsic schist

CW DC CW

SOUTHWEST TERRANE: BG10-11d 494100 6416800 BG10-11b 494100 6416800 BG10-5a 474900 6485500 BG10-8b 517928 6436376 BG10-10a 538023 6442092 BG10-11c 494100 6416800 BG10-10b 538023 6442092 BG10-10d 538023 6442092 BG10-5b 474900 6485500 BG10-9a 528500 6449900 BG10-5c 474900 6485500 BG10-2a 468363 6506176 BG10-4c 469671 6506828 BG10-28a 389800 6200700 BG10-28b 389800 6200700 BG10-27b 393940 6205050 BG10-28c 389800 6200700 BG10-17 395108 6321286 BG10-15a 404765 6324682 BG09-145a 723495 6384109 BG09-147c 711769 6371083 BG09-158a 689560 6236925 BG09-148 677045 6396500 BG09-147b 711769 6371083 BG09-149 675708 6393163 BG09-160 691073 6241925 BG09-158e 689560 6236925 BG09-151b 664470 6399479 BG09-157a 690844 6236286 BG09-147a 711769 6371083 BG09-153 657027 6345359 BG09-158b 689560 6236925 BG09-146a 711183 6371091 BG09-120 712535 6487051 BG09-119 711967 6495046 BG10-52 482924 6539251 BG10-35b 569703 6316611 BG10-35d 569703 6316611 BG10-37 621392 6323585 BG10-38 623873 6338445 BG10-33i 584193 6287982 BG10-39k 623327 6340436 BG10-39j 623327 6340436 BG10-34 581605 6298400 BG10-39p 623327 6340436 BG10-39h 623327 6340436 BG10-39d 623327 6340436 BG10-1a 471629 6503730 BG10-41 601726 6454993 BG10-45 672861 6529253 BG10-44 654857 6538602 BG10-56 467915 6688540 BG10-54 494300 6648900

Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Jimperding Balingup Balingup Balingup Balingup Balingup Balingup Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern Eastern East Transition East Transition Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace Lake Grace North Transition North Transition North Transition North Transition Northern Northern

mafic granulite mafic granulite Grt-Crd-pelite granulite Grt-Crd-pelite granulite mafic granulite Ol-Opx-Hbl ultramafic Grt-Opx-pelite granulite Grt-Crd-Opx granulite mafic granulite mafic granulite Grt-felsic granulite Grt-Opx-granulite Grt-pelite granulite Grt-mafic mylonite Grt-mafic mylonite Grt-mafic mylonite Grt-mafic mylonite Grt-amphibolite mafic granulite mafic granulite mafic granulite Grt-Bt migmatite mafic granulite mafic granulite mafic granulite mafic granulite Grt-Crd-felsic granulite mafic granulite mafic granulite Grt-Opx-BIF granulite mafic granulite Opx-felsic granulite mafic granulite mafic granulite Cpx-amphibolite amphibolite Grt-pelite migmatite Grt-Sill-pelite granulite mafic granulite mafic granulite mafic granulite Grt-Opx-pelite granulite Grt-Opx-pelite granulite mafic granulite Grt-Crd-Opx granulite Grt-Crd-Opx granulite Grt-Opx-pelite granulite Cpx-amphibolite Cpx-amphibolite Cpx-ep-amphibolite Ep-amphibolite mafic granulite mafic granulite

IBC

NARRYER TERRANE: C503101 505270 C503108 494546 BG10-61c 395164 BG10-69l 440465

Jack Hills Jack Hills Narryer Hills Narryer Hills

Grt-amphibolite Grt-Ged schist amphibolite Grt-Crd-pelite granulite

7105938 7105157 7029983 7068361

99

CW ACW?

CW IBC CW CW DC DC CW CW CW 2ME 2ME 2ME 2ME DC

577 ± 25 605 ± 25 617 ± 10 620 ± 35 665 ± 55 550 555 ± 23 539, 2.1, 73 THERMO 569 ± 14 573 ± 20 517 ± 35 600 ± 50

715, 7.2, 28 Spear 600, 6.0, 29 Spear

700, 8.4, 24 Spear

LH IBC LH IBC

CW 680, 7.5, 26 Spear CW 695, 8.2, 24 Spear ITL ITL CW > ITD CW IBC CW CW CW IBC DC IBC

CW CW IBC CW

495, 4.8, 30 Gaidies 620, 6.6, 27 Gaidies

3.6 ± 0.4 3.2 ± 0.6 3.4 ± 0.5 2.8 ± 1.0 5.9 ± 1.0 4.6 4.5 ± 0.5 3.3 ± 0.5 5.3 ± 0.4 3.8 ± 0.7 3.8 ± 1.0

45.8 54.0 51.8 63.3 32.2 34.2 35.2 49.3 30.9 38.9 45.1

cores cores rim-cores cores cores cores rims rim-cores rim-cores cores cores

THERMO THERMO THERMO THERMO THERMO Grt isople THERMO THERMO THERMO THERMO THERMO

• 750 ± 61 760 ± 20 790 ± 84 • 798 ± 71 803 ± 30 807 ± 25 814 ± 76 814 ± 29 820 ± 61 850 ± 38 868 ± 31 • 772 ± 70 • 850 ± 25 587 ± 9 597 ± 38 680 ± 20 705 ± 200 706 ± 56 775 ± 34 712 ± 25 749 ± 17 751 ± 33 768 ± 26 770 ± 12 771 ± 27 772 ± 25 776 ± 59 778 ± 25 780 ± 25 800 ± 25 800 ± 73 801 ± 50 815 ± 25 • 735 ± 35 645 ± 50 645 ± 20 750 ± 27 758 ± 34 771 ± 51 • 770 ± 39 798 ± 35 809 ± 112 810 ± 50 831 ± 69 831 ± 46 833 ± 42 852 ± 50 684 ± 56 615 ± 20 627 ± 27 680 ± 20 795 ± 31 800 ± 13

6.0 ± 1.5 4.0 ± 0.5 4.4 ± 1.2 6±1 5.4 ± 2.2 5.0 ± 1.0 6 ± 1.1 6.1 ± 0.6 5.5 ± 2 8.1 ± 1.5 7.5 ± 0.9 6.1 ± 1.4 7.4 ± 1.5 5.5 ± 0.5 6.6 ± 0.8 6.5 ± 0.5 6.8 ± 1.3 6.3 ± 1.2 7.9 ± 0.6 4.2 ± 0.5 5.3 ± 1.1 7.1 ± 1.1 4.9 ± 0.9 4.4 ± 0.5 4.6 ± 1.1 4.3 ± 0.5 4.0 ± 0.9 4.4 ± 0.5 4.3 ± 3.0 5.3 ± 1.0 4.2 ± 3.3 5.1 ± 1.0 5.6 ± 0.5 5.8 ± 0.5 5.2 ± 1.0 4.9 ± 0.4 6.9 ± 1.1 6.8 ± 1.1 6.7 ± 1.8 6.4 ± 1.8 8.5 ± 1.3 5.2 ± 1.5 5.9 ± 1.7 4.8 ± 3.5 5.2 ± 0.6 5.2 ± 0.5 6.4 ± 1.2 6.4 ± 0.4 4.4 ± 0.4 8.0 ± 0.9 6.8 ± 0.3 2.2 ± 1.3 4.7 ± 1.9

35.7 54.3 51.3 38.0 42.5 46.1 38.8 38.1 42.6 30.0 33.1 36.2 32.8 30.5 25.8 29.9 29.6 32.0 28.0 48.4 40.4 30.2 44.8 50.0 47.9 51.3 55.4 50.5 51.8 43.1 54.4 44.9 41.6 36.2 35.4 37.6 31.1 31.8 32.9 34.4 26.8 44.5 39.2 49.5 45.7 45.8 38.0 30.5 39.9 22.4 26.8 103.2 48.6

cores cores cores cores cores cores cores cores cores cores cores cores rim-cores cores rim-cores cores cores cores cores cores rims rim-cores rims rim-cores cores cores cores rim-cores rims cores cores cores rims cores cores cores cores cores cores rims cores cores cores cores cores cores rim-cores cores cores cores cores cores cores

THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO THERMO

533 ± 45 550 ± 20 610 ± 15 • 726 ± 25

2.5 ± 0.6 4.3 ± 0.7 4.0 ± 0.3 4.4 ± 0.8

60.9 36.5 43.6 47.2

cores cores cores cores

THERMO THERMO THERMO THERMO

Yilgarn Metamorphism, PCR Review: Goscombe, Foster, Blewett et al., 2019

BG10-70g BG10-70a BG10-63b BG10-65 BG10-69k BG10-70c BG10-62 BG10-71a BG10-69i BG10-70i BG10-71b BG10-63a

440332 440332 438441 437237 440465 440332 420139 426086 440465 440332 426086 438441

7068444 7068444 7066476 7065807 7068361 7068444 7034565 7040689 7068361 7068444 7040689 7066476

Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills Narryer Hills

Grt-Crd-pelite granulite Grt-Crd-pelite granulite Grt-Crd-pelite granulite meta-gabbro Grt-Crd-pelite granulite Grt-Crd-pelite granulite Crd-Opx-Cum-granulite granite Grt-Crd-pelite granulite Grt-Crd-pelite granulite mafic granulite Grt-Crd-pelite granulite

CW CW CW

615, 6.1, 29 Gaidies 620, 7.2, 25 Gaidies 560, 4.1, 39 Spear

CW CW

625, 7.3, 25 Gaidies 618, 6.3, 28 Gaidies

DC CW CW LH CW

530, 6.6, 23 Gaidies 615, 6.1, 29 Gaidies 562, 4.0, 40 Spear

• 784 ± 50 • 725 ± 50 681 ± 76 • 805 ± 50 • 737 ± 25 • 752 ± 50 734 ± 76 ◊ 760 ± 35 770 ± 34 770 ± 68 800 ± 45 808 ± 130

4.6 ± 0.8 4.2 ± 1.2 4.0 ± 1.3 3.0 ± 0.5 4.4 ± 1.2 4.8 ± 0.8 4.7 ± 1.5 4.8 ± 0.5 5.0 ± 0.5 5.4 ± 1.0 6.0 ± 2.1 4.8 ± 1.3

Highlights: • Middle Neoarchaean orogenic cycle involved four distinct metamorphic periods with different tectonic settings. • Three M1 high-P/moderate-T hairpin clockwise P-T path events at 2748±39, 2727±12 and 2706±20 Ma, during deep burial of magmatic arc margins during accretion events. • M2 low-P/moderate-T clockwise P-T path regional-contact metamorphism at 2671±6 Ma in backarc settings during subduction related granite bloom. • M3 low-P/moderate-T anticlockwise P-T path metamorphism at 2656±5 Ma in post-volcanic rift basins formed during lithospheric extension. • M4 low-P/high-T clockwise P-T path regional-scale thermal pulses at 2644±4 and 2629±7 Ma resulting from lower-crust and mantle lithosphere delamination at 2665-2668 Ma.

100

48.7 49.3 48.6 76.7 47.9 44.8 44.6 45.2 44.0 40.7 38.1 48.1

cores cores cores cores cores cores cores cores cores cores rims cores

THERMO THERMO THERMO THERMO THERMO THERMO THERMO phase sta THERMO THERMO THERMO THERMO