Gold Systems in the Faina Greenstone Belt and Crustal Evolution of The

Gold systems in the Faina greenstone belt and crustal

evolution of the southern Goiás Archean Block, central Brazil

Jessica Bogossian M.Sc., B.Sc.

This thesis is presented for the degree of Doctor of Philosophy of the University of Western Australia

School of Earth Sciences

Centre for Exploration Targeting

2019

Supervisors: Prof. Steffen G. Hagemann

Dr. Anthony Kemp

Lydia M. Lobato THESIS DECLARATION

I, Jessica Bogossian, certify that:

This thesis has been substantially accomplished during enrolment in the degree.

This thesis does not contain material which has been submitted for the award of any other degree or diploma in my name, in any university or other tertiary institution.

No part of this work will, in the future, be used in a submission in my name, for any other degree or diploma in any university or other tertiary institution without the prior approval of The University of Western Australia and where applicable, any partner institution responsible for the joint-award of this degree.

This thesis does not contain any material previously published or written by another person, except where due reference has been made in the text and, where relevant, in the Declaration that follows.

The work(s) are not in any way a violation or infringement of any copyright, trademark, patent, or other rights whatsoever of any person.

The work described in this thesis was funded by the “Conselho Nacional de Desenvolvimento Científico e Tecnológico” - CNPq (Project No. 207220/ 2014-0). Jessica Bogossian was supported by CNPq scholarship during the PhD candidature.

Technical assistance was kindly provided by Dr Malcom Roberts for assistance with the JEOL JXA-8530F ion probe at the Centre for Microscopy, Characterisation and Analyses at UWA (Chapter 2); Hao Gao, Cristina Talavera and Allen Kennedy for U-Pb isotope analyses on the SHRIMP II at the John De Laeter Centre, Curtin University (Chapter 3); Dr. Heejin Jeon and Dr Laure Martin for O isotope analyses in zircons using the Cameca 1280 at the Centre for Microscopy, Characterisation and Analyses at UWA (Chapter 3); Dr. Anthony Kemp for Hf isotope determination with the LA-MC-ICP- MS at the School of Earth Sciences, UWA; Prof Leonid V. Danyushevsky and Paul Olin for arsenopyrite LA-ICP-MS analyses at the Centre for Ore Deposit and Earth Sciences (CODES), in the University of Tasmania.

This thesis contains published work and/or work prepared for publication, some of which has been co-authored.

Signature:

Date: May 6th, 2019

ABSTRACT

Mineral systems associated with Archean cratons are typified by large gold, nickel, iron and base metal deposits. The crustal evolution of cratonic areas and its margins has an essential role controlling the formation and distribution of mineral deposits. In the southern Goiás Archean Block (GAB), gold systems associated with the Faina greenstone belt have been historically and currently explored, yet they lack detailed studies defining their structural controls, hydrothermal alteration characteristics and genesis. In a broader context, still unclear complex and polyphase evolution of the southern Goiás Archean Block makes this an attractive area to investigate long-lived regional magmatism. Isotopic and elemental variations acquired in this study were combined with available data to help unravel the petrogenesis of the GAB. Accordingly, empirical patterns based on the spatial distribution of mineral deposits are used to propose links between crustal growth and gold endowment. Therefore, a multi- disciplinary analysis was conducted at different scales in order to design and evaluate the history of this terrane. Located in the mid-west part of the Goiás state, the NW-trending Faina greenstone belt hosts the Cascavel and Sertão deposits, ~ 28 km away from each other. Historical records indicate approximately 3 Moz Au were produced from alluvial sand and gravels in the region over the past 200 years. Recent mining in the Sertão deposit produced 254 koz Au of oxidized sulfide-rich ore with an average grade of 24.9 g/t Au, and an underground mine is currently under development at the Cascavel deposit. The lithostratigraphy, regional metamorphism, structural setting, hydrothermal alteration mineralogy and zonation of the Cascavel and Sertão deposits are presented in Chapter 2. This investigation included optical petrography of drill core samples, scanning electron microscope (SEM) and electron ion probe microanalyses (EPMA) to characterize the nature and composition of hydrothermal minerals associated with gold mineralization. The Faina greenstone belt is composed of a lower sequence of mafic-ultramafic rocks unconformably overlain by a sedimentary sequence. Regional shortening during Paleoproterozoic deformation caused the development of early structures, e.g., isoclinal F1 folds, S1 foliation (S0 transposed into S1), barren S1-parallel parasitic veins, the inversion of the Faina volcano-sedimentary paleobasin into a synform, and greenschist facies metamorphism expressed by quartz-chlorite- biotite±alkali-feldspar assemblage. Subsequent D2 event enclose EW-striking, tight to isoclinal F2 folds, subtly S-dipping axial planar S2 foliation, hydrothermal fluid flow and gold mineralization in the

Cascavel and Sertão deposits. Additional D2-related structures comprise gently W-plunging L2 intersection lineation between S1 and S2 foliations and shallow EW-trending f2 thrust faults that caused the stacking of the stratigraphy. The preferred orientation of quartz and gold grains subtly plunging west define the Lm2 mineral stretching lineation that implies a general sense of transport in the NNW-

SSE direction. Progressive deformation during D3 resulted in the formation of moderately N-dipping S3 foliation and NW-trending shear zones. Earlier-formed structures are transposed into parallelism within

D3 shear zones. Subtly NW-plunging Lm3 mineral stretching lineation indicate mass-transport in the NW-SE direction. Minor introduction of hydrothermal fluid flow along V2 veins led to the precipitation of paragenetically late galena-pyrite-stibnite±tourmaline assemblage during the D3 event.

Orebodies in the Cascavel and Sertão deposits are oriented according to gently W-plunging F2 hinges

(parallel to L2 intersection lineation) and NW-trending D3 shear zones. The Cascavel deposit consists of folded V2 quartz veins hosted in quartzite. Hydrothermal alteration around V2 gold-bearing veins is characterized by variable (30 m to < 1 m) alteration zones including distal quartz-K- feldspar±muscovite, intermediate quartz-fuchsitic mica-biotite-pyrite, and proximal quartz-white mica- pyrite±chalcopyrite. Minor mineralization hosted by biotite schist displays distal quartz±calcite and proximal quartz-white mica-pyrite±chalcopyrite alteration zones. Gold occurs dominantly as free grains within quartz in the V2 veins. In the Sertão deposit, carbonaceous schist and BIF form distal quartz- ankerite-chlorite-pyrite±chalcopyrite and proximal quartz-siderite/ankerite-pyrite-arsenopyrite- chalcopyrite±pyrrhotite alteration zones. Gold occurs as inclusions and adjacent to arsenopyrite disseminated in proximal alteration zones or hosted in minor quartz-ankerite veins. Detailed mineralogical, structural and textural analyses suggest hydrothermal alteration and gold mineralization in both deposits were synchronous with D2 deformation, with minor to negligible input during subsequent D3 event. Chlorite geothermometry suggests temperatures of 330-410°C and 320-420°C for the circulation of hydrothermal fluids in Sertão and Cascavel deposits, respectively. Arsenopyrite geothermometry from gold-bearing veins in the Sertão deposit yields temperatures between 310-430°C. In order to better understand the tectono-magmatic evolution of southern GAB, whole rock geochemistry with in-situ zircon U-Pb SHRIMP and Hf-O isotopes were conducted in eight regional igneous samples. Sampling included Archean TTGs (Uvá, Caiçara and Paus de Choro orthogneisses), Pink Syenite intrusion in the Faina greenstone belt, besides Neoproterozoic granites from the eastern (Itapuranga suite), and western margins (Rio Caiapó and Serra Negra) of the southern GAB. The petrography, geochemistry, in-situ zircon geochronology and isotope chemistry of these intrusive rocks are presented in Chapter 3. This work included optical petrography, scanning electron microscope (SEM), U-Pb sensitive high resolution ion microprobe (SHRIMP), O secondary ion mass spectrometry (SIMS) and Lu-Hf laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) analyses to constrain the composition, genesis, and evolution of regional magmatism. Secular chemical and zircon Hf-O isotope composition highlight mechanism of differentiation, partial melting depth and changes in magma source. Transition from Na-rich Archean TTGs to K-rich Neoproterozoic arc magmatism is marked by Pink Syenite intrusion. The low REE content and La-Yb ratio of the latter indicate low degree of partial melting predicted for metasomatised slab-derived melts. After > 1000 Ma hiatus, ca. 600 Ma magmatism represents K-rich granites. LILE enrichment, HFSE depletion and negative Ti anomaly characteristic of this rocks are consistent with arc-related magmatism along continental margins. These magmatic events are coeval with peak of collisional

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processes in Western Gondwana. Younger Serra Negra intrusion attest high Sr/Y ratio, low Y content and Rb/Sr suggest Serra Negra derived from deep melting at convergent margins. Zircon Hf-O isotope compositions establish four major groups: (i) 2870-2820 Ma are largely near chondritic and mantle-like, with higher Hf δ18O values with decreasing age, i.e. Paus de Choro granite; (ii) 2060 Ma denote a steep shift to unradiogenic Pink Syenite with mantle-like signature; (iii) 630-610 Ma K-rich granites outline a wide range of unradiogenic Hf and 18O-rich values that culminate with (iv) 530 Ma steep trend for strongly unradiogenic and mantle-like Serra Negra granite. The near chondritic isotope evolution across Archean terranes designate ≤ 70 m.y. of relatively continuous juvenile magmatism that is consistent with early rifting and greenstone belt formation. Transition to more radiogenic signatures at ca. 2820 Ma reflects the onset of subduction near the Meso-Neoarchean boundary prior to collisional orogeny. Following crustal extension, sedimentation and subsequent collision during the Paleoproterozoic favoured reworking and minor recycling attested by Pink Syenite intrusion in the Faina greenstone belt. Analogous syn-tectonic intrusions correlated with the Atlantica assembly enhance the exploration potential, particularly for gold, for this intrusive event. Nearly coeval emplacement of 630-610 Ma K-rich granites with peak metamorphic condition in the Brasiliano/Pan- African orogeny support widespread supracrustal reworking and/or sedimentary recycling. Sharp isotopic shift marks the progression to juvenile magmatism expressed by the 530 Ma Serra Negra granite. Similar chemical and isotopic features of TTG (Archean) and the Serra Negra granite (Cambrian), support a co-parental relation for the two magmatic events. Evolutionary trend suggest mantle extraction at ca. 3300 Ma and link between Archean Uvá TTG and Neoproterozoic Serra Negra granite. Chemical and isotopic testimony reveal fairly constant sources throughout crustal evolution in the southern GAB and implies in formation of continental crust predominantly by in-situ additions rather than to the accretion of allochthonous terranes. This is consistent with isotopic patterns documented in other Archean cratons worldwide and argue for a global process. Combining the various scales of observation applied in the course of this PhD project, chapter four presents an integration of major outcomes obtained in this research with available information to propose an updated and reasonable comprehensive tectono-magmatic scenario for the GAB. Long-term controversy concerning growth of Archean crust is assessed in an attempt to reconcile available models. The data is further used to underpin feasible links between the crustal evolution and mineral systems, particularly gold, in the GAB. Geodynamics reflect geochemical and isotopical shifts suggest reworking of Archean crust along subduction zones followed by protracted generation of juvenile magmatism proposed as potential fingerprints associated with gold systems. Still, the exact mechanisms of hydrothermal fluid generation and transport during magmatism remains a fruitful area of research. This multi-disciplinary PhD research provides a descriptive model for gold mineralization in the Faina greenstone belt, crustal growth evolution and geodynamic settings in an attempt to link magma fertility to potential gold systems in the Goiás Archean Block. Practical implications of this research

are exemplified by regional exploration along cratonic discontinuities and/or margins and the evaluation of igneous rocks emplaced during D2, i.e., synchronous with gold mineralization.

TABLE OF CONTENTS

THESIS DECLARATION…………………………………………………………...…….…..……i

ABSTRACT……………………………………………………………………………...…………ii

TABLE OF CONTENTS………………………………………………………………………..….vi

LIST OF APPENDICES………………………………………………………………………...….xi

TABLE OF FIGURES………………………………………………………………….…...….....xiii

ACKNOWLEDGEMENTS……………………………………………………………………….xvi

AUTHORSHIP DECLARATION: CO-AUTHORED PUBLICATIONS……………………....xviii

Chapter 1: Introduction………………………………………………………………………………...1

1.1. Aims of the research…………………………………………………………………………….3

1.2. Methods and sampling………………………………………………………………..…………4

1.3. Organization of the thesis……………………………………………………………………….4

1.4. Published work………………………………………………………………………...………..5

1.5. THESIS DECLARATION…………………………………………………………...…………6

References………………………………………………………………...... …………………….6

Introduction to Chapter 2: “Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão orogenic gold deposits”………………………………………10

Chapter 2: Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão orogenic gold deposits………………………………………………………..12

Abstract…………………………………………………………………..……………………..….12

1. Introduction………………………………………………………………………………..…….13

2. The Goiás Archean Block………………………………………….…………………………….14

2.1. Exploration and Mining History…………………………….……………………………16

3. Sampling and Methods……………………………………………….………………………….16

4. Geological Setting of the Faina greenstone belt………………………………………………….17

4.1. Lithostratigraphy……………………………...………………………………………….17

4.2. Regional Metamorphism…..…………………………………….…………...…………..19

4.3. Structural Setting……………………………………………………………..…………..19

5. Hydrothermal Alteration and Mineralization at the Cascavel Deposit………………………….22

5.1. Quartzite………………………………………………………………………….………22

5.1.1. V2 veins in quartzite……………………………………………………………25

5.1.2. Wallrock alteration in quartzite……………………………………..…………25

5.2. Biotite schist……………………………………………………………………...………26

5.2.1. V2 veins in biotite schist………………………………………………………..26

5.2.2. Wallrock alteration in biotite schist……………………………………………27

5.3. Cascavel Mineral Chemistry……………………………………………………………...27

5.3.1. Chlorite…………………………………………………………..…………….27

5.3.1.1. Chlorite geothermometry……………………..…………………….28

5.3.2. Biotite………………………………………………………………….………28

5.3.3. White mica…………………………………………………………………….29

5.3.4. Gold-Silver ratio……………………………………………………….………31

6. Hydrothermal Alteration and Mineralization at the Sertão deposit………………………………31

6.1. Carbonaceous schist……………………………………………………………...………31

6.1.1. V2 veins in carbonaceous schist………………………………………..………34

6.1.2. Wallrock alteration in carbonaceous schist………………………...…….…….34

6.2. BIF……………………………………………………………………………………..…35

6.2.1. V2 veins in BIF………………………………………………………..……….35

6.2.2. Wallrock alteration in BIF…………………………………………..…………35

6.3. Sertão Mineral Chemistry……………………………………………………..…….……36

6.3.1. Chlorite………………………………………………………..……………….36

6.3.1.1. Chlorite Geothermometry…………………………………………..36

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6.3.2. White mica………………………………………………………...………….37

6.3.3. Carbonates……………………………………………………………….……38

6.3.4. Arsenopyrite……………………………………….………….………………38

6.3.4.1. Arsenopyrite geothermometry…………………….….…………….39

6.3.5. Gold-silver ratio………………………………………….….……….………..39

7. Discussion……………………………………………………………………………….………41

7.1. Paragenesis……………………………………………………………………………….41

7.2. Redox and pH conditions……...………………………………………………………….41

7.3. Transport and precipitation mechanisms of gold…………………………………..…….42

7.4. Descriptive model for gold mineralization in the Faina greenstone belt…………………43

8. Comparisons with other orogenic gold deposits……………………………...………………….46

9. Conclusions and Implications for exploration……………………………...……………………48

Acknowledgements…………………………………………………………...……………………49

References………………………………………………………………………………………….50

Introduction to Chapter 3: “U-Pb and Hf-O evidence for Mesoarchean crustal growth in central Brazil”………………………...………...……………………………………………………………..57

Chapter 3: U-Pb and Hf-O evidence for Mesoarchean crustal growth in central Brazil ……………...59

Abstract…………………………………………………………………………………………….59

1. Introduction…………………………………………………………………….....……………..60

2. Regional Geology………………………………………………………………………………..61

2.1. Geology of the Archean TTGs - Caiçara, Uvá and Paus de Choro……………………….63

2.2. Geology of the eastern border - Itapuranga I and II……………………………..……….64

2.3. Geology of the western border - Rio Caiapó and Serra Negra……………………………65

2.4. Geology of the Pink Syenite…………………………………………………………...…65

3. Methodology……………………………………………………………………………...……..65

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4. Results…………………………………………………………………………………...………67

4.1. Petrology of granitic rocks…………………………………………………………….…67

4.2. Whole-rock geochemistry………………………………………………………..………70

4.2.1. Major element geochemistry……………………………………….…….……70

4.2.2. Trace and rare earth element geochemistry………………………..…………...72

4.3. In-situ SHRIMP U-Pb zircon geochronology…………………………………………….74

4.3.1. Serra Negra granite……………………………………………………………74

4.3.2. Rio Caiapó granite……...... …………………………………………………..74

4.3.3. Itapuranga Suite…………………………………………..……………………74

4.3.4. Pink Syenite…………………………………………………..………………..77

4.3.5. Caiçara orthogneiss………………………………………………...………….81

4.3.6. Uvá orthogneiss……………………………………..…….…………….81

4.3.7. Paus de Choro granite………………………………………………………….81

4.4. Hf and O isotopes through time...……………………………………………...…………81

5. Discussion……………………………………………………………………………………….83

5.1. Geochemistry and igneous petrogenesis………………………………………..…..…….84

5.2. Geochronological framework and significance of inherited ages………………..….……85

5.3. Zircon Hf-O isotopes and magma sources…………………………….………………….86

5.4. Contentious Archean crustal growth………………………………..…………………….88

5.4.1. Plume-related oceanic plateaux model…………………..…………………….89

5.4.2. Subduction-related model….………..…………………..…………………….90

5.4.3. Implications for the geological evolution of the southern GAB.……………….91

6. Conclusions………………………………………………………………………………….…..96

Acknowledgements……………………………………………………………………………...... 98

References……………………………………………………………………………………….....99

Introduction to Chapter 4: “Linking Au systems to crust-mantle evolution of Archean crust in central Brazil”………………………………………………………………………………………………..116

Chapter 4: Linking Au systems to crust-mantle evolution of Archean crust in central Brazil……….117

Abstract………………………………………………………………………………………...…117 1. Introduction…………………………………………………….……………….…………………117

2. Tectonic Setting…………………………………………………………………….………….119

3. Geology of the Goiás Archean Block (GAB)…………………………………….……………119

3.1. TTG terranes………………….……………………………...…………………..….…..120

3.2. Greenstone belts…………………………………………..………………………...…..121

3.3. Deformation History……………………………………..…………………….....……..124

4. Gold in the Goiás Archean Block……………………………...……………………………..…127

4.1. Regional distribution…………………………………….……………………….……..127

4.2. Deposit Structural Controls……………………………………………………………..128

4.3. Hydrothermal Alteration and Mineralization Styles…………….………………………128

4.4. Fluid conditions……………………………………………………………………..…..131

4.4.1. Gold transport and precipitation mechanisms………………………………...132

4.5. Timing of Au mineralization…………………………………………………..………..134

4.6. Exploration fingerprints…………………………………………………………….…..135

5. Geological evolution of the GAB…………………………………………………………....…135

5.1. Towards a model…..…………………………………….……………………….……..139

5.2. Linking crustal architecture to gold mineralization.……………………..……….……..144

6. Conclusions…………………………………………………...……………………………..…146

References………………………………………………………………………………………...147

Chapter 5: Conclusions ………………...…………..……………………………………….………164

5.1. Lithostratigraphy and structural settings in the Faina greenstone belt …...……..…..…164

5.2. Hydrothermal alteration mineralogy and zonation of the Cascavel and Sertão gold deposits………………………………………………………………………………………………165

5.3. Petrogenesis of intrusive rocks in the southern GAB……………………………..…….165

5.4. Tectonic-magmatic evolution and geodynamics in the GAB ……………………..……166

5.5. Links between crust-mantle evolution and mineral systems ……………………..…….167

5.6. Comparisons with Archean cratons worldwide……………..……………………..……168

6. Future work……………….…………..……………..……...………………..…….….….169

6.1. Regional crustal and structural architecture….………………..…………….….169

6.2. Isotope geochemistry and geochronology of intrusive rocks…………….….….169

6.3. Nature of ore-forming fluids and timing of gold mineralization in the GAB..…169

References……………………………………………………………………………….…..170

LIST OF APPENDICES

Chapter 2

Table A2.1. Microprobe analyses of hydrothermal silicates at the Cascavel deposit………..………176

Table A2.2. Microprobe analyses of hydrothermal chlorite at the Cascavel deposit……..………….178

Table A2.3. Microprobe analyses of gold and sulfides at the Cascavel deposit………………………179

Table A2.4. Microprobe analyses of hydrothermal silicates at the Sertão deposit……..……….……180

Table A2.5. Microprobe analyses of hydrothermal chlorite at the Sertão deposit………..…….……181

Table A2.6. Microprobe analyses of gold and sulfides at the Sertão deposit……………...….………182

Chapter 3

Methodology………………………………….……………………..……………………………….183

Table A3.1. Mean weighted analyses of standards used for U-Pb geochronology……...... 184

Table A3.2. Results of standard materials used for the Lu-Hf analyses………….……...... 185

Table A3.3. Results of standard materials used for oxygen analyses…………….……...... 187

Fig. A3.1. Bi-log diagrams of compatible versus incompatible elements……………………………189

Fig. A3.2. εHf(t) versus geochemical signatures………………………..……………………………190

Fig. A3.3. δ18O versus U-Pb dataset...………….………………………..……………………………191

Table A3.4. Whole-rock geochemistry…………….…………..……………………...... 192

Fig. A3.5. CL images of Itapuranga I zircons……………………….…..……………………………194

Fig. A3.6. CL images of Itapuranga II zircons…………………………..……………………………195

Fig. A3.7. CL images of Rio Caiapó zircons……………………….…..……………………………196

Fig. A3.8. CL images of Serra Negra zircons……………………….…..……………………………197

Fig. A3.9. CL images of Pink Syenite zircons……………………….….……………………………198

Fig. A3.10. CL images of Pink Syenite zircons………………...……….…..……..…………………200

Fig. A3.11. CL images of Paus de Choro zircons………………………………………..……………201

Fig. A3.12. CL images of Uvá zircons…………..………………….…..……………………………202

Fig. A3.13. CL images of Caiçara zircons...………..……………….…..……………………………203

Fig. A3.14. Results of zircon U-Pb analyses.……………………….…..……………………………204

Fig. A3.15 Weighted average plots for U-Pb analyses.………..…….…..……………………………205

Table A3.5. Results of zircon U-Pb analyses.……………………….…..……………………………206

Table A3.6. Results of SIMS oxygen analyses….………………….…..……………………………210

Chapter 4

Table A4.1. Previous deformation history of the GAB….………….…..……………………………213

Chapter 5

Petrography, SEM and EPMA……………………………………………………………………….214

Fig. A5.1. Hand-sample and photomicrographs of arsenopyrite-bearing samples…...... 215

Fig. A5.2. EPMA compositional zoning of arsenopyrite grains……...... 216

LA-ICP-MS analyses………………………………………………………………………………...217

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Fig. A5.3. BSE images of arsenopyrite mounts…………………………………...... 218

Fig. A5.4. LA-ICP-MS elemental map of arsenopyrite grains…………………...... 219

Fig. A5.5. LA-ICP-MS elemental map of arsenopyrite grains…………………...... 220

Fig. A5.6. LA-ICP-MS elemental map of arsenopyrite grains…………………...... 221

Fig. A5.7. LA-ICP-MS elemental map of arsenopyrite grains…………………...... 222

Table A5.1. Results of LA-ICP-MS analyses...... 223

TABLE OF FIGURES

Chapter 2

Fig. 1 Geology of the Brasília fold belt, central Brazil...... 13

Fig. 2 Geological map of the Faina greenstone belt...... 18

Fig. 3 Main structural features observed in the Faina greenstone belt.……..…………...... 21

Fig. 4 Geological map and cross section of the Cascavel deposit………...... 23

Fig. 5 Hydrothermal alteration diagram for the Cascavel deposit...... 24

Fig. 6 Mineral chemistry at the Cascavel deposit...... 29

Fig. 7 Cascavel gold grains…………………...... 30

Fig. 8 Geological map and cross section of the Sertão deposit...... 32

Fig. 9 Hydrothermal alteration diagram at the Sertão deposit...... 33

Fig. 10 Proximal alteration samples and gold from the Sertão deposit……...... 37

Fig. 11 Mineral chemistry at the Sertão deposit…..……...... 39

Fig. 12 Elemental maps show asenopyrite zonation….…..………...... 40

Fig. 13 Proposed geological evolution of the Faina greenstone belt...... 45

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Chapter 3

Fig. 1 Geology of the Brasília fold belt…...... 62

Fig. 2 Geological map of the southern GAB………………………………………………………….…..64

Fig. 3 Hand sample and photomicrograph of Archean/ Paleoproterozoic rocks in the southern GAB…68

Fig. 4 Hand sample and photomicrographs of Neoproterozoic rocks in the southern GAB……...... 69

Fig. 5 Geochemistry-based classification diagrams.....……………..………………...... 70

Fig. 6 Harker diagrams...... 71

Fig. 7 Trace element and REE patterns………………….….………………………………………………73

Fig. 8 CL images of representative magmatic zircons……………………...…...... 77

Fig. 9 U-Pb concordia diagrams with inheritance indicated……………...... 80

Fig. 10 εHf-time plot of magmatic zircons of the southern GAB………………...... 82

Fig. 11 δ18O-time plot of magmatic zircons of the southern GAB………………...... 83

Fig. 12 Proposed geological evolution in the southern GAB...... 93

Chapter 4

Fig. 1 Geology of the Brasília fold belt...... 118

Fig. 2 Stratigraphy of the greenstone belts of the GAB…………...... 122

Fig. 3 Geology of the northern and southern greenstone belts of the GAB…………...... 123

Fig. 4 Styles of mineralization in orogenic gold deposits in the GAB……………...... 130

Fig. 5 Ore assemblage, occurrence and distribution of gold in the GAB deposits………...... 133

Fig. 6 Schematic block diagram evolution for the northern GAB...... 137

Fig. 7 Schematic block diagram evolution for the southern GAB...... 138

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Fig. 8 Map showing available εNd/Hf data in the Goiás Archean Block...... 140

Fig. 9 Diagrams with isotopic dataset and histogram with crystallization ages for the GAB...... 143

ACKNOWLEDGEMENTS

The realization of this project counted with the help of many people, who contributed in different ways and to whom I will always be grateful.

Thanks to my supervisors for the opportunity to work and learn with you. These last years have definitely shaped me. Steffen, thank you for accepting being my tutor and for all the support during this time. Along with an energic leading, is beautiful to see how much effort and devotion you give to your students. Tony, your patience and genuineness are striking. Thanks for the blunt honesty, for the freedom to think and create ideas, and for the good laughs. To Hardy Jost and Rocky Osborne, who despite not being my supervisors have helped either during field work or by making themselves approachable for discussion always when needed.

I am very grateful to the CET people I have met during the past four years. You have significantly enriched my life. Thank you so much for making things easier during such an intense time. I could never have done without you. To my CET mates: Greg Poole, Eun Joo Choi, Mike Tedeschi, Eliza Smith, Quentin Masurel, Laura Petrella, Linda Iaccheri, Katarina Björkman, Jason Bennet, Perla Pina, Marcelo Godefroy, Erika Santiago, Nico Meriaud, Jordan McDivitt, thanks to all of you for the good times and all the wine. It was awesome to be part of such a dynamic and varied research group as CET. I will take a bit of each of you with me, thanks for the love and unpretentious help.

To the “Conselho Nacional de Desenvolvimento Científico e Tecnológico – CNPq” for funding this research (Project n° 207220/ 2014-0). The local mining company, Orinoco Gold Limited, is thanked for their permission to publish information and for the support given during the field work campaigns.

To the ones who had kindly helped by providing me with valuable and patient revisions, including Wally Witt, Linda Iaccheri, Quentin Masurel, Eliza Smith, Mike Tedeschi, Daniel Wiemer, Lydia Lobato and Tyler Baril.

To my beloved friends that cheered me up after returning to finish this project back in Brasília: Tuquicamata (Júlia Pêra), Soni (Fellipe Lopes), Loo (Luana Fonseca), Zéloi Guimarães, Federico Cuadros, and Marcelo Matos. Special mention to Rafael Brant and all the Jazida crew. At your own way, each of you made a difference for this to come true.

Marcos and Cristina Bogossian, thank you for the unconditional love, for cultivating our curiosity, and for letting us pursue what we are passionate about. There are no words to thank you both enough. To my siblings Picrits (Paula Bogossian), Colungas Grill (Bianca Bogossian), David Bogossian and Eduarda Fischer for all the care and fun. At last but not least, I thank nana Kasukiari (Wanda Miranda), who unfortunately is not here anymore to see this, for the legacy of courage and strength.

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AUTHORSHIP DECLARATION: CO-AUTHORED PUBLICATIONS

This thesis contains work that has been prepared for publication.

“Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão orogenic gold deposits”. This text has been prepared for publication in Ore Geology Reviews.

Location in the Thesis: Chapter 2

The candidate wrote 100% of the text, carried out all field work and selected all samples utilized in the study. Interpretation of petrographic was undertaken entirely by the candidate. Editing and revisions were carried out by Prof. S.G. Hagemann.

Co-author signatures and ates:

06/05/2019

“U-Pb and Hf-O evidence for Mesoarchean crustal growth in central Brazil”. This text has been prepared for publication in journal Precambrian Research.

Location in the Thesis: Chapter 3

The candidate wrote 100% of the text, carried out all field work and selected all samples utilized in the study. Design of the research, sample selection and sample preparation was done by the candidate. The LA-ICP-MS data were acquired with instrument setup by Dr. A.I.S. Kemp, who also handled data reduction. Oxygen isotopes measurements and data reduction were performed under the supervision pf Dr. H. Jeon. Interpretation of all data was by the candidate with the assistance of Dr. A.I.S. Kemp.

Co-author signatures and dates:

06/05/2019

“Linking Au systems to crust-mantle evolution of Archean crust in central Brazil”. This text has been prepared for publication in the “Journal of South America Earth Sciences”.

Location in the Thesis: Chapter 4

The candidate wrote 100% of the text, with editing and revisions by S.G. Hagemann. The LA- ICP-MS data were acquired by the candidate with instrument setup and data reduction by Paul Olin

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and Prof Leonid V. Danyushevsky (CODES, Tasmania).

Co-author signatures and dates:

06/05/2019

Student signature:

Date: 6th, May 2019

I, Prof. S.G. Hagemann, certify that the student statements regarding their contribution to each of the works listed above are correct.

Coordinating supervisor signature:

Date:

06/05/2019

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Chapter 1: Introduction

Greenstone belts are one of the oldest records of the lithosphere’s geologic evolution and they have been widely associated with substantial ore deposits worldwide (Wit and Ashwal, 1995). The 70 km-long Faina greenstone belt is enclosed in the southern Goiás Archean Block (GAB), approximately 400 km west from the capital Brasília. Gold deposits hosted in the Goiás Archean Block include the Serra Grande mine (~ 7 Moz, Anglo Gold), Pilar (0.7 Moz) and Guarinos (6.5 Moz, both owned by Yamana Gold Inc.). Typically, gold mineralization in the Goiás Archean Block is hosted within greenschist facies metasedimentary and metavolcanic rocks, and are structurally controlled by brittle- ductile shear zones, including low angle thrust faults and strike-slip faults developed during the Paleoproterozoic (Pulz, 1995; Queiroz, 2000; Jost et al., 2008; Marques et al., 2013; Jost et al., 2014), and to a lesser extent in the Neoproterozoic (Rodrigues, 2011). The Faina greenstone belt area records several minor artisanal workings active for over 250 years, which exploited mostly alluvial placer deposits with an estimated production of ~ 3 Moz (Silva and Rocha, 2008). The largest known deposits in the Faina greenstone belt are Cascavel and Sertão, ca. 28 km apart from each other. They are hosted in greenschist facies volcano-sedimentary rocks and are characterized by structurally controlled gold-bearing quartz±carbonate veins and disseminated gold in sulfide-rich layers. This study provides the first systematic description of gold mineralization in the Faina greenstone belt, based on a dataset obtained from the Cascavel and Sertão deposits. The investigation helps better understand the nature, controls and genesis of the gold mineralization and it allows correlation with similar deposits worldwide, aiding future exploration efforts in the region. Extensive work is available for gold deposits of the nearby Crixás (Thomson, 1987, 1991; Thomson and Fyfe, 1990; Fortes et al., 1996, 1997, 2003; Portocarreto, 1996; Petersen Jr., 2003; Tassinari et al., 2006), Pilar de Goiás (Pulz, 1995; Santos et al., 2008), and Guarinos (Pulz, 1990; Pulz et al., 1993, Michel et al., 1994; Rodrigues, 2011) greenstone belts. In contrast, only limited studies emphasised the Faina greenstone belt (Resende et al., 1998, 1999; Carvalho et al., 2013; Borges et al., 2017). Historical and current gold exploration and mining, exposed open pit walls, accessible underground mine workings and extensive diamond drill core promoted a unique opportunity to investigate the lithological, structural, petrological and mineralogical characteristics of the Cascavel and Sertão deposits. At Cascavel, vein-hosted mineralization controlled by F2 hinges is hosted by quartzite, with minor disseminated mineralization in biotite schist parallel to D3 shear zones. At Sertão, gold is mostly disseminated within proximal alteration zones controlled by D3 shear zones and hosted by carbonaceous schist, with minor vein-hosted mineralization hosted in BIF. Nature and mechanisms controlling craton generation are the focus of enduring debate (Taylor, 1967; Armstrong, 1991; Taylor and McLennan, 1985; Hawkesworth et al., 2010, 2013; Kemp et al., 2009). These terranes testimony the transition from an overall high geothermal gradient regime with

massive production of continental crust into a regime of lower and more heterogeneous geothermal gradients marked by the predominance of crustal reworking, e.g., at the Archean-Proterozoic boundary (Taylor and McLennan, 1985; Martin, 1986, 1993, 1994; Sylvester, 1994). This is illustrated by the chemical variation of TTGs throughout the Archean (Martin and Moyen, 2002), and by the contrasting K-rich post-Archean granitic rocks derived from partial melting of older tonalitic crust with a variable contribution of mantle-derived magmas (López et al., 2006). Craton margins outline intensely reworked major lithospheric discontinuities that underpin diverse tectonic settings (Kirkland et al., 2015). Convergent plate boundaries, for example, can devise: (i) juvenile accretion during subduction with magma production by partial melting of fertilized mantle wedge (Martin, 1986), (ii) recycling of crustal material by subduction (Veizer and Jansen, 1979; Von Huene and Scholl, 1991), or by delamination of lithospheric roots (Hawkesworth et al., 2010), and (iii) reworking of pre-existing crust due to deformation, metamorphism and partial melting (Hawkesworth et al., 2010). Thus, cratons and their marginal orogens make up important sites to investigate the secular changes recorded throughout the formation and evolution of the continental crust. The complex tectono-magmatic evolution of the Goiás Archean Block has been addressed by a number of studies (Danni et al., 1986; Queiroz, 2000; Pimentel et al., 2003; Jost et al., 2005, 2013; Queiroz et al., 2008; Beguelli Jr., 2012), however, most research has been concentrated in the northern portion of the craton, with comparatively little work carried out in the south. The still unclear and intricate evolution of the southern GAB (Beguelli Jr., 2012; Pimentel et al., 2003; Jost et al., 2005, 2013), make this area an attractive case study to track magma sources and to evaluate geodynamic changes through time. An investigation involving combined whole-rock geochemistry, in-situ zircon U-Pb and Hf-O isotopes is used to pinpoint the nature, settings and sources of regional magmatic events in the southern GAB. The outcomes of this study provide evidence for Archean crustal growth marked by successive magmatic additions. This challenges disparate evolutionary paths predicted for the collage of exotic terranes (e.g. Jost et al., 2010, 2014). Isotopic consonance across terranes suggests successive magmatic input shared fairly similar sources that imply an autochthonous growth of the southern GAB. Combining the various scales of observation conducted during the course of this study, Chapter 4 presents a holistic overview of the geology related to gold deposits and models for the crust-mantle evolution of terranes in the GAB. Due to the significant control of crustal architecture on the development and distribution of ore deposits (Begg et al., 2009), links between crustal growth and gold endowment are further explored for the GAB. This contribution provides insights into the potential mineral systems available in the area and may enhance regional exploration. The timing of gold mineralization in the Goiás Archean Block is also poorly defined. Early geochronology based on Rb-Sr, K-Ar and Ar-Ar methods (Fortes, 1996; Fortes et al., 1997) has proven unreliable in constraining the age of mineralization from deposits hosted in the northern Goiás Archean Block due to effects from subsequent thermal events. More robust Pb-Pb on galena and Re-Os on

arsenopyrite (Pulz, 1990; Marques et al., 2013) have improved the understanding of the timing of gold mineralization event in the northern Goiás Archean Block. Until the moment, no geochronology is available for gold mineralization in the southern Goiás Archean Block. Therefore, there is abundant room for further progress in determining the timing and evolution of gold systems in the Goiás Archean Block, and particularly their link with regional tectono-magmatic events. Hydrothermal minerals related to gold mineralization in the Faina greenstone belt, e.g., arsenopyrite, provide a potential opportunity to constrain the age of mineralization. The integration of this data in the geological framework of the Goiás Archean Block may provide insights on the geodynamic history and gold systems.

1. Aims of the research The aims of this thesis are to further understand gold mineralization in the Faina greenstone belt, nature, mechanisms and sources prescribed for regional magmatic events, and possible links of geodynamics and the distribution and formation of gold deposits in southern Goiás Archean Block based on: (i) characterization of hydrothermal alteration and mineralization in the context of the deformation history of the area, (ii) propose a descriptive model for gold mineralization associated with the Cascavel and Sertão deposits, (iii) constraint the crustal evolution of the igneous rocks in the area and (iv) test the possible correlations with gold systems. These issues are addressed via multi- disciplinary investigation in deposit and regional scales, summarized in Chapters 2 to 4 with a brief description presented below. Chapter 2 presents a deposit scale study of gold mineralization at the Cascavel and Sertão deposits. This contribution focused on the hydrothermal alteration, lithological and structural controls of gold mineralization. Detailed core logging of representative mineralized zones in both deposits was carried out and followed by representative sampling collected for petrographic analyses and mineral chemistry. Optical petrography (transmitted and reflected) was used to characterize the mineralogy of major host rocks, micro-structures and textures. Electronic semi- and quantitative analyses (SEM and EMPA, respectively) were conducted in hydrothermal minerals to estimate composition and conditions of fluids associated with mineralization, e.g., chlorite and arsenopyrite geothermometry. Chapter 3 presents a combination of whole-rock geochemistry, in-situ zircon SHRIMP U-Pb and Hf-O isotopes from eight selected intrusive rocks to constraint settings, timing and sources of magmatism within the southern Goiás Archean Block area. Dataset provided by the geochemistry, geochronology and isotopic signature of intrusive rocks help better understand regional deformation events and geodynamic settings. Results presented in this chapter reveal distinct magmatic events shared similar sources during the evolution of the southern Goiás Archean Block and implies in a cogenetic relation of these terranes. This part of the research utilized U-Pb sensitive high resolution ion microprobe (SHRIMP) on zircon to establish the timing of regional magmatism. The syn- to late nature of Pink Syenite intrusion in the Faina greenstone belt is used to constrain the timing of deformation

events. The role of juvenile mantle-derived magmatism associated with accretionary orogens as potential sources for gold systems is also explored in this contribution. Chapter 4 presents a combination of analytical results from Chapters 2 and 3 and incorporates them into the geological context of the gold systems hosted in the Goiás Archean Block. This Chapter presents a temporal and spatial framework of magmatism, deformation and gold mineralization constraints based on available data of the Goiás Archean Block and results obtained in this study. Contentious formation of Archean crust is assessed to reconcile available models. The chemical, geochronological and isotopic dataset provided in this contribution is further used to delineate possible links between crust-mantle evolution with the formation and distribution of mineral deposits, especially gold. Detailed studies, e.g. EPMA, LA-ICP-MS analysis, of gold-associated arsenopyrite disseminated in proximal hydrothermal assemblages at the Sertão deposit were conducted in an attempt to constrain the timing of gold mineralization (please see appendices of Conclusions and future work chapter for details). However, the unfeasibility of arsenopyrite grains for the Re-Os method prevented to bracket the timing of gold mineralization event in the Faina greenstone belt.

2. Methods and sampling Samples for this project included 95 drill core specimens of least altered to mineralized material from each major lithology from the Cascavel and Sertão deposits for the deposit scale study. Samples were selected during logging of representative sections on four drill cores at each deposit. In addition, eight intrusive rock samples from surface outcrops were collected for the regional study, including samples for geochronology. Two fieldwork campaigns were conducted during the first two years of the PhD project. Analytical methods utilized in this study are described in their respective chapter and they included: - Quantitative electron microprobe analyses (EPMA) on hydrothermal silicates (e.g. micas, carbonates), sulfides and gold (Chapter 2) - Whole-rock major element oxide, trace and rare earth element analyses (Chapter 3) - U-Pb SHRIMP geochronology on zircons (Chapter 3) - Lu-Hf LA-ICP-MS analyses on zircons (Chapter 3) - Oxygen isotope analyses using the Cameca 1280 SIMS on zircon (Chapter 3) - LA-ICP-MS trace element analyses on arsenopyrite (Appendix Chapter 5)

3. Organization of the thesis According to the regulations for Research Higher Degrees of the University of Western Australia, a thesis may comprise a series of peer reviewed journal publications. As listed in the Post Graduate handbook of UWA, regulation 32.1 declares that a thesis may be presented as a typescript, published

book, a single or a series of papers that have been published in a journal, papers submitted for publication but not accepted as yet, or a paper to be submitted. This thesis is presented as a series of manuscripts with each of those consisting of a chapter of the thesis, which includes corresponding references. The Chapters are sequentially distributed with a transition page consisting of a brief overview of the content and contributions of co-author. Chapter 2 has been accepted for publication, whereas Chapters 3 and 4 will be submitted for publication after submission of the thesis. They include: Chapter 1: Introduction Chapter 2 (manuscript 1): “Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão gold deposits” Chapter 3 (manuscript 2): “U-Pb and Hf-O evidence for Mesoarchean crustal growth in central Brazil” Chapter 4 (manuscript 3): “Linking Au systems to crust-mantle evolution of Archean crust in central Brazil” Chapter 5: Conclusions and future work

4. Published work During the period of this project, the PhD candidate published two extended abstracts/posters and one unpublished abstract/poster. Additionally, the first manuscript was accepted for publication in high impact international journal. 2016 Bogossian, J., Hagemann, S.G., Kemp, A., Lobato, L.M. Hydrothermal alteration mineralogy and zonation of the Cascavel and Sertão deposits, Faina greenstone belt in Goiás, Brazil. Unpublished abstract/poster for the Gordon Research Conference, 2016 Jun 19-24; Les Diablerets, Switzerland. 2018 Bogossian, J., Hagemann, S.G., Rodrigues, V.G., Roberts, M., Lobato, L.M. Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão gold deposits, Goiás, central Brazil. In: Proceedings of the SEG 2018 Conference. Metals, Minerals, and Society. 2018 Sep 22-25; Keystone, United States. 2019 Bogossian, J., Kemp, A., Hagemann, S.G. Crustal Evolution in the Southern Goiás Archean Block: Evidence from In Situ U-Pb and Hf-O Isotopes. In: Proceedings of the SEG 2019 Conference. South American Metallogeny: Sierra to Craton. 2019 Oct 7-9; Santiago, Chile. Bogossian, J., Hagemann, S.G., Rodrigues, V.G., Lobato, L.M., Roberts, M., 2019. Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão orogenic gold deposits. Ore Geology Reviews. In press.

5. THESIS DECLARATION I, Jessica Bogossian, declare that I am responsible for all portions and contents of this thesis except for collaborative work outlined in Chapters 2, 3 and 4. I have not presented any material that is subject of this thesis for a graduate degree at this or any other University.

References Armstrong, R.L., 1991. The persistent myth of crustal growth. Australian Journal of Earth Sciences, 38:5, 613-630. DOI: 10.1080/08120099108727995 Borges, C.C.A., Toledo, C.L.B., Silva, A.M., Chemale Jr., F., Host, H., Lana, C.C., 2017. Geochemistry and isotopic signatures of metavolcanic and metaplutonic rocks of the Faina and Serra de Santa Rita greenstone belts, Central Brazil: Evidences for a Mesoarchean intraoceanic arc. Precambrian Research 292:350-377. Brant, R.A.P., Stremel, R.B., Souza, V.S., Neto, L.R., Rodrigues, V.G., Carvalho, M.J., Araújo, K.C., Jost, H., 2014. Mineralização aurífera Curral de Pedra, greenstone belt de Faina, Goiás. In: Proceedings of the 47° Brazilian Congress of Geology, Salvador, Bahia, p. 1643. Carvalho, M.J., Rodrigues, V.G., Jost, H., 2013. Formação Arraial Dantas: depósito aurífero detrítco glaciogênico do greentone belt de Faina, Goiás. In: Proceedings III Brazilian Symposium on Metallogeny, 2013 2-5 Jun; Gramado, RS, Brazil. Fortes, P.T.F.O., 1991. Geologia do Depósito Aurífero Mina III, Crixás, Goiás. Unpublished MSc Dissertation, University of Brasília, p. 194. Fortes, P.T.F.O., Pimentel, M.M., Teixeira, W., 1993. Geocronologia Rb-Sr das rochas encaixantes do depósito aurífero Mina III, Goiás. In: SBGq, 4 Brazilian Congress of Geochemistry, Niterói, Extended Abstract Volume, p. 250-252 Fortes, P.T.F.O., Giuliani, G., Takaki, T., Pimentel, M.M., Teixeira, W., 1995. Aspectos geoquímicos de depósito aurífero Mina III, greenstone belt de Crixás, Goiás. Geochi. Brasiliensis, 9:13-31. Fortes, P.T.F.O., 1996. Metalogênese dos depósitos auríferos Mina III, Mina Nova e Mina Inglesa, Greenstone Belt de Crixás, GO. Unpublished PhD Thesis, University of Brasília, p. 176. Fortes, P.T.F.O., Cheilletz, A., Giuliani, G., Féraud, G., 1997. A Brasiliano age (500 ± 5 Ma) for the Mina III gold deposit, Crixás Greenstone Belt, Central Brazil. Internat. Geol. Rev., 39:449-460 Fortes, P.T.F.O., Pimentel, M.M., Santos, R.V., Junges, S., 2003. Sm-Nd study of the Crixás greenstone belt, Brazil: implications for the age of deposition of the upper sedimentary rocks and associated Au mineralization. J. South Am. Earth Sci., 16:503-512. Hawkesworth, C.J., Dhuime, B., Pietranik, A.B., Cawood, P.A., Kemp, A.I.S., Storey, C.D., 2010. The generation and evolution of the continental crust. Journal of the Geological Society of London 167, 229–248. Hawkesworth, C., Cawood, P., Dhuime, B., 2013. Continental growth and the crustal record. Tectonophysics 609, p. 651–660. http://dx.doi.org/10.1016/j.tecto.2013.08.013

Jost, H. and Fortes, P.T.F.O., 2001. Gold deposits ad occurrences of the Crixás Goldfield, central Brazil. Mineralium Deposita 36: 358-376. Jost, H., Carvalho, M.J., Rodrigues, V.G., Martins, R., 2014. Metalogênese dos Greenstone belts de Goiás. In: Silva, M.G., Neto, M.B.R., Jost, H., Kuyumjian, R.M. (Orgs.), Metalogênese das Províncias Tectônicas Brasileiras, Belo Horizonte, CPRM, p. 141-168. Kemp, A.I.S., Hawkesworth, C.J., Collins, W.J., Gray, C.M., Blevin, P.L., 2009. Isotopic evidence for rapid continental growth in an extensional accretionary orogen: the Tasmanides, eastern Australia. Earth Planet. Sci. Lett. 284, 455–466. doi:10.1016/j.epsl.2009.05.011 Kirkland, C.L., Spaggiari, C.V. Smithies, R.H., Wingate, M.T.D., Belousova, E.A., Gréau, Y., Sweetapple, M.T., Watkins, R., Tessalina, S., Creaser, R., 2015. The affinity of Archean crust on the Yilgarn—Albany–Fraser Orogenboundary: Implications for gold mineralisation in the Tropicana Zone. Precambrian Research 266, p. 260–281. Kuyumjian, R.M. and Costa, A.L.L., 1999. Geologia, geoquímica e mineralizações auríferas da seqüência Mina Inglesa, greenstone belt de Crixás, Goiás. Brazilian Journal of Geology 29(3):313-318. López, S., Fernández, C. and Castro, A., 2006. Evolution of the Archaean continental crust: Insights from the experimental study of Archaean granitoids. Current Science, 91(5), 607-621. Retrieved from http://www.jstor.org/stable/24094364 Marques, J.C., Jost, H., Creaser, R.A., Frantz, J.C., Osorio, R.G., 2013. Age of arsenopyrite gold- bearing massive lenses of the Mina III and its implications on exploration, Crixás greenstone belt, Goiás, Brazil. In: III Brazilian Symposium of Metallogeny, Gramado, extended abstracts. Martin, H., 1986. Effect of steeper Archean geothermal gradient on geochemistry of subduction-zone magmas. Geology 14, 753–756. Martin, H., 1993. The mechanisms of petrogenesis of the Archaean continental crust comparison with modern processes. Lithos 30, 373–388. Martin, H., 1994. The Archean grey gneisses and the genesis of the continental crust. In: Condie, K.C. (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 205–259. Martin, H. and Moyen, J.-F., 2002. Secular changes in TTG composition as markers of the progressive cooling of the Earth. Geology 30, 319–322. Michel, D., Giuliani, G., Pulz, G.M., Jost, H., 1994. Multistage gold deposition in the Archean Maria Lázara gold deposit (Goiás, Brazil). Mineral. Deposita 29, 94-97. Montalvão, R.M.G., 1985. Evolução geotectônica dos terrenos granitóide-greenstone belts de Crixás, Guarinos, Pilar de Goiás-Hidrolina (Goiás). Unpublished PhD Thesis, University of São Paulo, p. 375. Petersen Jr., K.J.P., 2003. Estudo das mineralizações auríferas do Corpo IV e V da Estrutura IV do greenstone belt de Crixás, Goiás. Unpublished PhD Thesis, University of São Paulo, p. 175.

Portocarrero, J.L.P., 1996. Geologia da jazida aurífera Mina Nova, greenstone belt de Crixás, Goiás. Unpublished MSc Dissertation, University of Brasília, p. 109. Pulz, G.M., 1990. Geologia do depósito aurífero tipo Maria Lázara (Guarinos, Goiás). Unpublished MSc Dissertation, University of Brasília, p. 139. Pulz, G.M., Martins, E.S., Fuck, R.A., 1992. Morfologia dos minerais de ouro e arsenopirita no depósito Maria Lázara (Guarinos, Goiás) e suas implicações nos mecanismos de deposição. Brazilian Journal of Geology 22(3): 257-261. Pulz, G.M., Jost, H., Giuliani, G., Michel, D., 1993. Evidências mineralógicas e estruturais da percolação episódica de fluidos hidrotermais no depósito aurífero Maria Lázara, Goiás. Proceedings of the Brazilian Academy of Sciences 65:19-28. Pulz, G.M., 1995. Modelos prospectivos para ouro em greenstone belts: exemplo dos depósitos Maria Lázara e Ogó, na região de Guarinos e Pilar de Goiás, Goiás. Unpublished PhD Thesis, University of Brasília, p. 189. Queiroz, C.L., 2000. Evolução Tectono-Estrutural dos Terrenos Granito-Greenstone Belt de Crixás, Brasil Central. Unpublished PhD Thesis, University of Brasília, p. 209. Queiroz, C.L., Jost, H., Silve, H., McNaughton, N.J., 2008. U-Pb SHRIMP and Sm-Nd geochronology of granite-gneiss complexes and implications for the evolution of the central Brazil Archean terrain. Journal of South American Earth Sciences 26, 100-124. Resende, M.G., Jost, H., Osborne, G.A., Mol, A.G., 1998. Stratigraphy of the Goiás and Faina greenstone belts, Central Brazil: A new proposal. Brazilian Journal of Geology 28 (1):77-94. Resende, M.G., Jost, H., Lima, B.E.M., Teixeira, A.A., 1999. Proveniência e idades modelo Sm-Nd das rochas siliciclásticas arqueanas dos greenstone belts de Faina e Santa Rita, Goiás. Brazilian Journal of Geology, 29:281-290. Rodrigues, V.G., 2011. Geologia do depósito aurífero do Caiamar, greenstone belt de Guarinos: um raro depósito associado a albitito sódico. Unpublished MSc Dissertation, University of Brasília, p. 79. Santos, R.V., Oliveira, C.G., Souza, V.H.V., Carvalho, M.J., Andrade, T.V., Souza, H.G.A., 2008. Correlação isotópica baseada em isótopos de Carbono entre os greenstone belts de Goiás. In: SBG, Proceedings 44° Brazilian Congress of Geology, Curitiba, p. 52. Silva, M.P. and Rocha, C., 2008. A caracterização da mineração aurífera em Faina, Goiás, em um contexto ambiental histórico e atual. Ambiente & Sociedade, Campinas (SP), v. XI, n. 2 p. 373- 388. Sylvester, P.J., 1994, Archean granite plutons, in K.C. Condie, ed., Developments in Precambrian geology, Archean Crustal Evolution: Elsevier Amsterdam, p. 261–315. Taylor, S.R., 1967. The origin and growth of continents. Tectonophysics 4, 17–34. Taylor, S.R. and McLennan S.M., 1985. The continental crust: its composition and evolution. Blackwell, Oxford, p. 312.

Tassinari, C.G., Jost, H., Santos, J., Nutman, A., Bennell, M.R., 2006. Pb and Nd isotope signatures and SHRIMP U-Pb geochronological evidence of Paleoproterozoic age for Mina III gold mineralization, Crixás District, Central Brazil. 5th South American Symposium on Isotope Geology, Punta Del Este, Uruguay, Short Papers Volume, p. 527-529. Thomson, M.L., 1987. The Crixás Gold deposit; Brazil: metamorphism, metassomatism and Gold mineralization. Unpublished PhD Thesis, University of Western Ontario, Canada, p. 345. Thomson, M.L. and Fyfe, W.S., 1990. The Crixás Gold Deposit, Brazil: Thrust-Related Postpeak Metamorphic Gold Mineralization of Possible Brasiliano Cycle Age. Econ. Geol., 85(5): 928- 942. Veizer, J., Jansen, S.L., 1979. Basement and sedimentary recycling and continental evolution. J. Geol. 87, 341–370. Von Huene, R., Scholl, D.W., 1991. Observations at convergent margins concerning sediment subduction, subduction erosion and the growth of continental crust. Rev. Geophys. 29, 279– 316.

Introduction to Chapter 2: “Hydrothermal alteration and mineralization in the Cascavel and Sertão gold deposits in the Faina greenstone belt, central Brazil”

This chapter presents the results of an investigation on the hydrothermal alteration related to gold mineralization in the Cascavel and Sertão deposits, both located in the Faina greenstone belt. It was submitted to the Journal Ore Geology Reviews with the candidate J. Bogossian as the first author, followed by S. Hagemann (Centre for Exploration Targeting, University of Western Australia), V. Rodrigues (Orinoco Gold Ltda.), M.P. Roberts (Centre for Microscopy, Characterisation and Analysis, University of Western Australia) and L.M. Lobato (Universidade Federal de Minas Gerais) as co- authors.

The Faina greenstone belt area has been recently and historically been explored for gold, however, it lacks a systematic description of the geological setting, structural controls and hydrothermal alteration associated with gold mineralization. Compared to the three other similar, but more well- known, belts of northern Goiás Archean Block (GAB), limited understanding of gold systems in the Faina greenstone belt can in part explain why the area is still underexplored. Therefore, the Faina greenstone belt is an interesting target to study the characteristics and controls of hydrothermal alteration and associated gold mineralization in order to aid future exploration efforts. The research included detailed logging of diamond drill cores from the Cascavel and Sertão deposits to define host rocks, structural controls, mineralogy and zonation of Au-related hydrothermal alteration. The structures and associated veins under investigation during diamond drill core logging were correlated with structural measurements taken from underground mine workings at Cascavel and in an old open pit at Sertão deposit. Sampling targeted representative hydrothermal alteration zones, host rocks and vein types for petrographic analysis. Mineralogical changes through hydrothermal alteration zones were described using optical petrography (transmitted and reflected), in addition to quantitative and semi- quantitative analyses via electron microscopy (SEM and EPMA). In-situ analyses using EPMA aimed hydrothermal minerals of each alteration zone and host rock, including micas, carbonates, feldspars, sulfides and gold. Chlorite and arsenopyrite geothermometry were utilized to estimate temperature conditions of the hydrothermal fluids responsible for gold mineralization in the two deposits. The collected dataset was used to propose a descriptive model for gold mineralization in the Faina greenstone belt, linking hydrothermal alteration, lithological and structural controls, together with comparisons with similar deposits in Brazil and worldwide.

The investigations mentioned above involved the following analytical techniques: (i) optical petrography (transmitted and reflected light) supplemented by semi-quantitative analyses and imaging using the scanning electron microscope (SEM-EDS), (ii) quantitative mineral chemistry analyses using

the EPMA. Polished thin-section were prepared by Vancouver Petrographics of Vancouver Canada and by Fernando Soares de Souza Preparações Geológicas e Arqueológicas, Minas Gerais, Brazil. All petrography and sample preparation was carried out by the candidate and included optical microscopy, SEM and EPMA analyses. Optical microscopy was conducted in the microscope laboratory of the School of Earth Sciences at the University of Western Australia. The SEM and EPMA analyses were conducted at the Centre for Microscopy, Characterisation and Analysis, at the University of Western Australia with the instrument set up by Dr. M.P. Roberts. Design of analytical work, selection of samples and further interpretation of the collected dataset was conducted entirely by the candidate. All written work was completed by the candidate with discussions and revisions provided by Profs. S.G. Hagemann and Lydia Lobato.

Chapter 2

Hydrothermal Alteration and Mineralization in the Faina greenstone belt: evidence from the Cascavel and Sertão orogenic gold deposits

Abstract The Faina greenstone belt is located in central Brazil and contains the Cascavel and Sertão gold deposits (500,000 oz Au total resource and past production). The bulk of the gold is hosted in quartzite and biotite schist at Cascavel, and carbonaceous schist and banded iron formation at Sertão. The metamorphic grade of the host rocks is greenschist facies.

Cascavel is a vein-hosted gold deposit with orebodies mainly controlled by F2 fold hinges, with minor mineralization in D3 shear zones. Hydrothermal alteration in quartzite is characterized by distal quartz-K-feldspar±muscovite, intermediate quartz-fuchsitic mica-biotite-pyrite, and proximal quartz- white mica-pyrite±chalcopyrite. The biotite schist is characterized by distal quartz-chlorite-calcite and proximal white mica-biotite-tourmaline-siderite-pyrite. Native gold is located within intensely deformed quartz veins, and locally as free gold disseminated in proximal alteration zones.

Sertão is a shear zone-hosted gold deposit with orebodies controlled by D3 shear zones including gold-bearing shear veins and surrounding hydrothermally altered wallrocks. Hydrothermal alteration is characterized by distal quartz-ankerite-chlorite-pyrite±chalcopyrite and proximal quartz-white mica- siderite-pyrite-arsenopyrite±chalcopyrite. Gold is in equilibrium with arsenopyrite and within arsenopyrite disseminated in the proximal alteration zone.

Chlorite geothermometry suggests temperatures of 335-407°C and 328-397°C for the gold- bearing hydrothermal fluids in the Sertão and Cascavel deposits, respectively. Arsenopyrite geothermometry on arsenopyrite within gold-bearing veins in the Sertão deposit yields temperatures between 310-480°C.

Structures controlling hydrothermal alteration and gold mineralization in both deposits are developed during a regional D2 event and subsequently deformed during a D3 event. Hydrothermal alteration minerals are consistently replacing metamorphic minerals. The alteration assemblages in both deposits are similar, albeit their modal mineralogy is controlled by the host rock composition. Therefore, both the Cascavel and Sertão deposits are interpreted as orogenic, mesozonal gold systems comparable with other orogenic gold systems in Brazil and worldwide.

Keywords: Faina greenstone belt, Goiás Archean Block, gold mineralization, Cascavel and Sertão deposits, hydrothermal alteration, central Brazil

1. Introduction Archean and Paleoproterozoic greenstone belts are recognized as fertile terranes for the concentration of gold and base metals (deWit and Ashwal, 1995), and have been the focus of research for decades. In Brazil, gold mineralization in greenstone belts is mostly located in the São Francisco Craton (e.g., Rio das Velhas and Rio Itapicuru), in the Amazonian Craton (e.g., Rio Maria), and in the Goiás Archean Block (e.g., Faina), where the Cascavel and Sertão deposits are located (Jost et al., 2014; Fig. 1, inset C).

Fig. 1. (A) Location of the Tocantins Province in Brazil. (B) Geological map of the Brasília fold belt (modified after Pimentel et al., 2004), and (C) TTG granite-gneiss complexes and greenstone belts of the Goiás Archean Block. Abbreviations correspond to Gr. – Group, seq. – sequence, TTG – tonalite- trondhjemite-granodiorite.

The gold endowment of the Goiás Archean Block is exemplified by greenstone belt-hosted deposits including Crixás (7 Moz Au, owned by Anglo Gold), Pilar (0.7 Moz), and Guarinos (6.5 Moz, both owned by Yamana Gold Inc.). Gold mineralization in greenstone belts of the Goiás Archean Block is typically hosted in metasedimentary and metavolcanic rocks, structurally controlled by brittle-ductile shear zones, including low-angle thrusts and strike-slip faults, and associated with orogenic processes

during the Paleoproterozoic (Pulz, 1995; Queiroz, 2000; Jost et al., 2008; Marques et al., 2013; Jost et al., 2014), with local expressions in the Neoproterozoic (Rodrigues, 2011). In the southern Goiás Archean Block, the Faina greenstone belt (~3 Moz; Silva and Rocha, 2008) forms an elongated, NW-trending synform with, from the bottom to top, metamorphosed ultramafic, mafic and sedimentary rocks, culminating with chemically precipitated sedimentary rocks. The belt is host to the Cascavel and Sertão gold deposits, located about 28 km apart from each other (Fig. 2). Limited investigations have recognized distinct types of gold mineralization styles in the Faina greenstone belt, including (i) vein-hosted gold mineralization in the Cascavel deposit (Jost et al., 2014; Brant et al. 2014), (ii) disseminated gold mineralization in the Sertão deposit (Internal report Troy Resources Ltd, 1998), and (iii) conglomerate-hosted gold (Carvalho et al., 2013). However, these independent studies lacked regional integration to propose a coherent genetic model for gold mineralization in the Faina greenstone belt. The role of the Brasiliano orogeny within the Goiás Archean Block is still controversial and studies in greenstone belts affected by this orogeny, such as the Rio das Velhas, have suggested gold remobilization during this orogenic event (Thorpe et al., 1984; Noce et al., 2007). However, the only manifestation of Neoproterozoic gold mineralization in the Goiás Archean Block (in the Guarinos greenstone belt; Fig. 1, inset C), is interpreted as intrusion-related Au by Rodrigues (2011). In this paper, we present new data on the hydrothermal alteration and mineralization in both the Sertão and Faina deposits. An integrated structural-hydrothermal alteration model for gold mineralization in both deposits is proposed and gold deposits in the Faina greenstone belt compared to other good deposits in Brazil and worldwide.

2. The Goiás Archean Block Five greenstone belts occupy the midwest portion of Goiás state, central Brazil (Inset C in Fig. 1). These greenstone belts are hosted in an allochthonous exotic fragment of Archean crust of uncertain origin and are known as the Goiás Archean Block. The latter Archean crustal fragment was amalgamated during the Brasiliano/Pan-African orogeny onto the western margin of the Brasília fold belt (Pimentel, 2016; Pimentel et al., 2000; Jost et al., 2008). The belt comprises one of three fold belts of the Tocantins Province (Almeida et al., 1981; Marini et al., 1984), formed during the Neoproterozoic convergence between the Amazonian, São Francisco-Congo and Paranapanema cratons (Brito Neves and Cordani, 1991; Fuck et al., 2014). The Goiás Archean Block comprises Archean TTG granite-gneiss complexes (Queiroz et al., 2008; Jost et al., 2008) with inliers of Archean/Paleoproterozoic greenstone belt sequences (Fortes et al., 1995; Tassinari et al., 2006; Santos et al., 2008; Jost et al., 2010), and rare Neoproterozoic intrusions (Rodrigues, 2011; Pimentel et al., 2003). Crustal extension in the Goiás Archean Block is recorded by two main episodes of mafic magmatism: (i) NW-NE trending bodies with hornblende 40Ar/39Ar crystallization ages of 2490 ± 40 Ma (Corrêa da Costa et al., 2006), and (ii) EW-trending

Jurassic/Cretaceous diabase dikes emplaced during a late extensional phase after the West Gondwana break up (Araújo Filho, 2000). The five greenstone belts in the Goiás Archean Block are 40 to 50 km long, and 6 to 15 km wide, and include Faina, Goiás (former Serra de Santa Rita), Crixás, Guarinos, and Pilar de Goiás (Fig. 1, inset C). They share a similar lower stratigraphic sequence, represented by basal komatiites and tholeiitic basalts, which is overlain by distinct sedimentary packages that reflect contrasting depositional settings (Jost and Oliveira, 1991; Resende et al., 1998). The rock sequences underwent greenschist to amphibolite-facies metamorphism (Resende et al., 1999; Jost et al., 2014). The two southernmost greenstone belts, Faina and Goiás, were initially interpreted as a single NW-SE trending belt, approximately 100 km long and 7 km wide, forming a synformal structure (Resende et al., 1998). Teixeira (1981) divided the two greenstone belts based on their different rock associations and their offset by the regional Faina strike-slip fault (Fig. 2). Based on the stratigraphy proposed by Resende et al. (1998), the sedimentary sequence at Faina consists of two cycles of lower conglomerate, overlain by sandstone, pelite, dolomite and iron formation, whereas at Goiás the sequence starts with carbonaceous pelite that grades into chert, iron formation and dolomite, discordantly overlain by turbidite. The contrasting sedimentary records of these two southern belts were interpreted by Resende et al. (1999) as the result of the deposition of two paleobasins under distinct geographical settings and depositional regimes. At Goiás, the sedimentation was considered to be related to a deep marine platform that advanced into a shallower depositional regime, whereas at Faina the sedimentation is interpreted as a platform setting with evidence of two transgressive cycles that culminate with chemically precipitated sediments (Resende et al., 1998). Metamorphic studies conducted by Resende et al. (1999) on the metasedimentary rocks of the Faina greenstone belt document typical greenschist facies assemblages. A Paleoproterozoic (Rhyacian) age is indicated for the sedimentation of all greenstone belts in the Goiás Archean Block (e.g., Fortes et al., 2003; Tassinari et al., 2006; Jost et al., 2008; Brant et al., 2015), and for the volcanic sequences in Guarinos and Pilar de Goiás. In contrast, Archean ages are indicated for the volcanic rocks from Crixás, Faina and Goiás (Jost et al., 2014; Borges et al., 2017). Model Sm-Nd provenance ages of the detrital load from the Faina greenstone belt suggest a maximum age between 3200 and 2800 Ma for protoliths of the sedimentary rocks (Resende et al., 1998; Brant et al., 2015). Carbon isotope analyses of marble from the northern greenstone belts and from the first sedimentary package in the southern belts have positive δ13C values between +10 to +14‰ (Fortes, 1996; Resende et al., 1998; Jost et al., 2008; Santos et al., 2008). The latter values suggest that the sedimentation was completed during the Huronian glaciation, at the beginning of the Rhyacian (Jost et al., 2014). In the Faina greenstone belt, the δ13C values of marble from the second sedimentary cycle (Resende et al., 1999) range between -0.6 and +0.6‰, suggesting a minimum age of deposition during the end of the Rhyacian and start of the Orosirian times (Resende et al., 1999; Jost et al., 2014).

2.1. Exploration and Mining History The Goiás Archean Block has been historically explored for gold. For instance, the Faina greenstone belt is adjacent to the area where gold was first discovered in central Brazil. It displays widespread historical and recent artisanal mining evidence, with records dating back until the 18th century (Silva and Rocha, 2008). For ca. 200 years, approximately 3.3 Moz Au were produced from alluvial sand and gravels in the region (Silva and Rocha, 2008). In the last decades, several mining companies have conducted exploration in the area (METAGO, Western Mining Co., Amazônia Mineração, Troy Resources Brazil, Yamana Gold Inc.). From 2000 to 2007, the Sertão Mineração Ltda., a joint venture between Troy Resources and Amazônia Mineração, exploited 254 Koz Au from oxidised sulfide-rich ore with an average grade of 24.95 Au g/t (DNPM, 2007). In 2016 Orinoco Gold Ltda. was developing an underground mine at the Cascavel deposit and conducted exploration at the Sertão deposit. As of June 2017, the Sertão deposit recorded measured, indicated and inferred resources of 223 t of Au at 6.9 g/t for approximately 50 Koz of contained gold, with mineralization open at depth (Orinoco Gold Ltda., June 2017 ASX report). Faults, veins (±breccias) and hydrothermal alteration associated with anomalous element enrichment (e.g., Ag, Mo, Pb, W, U) comprise a late mineral system known as Tinteiro, which is exposed in several locations in the Faina greenstone belt. Apparently devoid of gold, the metal endowment of the Tinteiro system is presently not constrained (Orinoco Gold annual report, 2014).

3. Sampling and Methods Field work included geological mapping and detailed logging of eight representative diamond drill holes from the Cascavel and Sertão deposits in order to characterize their main lithological, structural and mineralogical variations. Four mineralized samples were collected from underground workings in the Cascavel mine. Despite the emphasis on these two deposits, selected other prospects or small open pits in the Faina greenstone belt were investigated for comparison purposes (e.g., Antena, Xupé, Eliseu targets – see Fig. 2). All structural readings are given in true north coordinates, planes are presented by dip direction/dip (e.g., S1=085°/30°), whereas linear features use the plunge → trend convention (e.g., L2=20°→260°). Measurements in Table 1 show dip and dip direction. Petrography was performed on approximately 100 polished thin sections that included host rocks with various alteration types and intensities from the Cascavel and Sertão deposits. Petrographic studies were conducted using optical microscopy (transmitted and reflected light), supported by semi-quantitative (EDS) analyses and backscattered images (BSE) by scanning electron microscopy – SEM, at the Centre for Microscopy and Microanalyses (CMCA) at the University of Western Australia. The analytical software employed for the SEM was the Tescan Vega. Quantitative microanalyses were performed on the JEOL JXA-8530F microprobe, coupled with five wavelength dispersive spectrometers (WDS), and they were accomplished using proprietary standards. The microprobe data were processed using

software packages, e.g., Probe for EPMA®, CalcImage, Surfer14 and Squid. Microprobe data focused on mineral chemistry of hydrothermal minerals, and their variations displayed as elemental maps (e.g., for arsenopyrite). Operating parameters varied according to the type of mineral phase analyzed. For sulfides and gold, the following set up was used: accelerating voltage of 20-25 kV; working distance of 15 mm; beam current of 40 nA, and a counting time of 50 s on peak position. Elements included: S, Sb, As, Bi, Cu, Au, Fe, Co, Ni, Ag, Pb, Hg, and Zn, whereas for gold Au, Ag, Sb, Fe, Se, Cu, Te, Co, Pb, Bi, Hg, Ni and Pt were analysed. For silicates: accelerating voltage of 15 kV; working distance of 15 mm; beam current of 15 nA, and counting time of ca. 30 s on peak position. Analyses of silicates included the elements F, Cl, Ca, Rb, Si, Al, Mg, Fe, Mn, Na, Ti, K, Ba, Cr, Ni, V, O and H. Mean atomic number background and ZAF corrections were used. The formula of chlorite was calculated on the basis of 18 oxygens, whereas the formula of muscovite and biotite were calculated on the basis of 24 oxygens. All Fe is considered as Fe2+.

4. Geological Setting of the Faina greenstone belt 4.1. Lithostratigraphy The Faina greenstone belt consists of metasedimentary rocks unconformably overlying metavolcanic rocks. The metasedimentary sequence is represented by detrital clastic rocks, with local carbonate lenses, overlain by pelites, followed by carbonaceous schist and banded iron formation (BIF). Based on the original stratigraphic classification by Resende et al. (1998), the rocks described at the Cascavel and Sertão deposits are part of the lower Furna Rica Group (Fazenda Tanque Formation), which overlies metavolcanic rocks from the Serra de Santa Rita Group (Digo-Digo and Manoel Leocádio Formations; Fig. 2). Although dominated by different host rocks, regional stratigraphic correlations suggest that the Sertão deposit is stratigraphically located above the Cascavel deposit. Lower immature siliciclastic and pelitic sedimentary rocks intercalated with minor carbonate lenses form the stratigraphy at Cascavel, whereas dolomitic marble, carbonaceous schist, Fe-rich chert, BIF, metasedimentary rocks and metabasalts dominate the stratigraphy at Sertão. A subvolcanic, ultramafic potassic rock is recorded as undeformed late dikes that crosscut the stratigraphy at the Sertão deposit. These dikes are here referred to as lamprophyre. The stratigraphic units present at the Cascavel and Sertão deposits are described, from bottom to top. Quartzite: clastic sedimentary unit consisting of a thick succession (≤ 150m) of metamorphosed sub-arkose, mica-rich ortho-quartzite with rare lenses of microconglomerate. These quartz-rich immature rocks are enveloped by mica schist (±carbonaceous schist). The white to light grey quartzite typically comprises a medium- to coarse-grained massive rock composed mainly of quartz (55-89 vol. %), with variable amounts of feldspar (8-17 vol. %, mainly microcline with minor plagioclase) and minor muscovite. Seriate, granular aggregates are cemented by recrystallized quartz, filling inter-grain

contacts (Figs. 3A, B). This lithotype can contain thin lenses (≤ 30 m) of marble or impure quartzite. Accessory minerals include rutile, zircon (≤ 200 µm), monazite and rare apatite.

Fig. 2. Geological map of the Faina greenstone belt showing the location of the Cascavel and Sertão gold deposits (modified after Orinoco Gold Ltda., internal report 2017). The rectangles indicate the location of Figures 4 and 5. Structural data collected for the Faina greenstone belt is plotted in an equal area stereonet with lower hemisphere projection.

Metapelite: dark green to light grey, finely laminated and bedded micaceous schists (Figs. 3I and J), mainly composed of quartz (35 vol. %), muscovite (40 vol. %), alkali-feldspar (≤ 13 vol. %), and calcite (9 vol. %). Where present, sulfides include pyrite and chalcopyrite; rutile and hematite are common as minor constituents. Original sedimentary bedding is distinguished by alternating mm- to cm-wide bands of quartz±mica. BIF: narrow (≤ 5 m) intervals of alternating chert and carbonate-rich bands (Figs. 3G, H). It is described as a carbonate facies iron formation displaying typical mesobands (≤ 2.5-4.0 cm width). The carbonate-rich bands are formed by fine-grained, granoblastic Fe-rich dolomite that is intercalated with

recrystallized quartz aggregates with polygonal inter-grain contacts and oscillatory extinction. The BIF is commonly enveloped by narrow (≤ 4m) intervals of rheologically weak carbonaceous schist. The contact between both is intensely deformed in most places, and characterized by comminution of grain- size (≤ 5 µm), oscillatory and textural zoning of quartz. Carbonaceous schist: it is irregularly distributed in narrow (≤ 4 m wide) intervals of fine-grained (< 0.05 mm), and strongly foliated chlorite-quartz schists with carbonaceous matter, commonly associated with dolomitic marble, BIF and Fe-rich chert. A well-developed schistosity is defined by phyllosilicate minerals with very fine-grained black stringers of carbonaceous matter (≤ 10 vol. %) (Figs. 3E, F). These lepidoblastic domains are dominated by chlorite (up to 70 vol. %) and intercalated with seriate, granular aggregates of medium-grained subhedral quartz (≤ 20 vol. %) and Fe-dolomite (≤40% vol.). Minor to trace sulfides include pyrite, chalcopyrite and pyrrhotite, with local sphalerite inclusions in pyrite and rare pentlandite. Other accessory minerals include rutile, monazite, zircon, and apatite. Marble: light grey to pink, massive to laminated dolomite-rich rock that displays dominantly granoblastic texture with local lepidoblastic inliers. It is characterized by polygonal, equigranular aggregates of dolomite (up to 75 vol. %) with typical rhombohedral cleavage. Muscovite laths (≤ 10 vol. %) locally define an incipient foliation. Other minor phases are chlorite, apatite, titanite, rutile and rare carbonaceous matter. Locally, primary sedimentary structures such as graded bedding and/or compositional stratification are preserved. Lamprophyre: black rock consisting of cm-wide phlogopite flakes in an aphanitic matrix of fine- grained phlogopite, pyroxene, olivine, and subhedral to anhedral Fe-Ti oxides. The mineralogy of this rock is compatible with an alnöite.

4.2. Regional Metamorphism Metamorphic assemblages in schists and quartzite are characterized by biotite-chlorite- muscovite-quartz and quartz-alkali-feldspar-muscovite±biotite, respectively. Rare garnet marks tectonic contacts in metapelite, and comprise the assemblage quartz-biotite-chlorite-muscovite-alkali- feldspar-garnet. The sedimentary rocks display a well-developed peak-metamorphic S1 fabric defined by lepidoblastic domains of muscovite, chlorite and biotite. The metamorphic assemblages are indicative of mid-greenschist facies conditions (Miyashiro, 1973; Winkler, 1976).

4.3. Structural Setting The structural evolution of the Faina gold greenstone belt is characterized by a complex and polyphase deformation history. Based on cross-cutting relationships, four major deformation events, i.e., D1, D2, D3 and D4 are recognized (Table 1). Hydrothermal veins are discussed here as part of the deformation events. Those that relate to the gold mineralization at the Cascavel and Sertão gold deposits are characterized in the “Hydrothermal Alteration and Mineralization” section. Identification of primary

sedimentary structures is hindered by intense deformation. Only locally, primary bedding (S0) is preserved as compositional variations in quartzite and the rhythmic alternation of quartz- and carbonate- rich bands in BIF.

The D1 deformation event is characterized by a steeply N-dipping S1 foliation parallel to the original bedding (S0 transposed into S1). This S1 foliation is poorly defined by oriented quartz and muscovite with mm spacing. Additional structures developed include rare NNW-striking, parasitic F1 folds and S1-parallel veins, classified as V1 veins.

TABLE 1. Summary of structural evolution at the Faina greenstone belt

The D2 deformation event is represented by ~E-W-striking, tight to isoclinal F2 folds and subtly

S-dipping axial planar S2 cleavage. In quartzite, S2 cleavage is recorded as mm- to cm-wide spaced cleavage defined by oriented laths of white mica, biotite and fuchsitic mica. In quartz-mica schists, S2 forms a penetrative foliation with mm-wide spacing expressed by oriented white mica, biotite and minor fuchsitic mica (Fig. 3A, B). A L2 intersection lineation between S1 and S2 foliations gently plunges to the west subparallel to F2 fold hinges. Progressive shortening led to the development of local ~E-W- trending, gently S-dipping thrust faults (f2) that caused the stacking of the stratigraphy (Fig. 3C). The preferred orientation of hydrothermal minerals, such as quartz, and fuchsitic mica (±gold) define the gently W-plunging Lm2 mineral stretching lineation (Fig. 3D), which indicates transport to the west.

During the D2 event, deformation of earlier formed V1 veins results in the development of gold-bearing

V2 shear veins formed parallel to the thrust planes. Structures related to D2 are best preserved in the central and northern portions of the Faina greenstone belt.

The D3 deformation event is characterized by NW-striking D3 shear zones (≤ 3m wide), locally defined by LS-mylonites and a moderately N-dipping S3 foliation demarcated by anastomosed white mica, minor biotite and alkali-feldspar with cm-wide spacing. The sub-horizontal orientation of white

mica and quartz plunging north outlines the intersection between S2 and S3 foliations that define the L3 intersection lineation. The orientation of kinematic indicators such as σ-feldspar porphyroclasts in quartzite (Fig. 3E) indicates a reverse sense of displacement. Subtly NNW-plunging Lm3 mineral stretching lineation is indicative of mass transport in the N direction. Local development of subtly S- dipping f3 faults is recorded in the Faina greenstone belt. Structures related to D3 transpose earlier formed structures into parallelism within the NW-trending shear zones (Fig. 3F, G). The reactivation of

~E-W-trending, gently S-dipping V2 veins during D3 is associated with minor hydrothermal fluid flow that led to the formation of paragenetically late galena, pyrite, tourmaline and minor stibnite.

Fig. 3. Photoplate displaying the main structural features observed in the Faina greenstone belt. (A) Photomicrograph showing the orientation of hydrothermal minerals according to S2 (±S3) planes in carbonaceous schist observed in the Cascavel underground stope (cross-polarized light). The S2 foliation is defined by white mica (±biotite), whereas spaced S3 cleavage is defined by shear bands. (B) Fuchsitic mica parallel to S2 cleavage in the Cascavel underground level. (C) The relation between the main fabrics and their cross-cutting relationships exposed in the Cascavel underground stope. At bottom left, thrust fault (f2) offsets the Au-bearing V2 vein. (D) Elongated gold grains (up to cm-wide) oriented parallel to F2 hinges. (E) Sigmoidal alkali-feldspar with dextral sense of displacement recorded in the Cascavel underground stope. (F) Folded V2 vein in the Cascavel underground stope showing the relationship between axial planar S2 and S3 fabrics. At top right, stereonet projection of S1 (in black), S2 (in red), and S3 (in green) fabrics. (G) Weathered mica schist sampled near the Sertão deposit showing the relationship between the S2 and S3 foliations.

The D4 deformation event is characterized by moderately S-dipping f4a thrust fault planes and associated sub-horizontal S-plunging slickenlines. Additional structures produced during D4 include steeply W-dipping f4b reverse faults and associated steeply S-plunging slickenlines. Common V4 fault- fill veins and local breccias are commonly attributed to D4 structures in the Faina greenstone belt. The relative timing between D4 faults is unclear due to mutual crosscutting relationships. Structures related

to this late deformation event are regionally distributed and crosscut all previous structural elements and associated veins.

TABLE 2. Veins described at Cascavel and Sertão deposits. Veins related to the gold-related D2 hydrothermal alteration event are described as V2d, V2i and V2p, whereas gold-bearing V2 veins are discriminated based on the host rock

5. Hydrothermal Alteration and Mineralization at the Cascavel Deposit

Gold mineralization at the Cascavel deposit (Fig. 4A-B) is associated with V2 quartz veins typically developed in quartzite and, locally, in biotite schist near or within sheared lithological contacts. Irregular alteration zones are defined by distinct hydrothermal alteration assemblages around these veins in each of the host rocks (Fig. 5A). The width of the alteration zones is mostly dependent on the permeability of wall rocks and proximity to veins and faults zones. In order to better understand the relationship between deformation and hydrothermal alteration in the Cascavel deposit, a detailed study of hydrothermally altered wallrock and associated veins was conducted (Table 2). Veins formed during D2 are defined with respect to the alteration zones and classified as distal (V2d), intermediate

(V2i), and proximal (V2p).

5.1. Quartzite Hydrothermal alteration of quartzite encompasses distal, intermediate and proximal alteration zones that are defined by mineral assemblages as indicated in Fig. 5A.

Fig. 4. Geological map of the Cascavel deposit area (A) shows the main lithologies, structures with mapped and projected quartz veins that define the orebody, as well as the area of old workings. (B) Cross-section AB (modified from Orinoco Gold Ltda., internal report 2017).

Fig. 5. Hydrothermal alteration diagram (A), hand samples (B, D, F, H, L) and photomicrographs (crossed polarizers: C, E, G, I; parallel polarizers: M, N; reflected light: J) in the Cascavel deposit. Photos B to G represent the hydrothermal alteration zones in quartzite, whereas photos H to N represent samples from the biotite schist. (A) Distribution of hydrothermal minerals in distal, intermediate and proximal alteration zones in quartzite and biotite schist. (B) Distal quartz-muscovite alteration zone in quartzite with sparse metamorphic feldspar porphyroclasts. (C) Coarse (30-500 µm) quartz feldspathic host rock, enlarged from figure 5B. (D) Fuchsitic mica bands typical in quartzite within intermediate alteration zones. (E) V2i vein enveloped by fuchsitic mica in intermediate alteration zone. (F) Dark coloured fuchsite-rich bands typical of quartzite in proximal alteration zone. (G) Intensely deformed quartz-white mice-K-feldspar in proximal alteration zones of quartzite. (H) Distal alteration zone in biotite schist displays quartz veinlets surrounded by green chlorite-rich envelopes and sparse quartz- carbonate veinlets parallel to regional S2 foliation. (I) Enlargement of inset I in Figure 5H showing

distal quartz-biotite-chlorite±calcite alteration assemblage. (J) Replacement of xenoblastic Py1 by pyrrhotite, which is also being overprinted by chalcopyrite. (K) V2d vein enveloped by hydrothermal biotite in distal alteration zones of biotite schist. (L) Proximal alteration zone in biotite schist with discreet ochre bands reflecting weathering of hydrothermal minerals as ankerite±siderite and pyrite. (M) Enlargement of inset M in Figure 5L shows proximal alteration zone of biotite schist. Fine-grained metamorphic biotite is replaced by white mica and/or reddish flakes of coarse-grained hydrothermal biotite that is locally replaced by retrograde metamorphic chlorite. (N) Hydrothermal tourmaline and biotite in proximal alteration zone of biotite schist. Abbreviations: Qtz: quartz; Ms: muscovite; Fm: fuchsitic mica; Bt: biotite; WM: white mica; Py: pyrite; Au: gold.

5.1.1. V2 veins in quartzite

The V2 veins hosted in quartzite comprise up to 1 m wide massive to laminated veins with translucent to milky quartz. The veins are locally hosted within shear zones defined by LS-type mylonites. Commonly folded, the veins display structures as boudins and pinch and swell structures.

Narrow slivers of strongly foliated and hydrothermally altered wallrock are often observed within V2 veins. They are composed of recrystallized quartz, minor alkali feldspar, white mica and traces of chalcopyrite (-pyrite) and free gold. Gold forms round- to rod-shaped grains, up to a few cm in size, oriented parallel to the F2 fold hinges and L2 intersection lineation. Note that the V2 veins are loci to the main gold mineralization at Cascavel and are typically higher grade (~ 4.0 Au g/t) compared to the V2 veins hosted in biotite schist (≤ 2 Au g/t; Orinoco Gold Ltd., 2014).

5.1.2. Wallrock alteration in quartzite The distal alteration zone in quartzite is defined by the assemblage quartz-alkali-feldspar- muscovite and it extends outwards 5-30 m from the V2 veins (Fig. 5B, C). The transition from least altered to distal alteration zone is poorly defined by diffuse/gradual contacts, marked by the increase of quartz veinlets. Quartz (88 vol. %) presents typical oscillatory extinction. Feldspar (10 vol. %) includes mostly alkali feldspar porphyroclasts that are commonly saussuritized with discrete development of muscovite and quartz along its irregular borders with minor albite as finer grains. Lath-like muscovite comprises sparse (≤ 2 vol. %), colorless grains. Pyrite is present in trace amounts.

The intermediate quartz-alkali-feldspar-fuchsitic mica-biotite alteration zone ranges between 5 cm to 15 m in width. Transition from the distal alteration is marked by sharp contacts characterized by the appearance of fuchsitic mica (Fig. 5D, E), a decrease of the modal percentage of quartz (70 vol. %), and by an increase in alkali feldspar (15 vol. %). Quartz displays dynamic recrystallization features and oscillatory extinction. Alkali feldspar forms subhedral to anhedral phenocrysts (≤ 0.6 cm), and commonly forms untwined and saussuritized micro phenocrysts. The crystals are mostly untwined (< 100 µm), but they can also be faintly (400 µm) and locally twinned. Albite occurs as anhedral grains that form fine-grained, polygonal aggregates with quartz. Pervasive fuchsitic mica (11 vol. %) is commonly in equilibrium with biotite and white mica. Minor phases include biotite, pyrite, chalcopyrite

and hematite (after pyrite), with the latter being responsible for fine, dark brown inliers and reddish dusting of feldspar grains. Trace accessory minerals include monazite, rutile, and zircon. The proximal alteration zone (quartz-white mica-alkali-feldspar-pyrite±biotite) is characterized by 5 to 50 cm irregular halos developed around V2 veins. Polygonal quartz (63 vol. %) shows strong undulatory extinction and conspicuous elongation. The alkali feldspar (≤ 13 vol. %) tends to form sigmoidal shapes as the deformation intensity increases and is commonly surrounded by pressure shadows filled with white mica and quartz. White mica (14 vol. %) is more pervasive next to the V2 veins. Comminution of grain size (physical effect of intra-crystalline deformation), and plastic recrystallization (sub-grain rotation, bulging) are diagnostic features observed proximal to the D3 shear zone (Fig. 5F, G). Fine-grained, free gold, typically less than 0.3 mm in size, is disseminated throughout proximal wall rock alteration zone in quartz ±alkali feldspar veins.

TABLE 3. Representative chemical reactions involved in the hydrothermal alteration observed in the Cascavel and Sertão deposits

5.2. Biotite schist Hydrothermal alteration of biotite schist encompasses distal and proximal alteration zones that are defined by distinctive mineral assemblages (Fig. 5A).

5.2.1. V2 veins in biotite schist

The V2 veins comprise narrow (≤ 5 cm) quartz veinlets typically hosted in biotite schist. They are composed of recrystallized quartz (up to 95 vol. %), with variable amounts of minor white mica, biotite (±alkali feldspar), pyrite, chalcopyrite, with traces of sphalerite, galena, pentlandite and free gold. Gold is irregularly distributed as anhedral to rounded grains (≤ 30 µm) filling inter-grain space in granular quartz aggregates. Quartz veinlets are characterized by up to 2 g/t Au (2017 internal report Orinoco Gold Ltda.).

5.2.2. Wallrock alteration in biotite schist

The distal alteration zone (quartz-chlorite-calcite) ranges 2-15 m outwards from the V2 veins. Transition from least altered to the distal zone is subtle and characterized by mm- to cm-scale quartz- calcite (±albite) veins, in more intensely deformed wallrock with discrete ochre staining (Fig. 5H). Fine- grained metamorphic biotite is progressively replaced by white mica (30 vol. %), ragged hydrothermal biotite (9 vol. %), and chlorite (14 vol. %). The latter is commonly intergrown with white mica and/or biotite as fine-grained laths disseminated in the wallrock, within or surrounding veins (Fig. 5I). Additional minerals include quartz (13 vol. %), alkali feldspar (≤ 7 vol. %), minor calcite (≤ 4 vol. %), and rare titanite. Sulfides include pyrrhotite, pyrite, chalcopyrite, trace pentlandite and sphalerite. Pyrite occurs as (i) deformed, fine-grained Py1 intergrown with hydrothermal micas, and (ii) subhedral, medium to coarse-grained Py2 (30-400 µm) with local inclusions of galena (≤ 30 µm). Xenomorphic Py1 is overgrown by chalcopyrite, pyrrhotite and/or Py2 (Fig. 5J). Trace apatite (Fig. 5K), monazite, rutile, galena and rare cobaltite are in equilibrium with quartz-chlorite-calcite in the distal alteration zone in biotite schist. The proximal alteration zone (white mica-biotite-quartz-chlorite-ankerite) extends up to 1 m from the V2 veins. The assemblage is composed of ubiquitous white mica (42 vol. %; (Fig. 7L-N), biotite (15 vol. %), alkali-feldspar (10 vol. %), quartz (10 vol. %), chlorite (7 vol. %) and variable amounts of ankerite (Fig. 5L). From distal to proximal alteration zones, calcite is replaced by ankerite (Eq. 4; Table 3), and minor siderite. The replacement of calcite by ankerite can be associated with apatite, which increases in modal abundance towards the proximal zone. Other minor to trace minerals include epidote and titanite, the latter gradually replaced by rutile (Eq. 7; Table 3). Brown, fine-grained metamorphic biotite is replaced by white-mica and, to a minor extent, by medium to coarse-grained red flakes of hydrothermal biotite (Fig. 5M), with the latter being typically overprinted by retrograde chlorite. Titanite is replaced towards the proximal alteration zone by ilmenite and rutile (Eq. 7; Table 3). Paragenetically late tourmaline is common as prismatic or semi-round basal sections in the proximal zone (Fig. 5N). Trace monazite (≤ 450 µm), zircon (≤ 50 µm) and xenotime (≤ 80 µm) are common in the proximal alteration zone.

5.3. Cascavel Mineral Chemistry Quantitative analyses of hydrothermal minerals, including K-feldspar, biotite, white mica, calcite, epidote (Table A.1), and chlorite (Table A.2) were conducted on samples representative of the hydrothermal alteration zones at the Cascavel deposit. Additional analyses were completed on sulfide mineral phases e.g. pyrite, chalcopyrite, arsenopyrite, galena, stibnite and gold (Table A.3).

5.3.1. Chlorite

Microprobe analyses on chlorite were conducted on representative samples from distal, intermediate and proximal alteration zones in biotite schist. A total of 120 chlorite grains from five samples of hydrothermally altered biotite schist were analysed (Table A.2). Major chemical variations are represented by the contents of Si (2.47-2.80 apfu), Aliv (1.14-1.52 apfu), Alvi (1.19-1.50 apfu), Mg (1.18-2.3 apfu), Fe (2.17-3.31 apfu), and Fe/(Fe+Mg) (0.48-0.73). Chlorite classification based on Foster (1962) shows that most of the hydrothermal grains plot within the ripidolite field. A few grains from least altered garnet-bearing schist and V2i vein in biotite schist plot in the brunsvigite field (Fig. 6A). Major chemical changes in chlorite are displayed by an increase in the total amount of Fe of disseminated grains in intermediate alteration zone, with higher values reported in chlorite from V2p veins in the proximal alteration zone. The Fe/(Fe+Mg) ratios range from 0.48 to 0.73, with lower values for metamorphic chlorite associated with the least altered samples. The FeOT content of disseminated chlorite in proximal alteration zone is lower (26.98 wt. %) than that in vein-hosted chlorite (32.01-34.87 wt. %) in proximal alteration zone. The precipitation of iron-rich minerals such as ankerite/siderite and sulfides may have hampered the incorporation of iron in disseminated chlorite from the proximal alteration zone.

5.3.1.1. Chlorite geothermometry: Temperature constraints based on chlorite chemistry utilizes the empirical thermometer proposed by Kranidiotis and MacLean (1987) using the formula T (°C) = 106 * (Aliv + 0.7 * Fe/ (Fe+Mg) + 18). At the Cascavel deposit, chlorite from least altered biotite schist (n=58), interpreted as metamorphic, returns temperatures in the range 335-380°C, with an average of

362±10 °C. Chlorite hosted in intermediate V2i vein (n=37) yields temperatures between 360-393°C, averaging 385±6°C. Similar temperatures derive from chlorite disseminated in intermediate alteration zone (n=10), with values between 397° and 413°C, averaging 406±4°C. Chlorite associated with V2p proximal veins (n=7) yields temperatures between 391° and 428°C, with an average temperature of 417±11°C. Disseminated chlorite in proximal zones yields similar temperatures, ranging from 396° to 421°C, with an average of 407±6°C (Fig. 6B).

5.3.2. Biotite Spot analyses (n=197) of disseminated (n=3) and vein-hosted (n=3) biotite are presented in Table A.1 and in Fig. 6C-D. The main compositional variations observed in the chemistry of biotite include: Mg (0.4-2.6 apfu), Ti (0.0-0.4 apfu), Fe (0.5-4.0 apfu), and Fe/(Fe+Mg) ratio (0.46-0.76). The increase in the Ti content is associated with the temperature of equilibration with Fe-Ti phases, e.g., rutile, which slightly increases towards proximal alteration zone (Table A.1). The Mg content is higher in distal samples, with Mg/(Mg+Fe) ratio average of 0.03 apfu (Fig. 6C). Vein-hosted and disseminated biotite in intermediate and proximal zones are considered slightly Al-rich, with overall total Al content of 3.13 apfu (average). Biotite hosted in V2d and V2i veins include annite, whereas disseminated biotite in the

proximal zone is mostly siderophyllite (Fig. 6C). The decrease of Fe observed in hydrothermal biotite preferentially hosted in veins and in proximal alteration zone can be explained by the co-existence with pyrrhotite. According to Marmo (1957), biotite co-existing with pyrrhotite is more Mg-rich than in sulfide-free pelites due to reaction in which Fe from biotite reacts with sulfur to form pyrrhotite.

Fig. 6. Mineral chemistry diagrams from hydrothermal alteration minerals observed at the Cascavel deposit. (A) Classification diagram (Foster, 1962) applied for chlorite in metasedimentary rocks plots most of the hydrothermal chlorite in the ripidolite field, except for some grains hosted in the least altered rocks and V2i veins that are located in distal alteration zones. (B) Fe/(Fe+Mg) vs temperature diagram of chlorite shows a positive correlation from distal to proximal alteration zone. Note that vein-hosted chlorite has higher Fe/(Fe+Mg) ratios than disseminated grains, followed by a fairly increase in their Fe content from distal to proximal alteration zones. (C) Higher Mg contents are related to hydrothermal biotite hosted in biotite schist (annite) when compared to the more Al-rich biotite hosted in quartzite (siderophyllite). (D) Variation of Si versus Cr in white mica hosted in biotite schist and quartzite shows a significant increase of Cr in the proximal alteration zone in quartzite. Abbreviations - BS: biotite schist and QTO: quartzite.

5.3.3. White mica The chemical composition of white mica hosted in quartzite (n=3) and biotite schist (n=3) from the Cascavel deposit was measured by EPMA (Table A.1, Fig. 6D). Quantitative analyses indicate

muscovite and phengite (lower Al and higher Si) compositions. The total Al content appears to decrease from distal to proximal alteration zone, followed by an increase of Si. The extent of phengite substitution is constrained by the sum of siderophile elements (Mg, Fe2+, Cr, Mn, Ti, Li) that potentially replace Alvi (values displayed in apfu at Table A.1). The more phengitic samples (≥ 0.82) are associated with proximal alteration assemblages. The Fe content is higher in white mica hosted in biotite schist compared with those hosted in quartzite, which likely reflects the host rock chemistry. The Fe tends to increase from distal to proximal alteration zones in both host rocks. Chromium commonly fills the octahedral site in white micas, forming green micas pervasive in the proximal and intermediate alteration zones. Samples with high Cr (± Fe) contents (≥ 0.1 wt.% Cr2O3) are termed fuchsitic micas and include samples hosted in quartzite of the intermediate and proximal alteration zones (Fig. 6D).

Fig. 7. Photomicrographs (reflected light; A, C, E, F) and BSE images (B, D) of Cascavel gold grains (A) Free gold (~ 500µm) with hematite-rich corroded borders next to weathered galena. (B) Enlargement of inset B in Figure 7A with element distribution evidence the replacement of galena by chalcopyrite (± hematite) and gold at the bottom left. Local formation of covellite is considered as supergene. (C) Subhedral arsenopyrite within fracture next to anhedral galena overprinted by hematite/covellite. (D) Enlargement of inset D in Figure 7C shows covellite replacing galena and

botryoidal to radial hematite filling fractures. (E) Embayed contours of free gold (300µm) suggest intergrain precipitation of these grains. (F) Galena being replaced by chalcopyrite (-iron oxides) adjacent to free gold. Textural relationships, e.g. lack of equilibrium assemblages, indicate galena and gold are not coeval. Note trails of fluid inclusions in quartz grains on top of galena grain.

5.3.4. Gold-silver ratio Gold occurs as anhedral to bleb-like ‘free’ grains (0.2-3 cm wide) commonly observed in close spatial association, but not in equilibrium, with galena (Fig. 7A-F). Twenty-six native gold grains (≤ 2 wt. Ag) from two samples of V2 vein hosted in quartzite were analysed (212 spot analyses in total). The Au content varies between 97.9-101.4 wt. %, averaging 100.3 ± 0.7 wt. %, whereas silver varies from 0.7-1.3 wt. %, averaging 0.7 ± 0.1 wt. % (Table A.3). The gold fineness, given by the ratio Au/ (Au+Ag) * 1000 wt. %, varies between 986 and 993, averaging 991 ± 0.9 (Fisher, 1945).

6. Hydrothermal Alteration and Mineralization at the Sertão deposit Gold mineralization at the Sertão deposit (Fig. 8A, B) is located mostly in proximal alteration zones and in V2 quartz-siderite-white mica shear veins. The shear veins are typically surrounded by irregular, 0.1-4 m wide zones of intensely deformed and hydrothermally altered wall rock, comprising carbonaceous schist and BIF. The distribution and width of alteration zones are controlled by the intensity of deformation, distance to the vein and the host rock composition. The hydrothermal alteration in both host rocks is characterized by distal and proximal alteration zones that display similar modal mineralogy (Fig. 9). The distal alteration zone is dominated by metamorphic assemblages, with rare relicts of original sedimentary textures preserved. Gold-bearing veins change their shape, mineralogy and texture according to the host rock in which they form. Veins formed during D2 are defined based on their relative distance to the mineralized zone or orebody, and are thus classified as distal (V2d), intermediate (V2i), and proximal (V2p).

6.1. Carbonaceous schist Hydrothermal alteration of carbonaceous schist encompasses distal and proximal alteration zones with discreet veins (Fig. 9B-E). Transition from distal to the proximal zone is characterized by subtle color change, massive sulfide-rich wallrock alteration and intense deformation.

Fig. 8. Geological map (A) and AB cross-section (B) of the Sertão deposit area with location of the open pit mined by Troy Resources Ltd. during the early 2000s and the projection of the ore shoot (in red) (modified from Orinoco Gold Ltda., internal report 2015).

Fig. 9. (A) Diagram showing the distribution (in volume %) of hydrothermal minerals in carbonaceous schist (B to E) and BIF (F to J) at the Sertão deposit. (B-E) Photomicrograph and hand samples of hydrothermal zoning characterized by an increase in sulfide content (pyrite, arsenopyrite, ±chalcopyrite) from distal (B, C) to proximal (D, E) alteration zones in carbonaceous schist. The main changes observed in the transition from distal to proximal alteration zones involves the replacement of chlorite by ankerite and white mica. (F and G) Hand sample and photomicrograph of BIF in distal alteration zones shows regular intercalation of mesobands composed of dark Fe-dolomite/ankerite and white chert. (H) Hand sample of transition between carbonaceous schist and BIF. Note the reddish stain from weathering of siderite±ankerite typical for the proximal alteration zone. (I and J) Photomicrograph displays white mica-pyrite intergrowth adjacent to V2 veins of BIF in the proximal alteration zone and intensely deformed hand sample of BIF showing isoclinal F2 folds, respectively. Abbreviations: Qtz: quartz, Chl: chlorite, Wm: white mica, Py: pyrite.

6.1.1. V2 veins in carbonaceous schist

The V2 shear veins (quartz-siderite-white mica-pyrite-arsenopyrite-chalcopyrite) are narrow (0.5- 1.5 cm) and preferentially located near lithological contacts. Texturally, they comprise an interlobate aggregate of recrystallized quartz (40 vol. %), subhedral to anhedral Mg-siderite (35 vol. %), lepidoblastic domains of white mica (15 vol. %), and anhedral to subhedral pyrite (≤ 7 vol. %). Pyrite is locally observed in equilibrium with stibnite (≤ 3 vol. %) and/or galena. Intergrain contacts and textural evidence suggest the latter are syn to late with respect to the main Py-Asp-Cpy±Po assemblage.

Fig. 10. Photomicrographs of proximal alteration samples from the Sertão deposit. (A) BSE image illustrating typical sulfide-rich white mica-quartz-siderite-pyrite-arsenopyrite assemblage. Note the mutual crosscutting relationship between arsenopyrite and pyrite. (B) Xenomorphic Py1 is overprinted by coarse Py2. Precipitation of the latter after ankerite/siderite developed typical poikiloblastic texture. (C) Well-developed arsenopyrite with local pyrrhotite inclusions is in equilibrium with chalcopyrite. (D) Idiomorphic arsenopyrite with gold inclusion (ca. 10 µm) overprints Py2 in quartz-siderite intergrowth. (E) Gold (≤ 50 µm) in equilibrium with arsenopyrite. (F) Free gold (ca. 20 µm) adjacent to subhedral, arsenopyrite grain in equilibrium with pyrite and chalcopyrite.

6.1.2. Wallrock alteration in carbonaceous schist The distal alteration zone (quartz-chlorite-ankerite-pyrite) is narrow and extends 0.1-4 m away from the V2 veins with poorly defined contacts with the least altered rocks. It consists of mm-wide lepidoblastic domains (Fig. 9B, C) comprising chlorite (20 vol. %) with scattered white mica laths (≤ 5 vol. %) intercalated with granoblastic domains of polygonal quartz (50 vol. %) intergrown with subhedral to anhedral ankerite±Fe-dolomite grains (≤ 25 vol. %). Pyrite (3 vol. %) occurs as subhedral grains. Minor anhedral chalcopyrite (≤50 µm) in equilibrium with pyrrhotite fills cleavage planes in well-developed V2d ankerite (±Fe-dolomite) veins. Trace pentlandite is in equilibrium with pyrrhotite. The proximal alteration zone (quartz-white mica-siderite-pyrite-arsenopyrite±chalcopyrite) extends 0.5-2 m outwards from the V2 shear veins. It consists of white mica (40 vol. %), Mg-rich siderite (≤ 35 vol. %), pyrite (5 vol. %), arsenopyrite (4 vol. %), minor chalcopyrite (≤ 1 vol. %), trace pyrrhotite,

pentlandite and sphalerite (Fig. 9D, E). Other typical minerals include tourmaline and rod-shaped ilmenite, mostly enveloped by rutile rims. Carbonaceous schist typically adjacent to BIF with disseminated arsenopyrite represents the main ore at the Sertão deposit. Pyrite is in equilibrium with arsenopyrite and chalcopyrite. Two types of pyrite are observed: a fine-grained (≤ 100µm) deformed Py1, and a coarse (≤ 500µm) subhedral Py2. Arsenopyrite (≤ 400µm) comprises isolated euhedral to subhedral grains that overprint earlier formed sulfides such as pyrite, chalcopyrite, and minor pyrrhotite. Gold is in equilibrium with arsenopyrite (refractory gold), and rarely as ‘free’ gold grains (≤ 20 µm) between white mica laths or filling intergrain spaces of quartz aggregates. Gold in equilibrium with arsenopyrite forms inclusions, anhedral grains, and rarely fills cracks in arsenopyrite. The highest gold grades reported for the Sertão deposit (13.55 g/t Au in STO 5: 130.1-132m; Orinoco, annual report 2016) are related to disseminate sulfides hosted in proximal alteration zone of carbonaceous schist.

6.2. BIF Hydrothermal alteration of BIF encompasses distal and proximal alteration zones that grade towards sulfide-rich, intensely deformed bands (Fig. 9F-J). Compared with carbonaceous schist, the modal mineralogy of hydrothermal alteration in BIF is characterized by higher contents of quartz and, to a minor extent, chalcopyrite.

6.2.1. V2 veins in BIF

The V2 veins (quartz-siderite-white mica-pyrite-arsenopyrite-chalcopyrite-gold) are narrow (≤ 30 cm), massive, foliation (sub-) parallel, and typically pinch and swell along strike. They are hosted by intensely deformed BIF with thin (≤ 10 cm) bands of alternating carbonaceous schist. The V2 veins are composed by white, interlobate quartz (47 vol. %), siderite (30 vol. %), white mica (15 vol. %), pyrite (5 vol. %) and arsenopyrite (3 vol. %). The limits between hydrothermally altered wallrock and

V2 veins commonly display comminution of quartz grains towards the vein core and slivers of white mica-siderite-pyrite.

6.2.2. Wallrock alteration in BIF

The distal alteration zone (quartz-ankerite±chlorite) in BIF extends 4-30 m outward from V2 veins and displays poorly defined contacts with least altered samples (quartz-Fe-dolomite; Fig. 9F, G). The transition between alteration zones is marked by discreet textural and grain size change (Fig. 9H). The distal zone is characterized by granular aggregates of ankerite (±Fe-dolomite) (≤ 45 vol. %) intercalated with quartz (43 vol. %), minor white mica (7 vol. %) and chlorite (≤ 5 vol. %). Pyrite and chalcopyrite located within cleavage planes and/or as poikiloblastic blebs of ankerite/siderite suggest

the replacement by the former. Traces of arsenopyrite, pyrrhotite, pentlandite and sphalerite are disseminated within the wallrock. The proximal alteration zone (quartz-siderite-white mica-pyrite-arsenopyrite-chalcopyrite

±pyrrhotite) extends approximately 0.1-2.0 m outward from the V2 veins (Fig. 9I). It is characterized by high volume % of disseminated Py-Cpy-Asp-Po in shear zones and hinges of tight isoclinal folds (Fig. 9J). Pyrite, the dominant sulfide (7 vol. %), is observed as fine-grained Py1 or as subhedral Py2 (≤ 500 µm). Xenoblastic Py1 is typically intergrown with white mica and locally overprints chalcopyrite (Fig. 10A). Py2 forms a poikiloblastic texture with quartz and siderite (±ankerite; Fig. 10B). Minor to trace chalcopyrite is in equilibrium with arsenopyrite and, locally, with pyrite±pyrrhotite. Arsenopyrite (4 vol. %) is described as (i) individual idiomorphic, foliation (sub-) parallel grains or (ii) epitaxial crystals overprinting pyrite-chalcopyrite intergrowths. Locally, pyrrhotite inclusions (≤ 20 µm) occur in well-developed arsenopyrite (Fig. 10C). Gold is observed with arsenopyrite as inclusions (≤ 50 µm) and/or anhedral grains (Fig. 10D-F).

6.3. Sertão Mineral Chemistry Quantitative microanalyses of hydrothermal minerals, including silicates (chlorite, white mica), carbonates (calcite, ankerite, siderite), sulfides (pyrite, chalcopyrite, arsenopyrite, galena, stibnite) and gold were conducted on samples representative of the hydrothermal alteration zones at the Sertão deposit (Tables A.4, A.5 and A.6).

6.3.1. Chlorite

Chlorite analyses (n=113) included samples of V2d veins hosted in marble and carbonaceous schist, V2p veins hosted in metapelite and, for comparison purpose, V4 veins hosted in BIF (Table A.4.). Major chemical variation are displayed by Fe/ (Fe+Mg) ratios (0.24-0.68), Si (2.54 to 2.88 apfu), Aliv (1.11 – 1.45 apfu), Alvi (0.97-1.46 apfu), and Mg (1.54-3.46 apfu). According to the classification scheme proposed by Foster (1962), most hydrothermal chlorite is ripidolite, few distal samples fall in the brunsvigite and sheridanite fields, whereas chlorite from V4 veins plots entirely in the brunsvigite field (Fig. 11A). Chlorite hosted in V2d veins presents lower Fe/(Fe+Mg) ratios (≤ 0.4) compared with those hosted in V2p veins (~ 0.6), with higher ratios for late V4 chlorite (Fig. 11B). Thus, chlorite shifts from Mg-rich in distal alteration zones to more Fe-rich composition in proximal zones including the

V2p veins.

6.3.1.1. Chlorite geothermometry: Chlorite bordering V2d veins hosted in marble yield temperatures between 328-346°C, averaging 337±4°C (Fig. 11B). Calculated temperatures from chlorite bordering veins in distal carbonaceous schist (sample 2:98) returned slightly higher temperatures between 362-386°C, averaging 377±5°C. Chlorite hosted in veins from proximal alteration zones in metapelite (sample 5:69) range between 361-397°C, with a median temperature of

385±8°C. Late chlorite from V4 veins (sample 2:120) in BIF displays temperatures between 344-380°C (n=45), with an average of 363±8°C.

Fig. 11. Mineral chemistry diagrams showing hydrothermal alteration minerals observed in the Sertão deposit. (A) Classification scheme (Foster, 1962) plots most of the analyses within the ripidolite field, with few samples representative of least altered rocks and distal alteration zones that fall in the brunsvigite field. (B) Temperature versus Fe/(Fe+Mg) ratio of chlorite grains indicates a correlation between host rock and Fe content. Note samples are illustrated according to the same colour schemes as in Fig. 11A. Widespread compositions and temperature ranges (340-400°C) are attributed to chlorite hosted in V4 veins. (C) Ternary diagram of carbonate composition displays calcite in distal alteration samples, Fe-dolomite in intermediate alteration samples, siderite in proximal alteration samples and ankerite (-Fe-dolomite) in late mafic intrusion. (D) Electronic images outline compositional zoning of carbonate minerals in distal (-intermediate) alteration zones and in (E) vein-hosted mineralized BIF. (F) Sulfur activity vs temperature applied to arsenopyrite geothermometry (modified after Kretschmar and Scott, 1976). Thin lines represent arsenic isopleths (atomic %). Dashed areas indicate buffer assemblage relative to at. % As values of Asp II and Asp III. Abbreviations: Asp: arsenopyrite, Py: pyrite, Po: pyrrhotite, Lö: loëllingite, Stb: stibnite; Po: pyrrhotite; Bor: bornite; Ccp: chalcopyrite.

6.3.2. White mica Analyses of white mica (n=42) hosted in carbonaceous schist, metapelite and marble are presented in Table A.5. Major changes are represented by a decrease in the contents of Mg and Fe towards the proximal alteration zone, probably due to the precipitation of ankerite and Mg-siderite. Silicium and Aliv in white mica tend to increase from distal to proximal alteration zone. Greater

phengitic component (≥ 0.5) is observed in the proximal carbonaceous schist and marble, whereas white mica hosted in BIF shows a more Na-rich (paragonite) composition. Green Cr-rich white mica (≥ 0.4

Cr2O3 wt. %) is ubiquitous in the transition between distal and proximal alteration zones. Lower Fe content in white mica from the proximal alteration zone could be related to the co-precipitation of Fe rich sulfides such as pyrite and pyrrhotite.

6.3.3. Carbonates Ankerite and siderite phenocrysts significantly increase towards proximal alteration zones (Fig. 11C-E). Analysed samples (n=5) include carbonate hosted in distal alteration zones in marble and metapelite and BIF-hosted V2 vein. Carbonate minerals include Fe-dolomite, ankerite, (Mg-) siderite and calcite. In distal alteration zones in marble, disseminated, fine-grained Fe-rich dolomite (≤ 0.5 mm) grades to vein-hosted, coarse (0.5-1.2 mm) Fe-Mg carbonate with compositions spanning breunnerite, pistomesite, sideroplesite (Fig. 11C). This is expressed by an outward zonation from Fe-rich cores to

Mg-rich rims (Fig. 11D). In V2 veins hosted in mineralized BIF, carbonate composition varies between siderite and ankerite (Fig. 11E), whereas in V2p veins calcite (≤ 0.5 mm) is typically in equilibrium with quartz-galena-pyrite. Therefore, major oxide composition from analysed carbonates shows enrichment of FeO (0.86-53.68 wt. %) and MgO (0.03-21.63 wt. %), whereas CaO (0.23-54.83 wt. %) is depleted as siderite is precipitated (except for retrograde calcite in equilibrium with galena and pyrite). Overall low contribution of MnO (0.33-1.12 wt. %) is registered, with higher contents in the proximal alteration zone. The Fe/(Fe+Mg) ratio is low (0.13) in disseminated Fe-rich dolomite from distal alteration zone hosted in marble and lamprophyre (0.21), with high values (0.45-0.87) reflected by hydrothermal ankerite±Mg-siderite from V2d or V2 veins hosted in metapelite and BIF, respectively (Table A.5).

6.3.4. Arsenopyrite Representative samples from mineralized BIF (n=2) and carbonaceous schist (n=1) were selected for EPMA microanalysis of arsenopyrite. Temperature constraints, considering the arsenopyrite- pyrite±pyrrhotite buffer and the arsenopyrite chemistry, were conducted based on thermodynamic work by Kretschmar and Scott (1976), and estimation of sulfur fugacity was conducted using the work from Barton and Skinner (1967; Fig. 11F). Arsenopyrite composition (146 analysis, 35 grains) comprises a limited range in the As-S-Fe system, with similar chemistry of grains hosted in both carbonaceous schist and BIF (Table A.6). Major elemental trends, from core to rim, reflect enrichment of As (28-33 at. %), decrease in S (33-39 at. %) and subtle changes in Fe (31-33 at. %). Prominent chemical zoning, typically displayed by As-rich rims and S-rich cores, is commonly observed in large (≥ 500 µm) arsenopyrite grains (Fig. 12A-E).

Fig. 12. Microprobe elemental maps show changes in major (As, S, Fe) and trace (Ni, Sb, Co) elements in arsenopyrite. A and B refer to arsenopyrite hosted in two samples of carbonaceous schist, whereas C, D and E refer to arsenopyrite hosted within the same sample of BIF. Note that arsenopyrite hosted in BIF appears to show anomalous Sb(-Ni) cores and grains hosted in the carbonaceous schist develop Co(-Ni)-rich rims.

6.3.4.1. Arsenopyrite geothermometry: The arsenopyrite geothermometer (Kretschmar and Scott, 1976) was applied on three arsenopyrite bearing samples in equilibrium with pyrite. Arsenopyrite analyses with Fe/(As+S) < 0.5 and/or Ʃ (Co,Ni,Sb) > 1 wt. % were omitted. From core to rim, As at. % variations are correlated with temperature changes represented by local Asp I (28-30 As at. %; 300- 350°C), pervasive Asp II (30-32 As at. %; 360-430°C), and minor Asp III (30-32 As at. %; 400-490°C). Due to higher trace amounts of Ni, Co and Sb in Asp III, seven out of 41 spot analyses of Asp III and two out of 19 analysis of Asp I were dismissed for the application of the arsenopyrite geothermometer. The compositional changes in arsenopyrite not only reflect an increase in temperature but also in the activity of S2 (Kretschmar and Scott, 1976; Barton, 1969). Thus, precipitation of arsenopyrite provides sulfur fugacities approximately between -9.5 and -5.0 and temperature ranges between 300-480°C.

6.3.5. Gold-silver ratio

Twenty-six microprobe spot analyses were conducted on gold grains (n=4) hosted in BIF and carbonaceous schist. The gold content varies between 91.7-94.7 wt %, averaging 93.4 ± 1.5 wt. %, whereas silver varies between 5.0-6.6 wt. %, averaging 5.7 ± 0.5 wt. %. Detection limits for gold and silver are 0.03 and 0.06 wt. %, respectively. Additional elements present in gold include As (3.63 wt. %), Sb (0.46 wt. %), Fe (0.78 wt. %), Pb (0.73 wt. %), Co (0.41 wt. %) and Bi (1.22 wt. %). The gold fineness varies between 929-956, averaging 945 ± 4. Lower fineness ratios are explained by the occurrence of refractory gold rather than free gold (Table A.6).

Fig. 13. Redox and pH conditions. (A) Stability of muscovite and K-feldspar combined with chlorite- based temperature estimates (350°C, 2kbar; modified after Mickuki and Ridley, 1993 and references therein) indicated by dashed area suggest near-neutral fluid composition for the Cascavel and Sertão deposits. (B) Stability of Fe-bearing sulfides and oxides at 350°C, reaction boundaries between Cu- and As-bearing phases (modified after Mickuki and Ridley, 1993 and references therein) shows the evolution from gold-related reduced fluid indicated as I to more oxidizing conditions indicated as II

(aka Tinteiro system). (C) Log activity of sulfur vs temperature (modified after Thorne et al., 2008), outlines buffer assemblages involving major ore-related fluids (I) and paragenetically late alteration (II). Abbreviations: apy: arsenopyrite, py: pyrite, po: pyrrhotite, lö: loëllingite, stb: stibnite; bor: bornite; ccp: chalcopyrite.

7. Discussion Detailed petrographic analyses of the hydrothermal alteration zones and mineral chemistry of key hydrothermal alteration minerals, in both the Cascavel and Sertão deposits, provide constraints on the mineral paragenesis, ore fluid characteristics and gold precipitation mechanism. Combined with lithostratigraphic and structural constraints, this information is used to propose a descriptive model for gold mineralization at the Cascavel and Sertão deposits within the Faina greenstone belt.

7.1. Paragenesis In the Cascavel deposit, two types of pyrite with unequivocal timing relationships are distinguished in distal alteration zones in the biotite schist. Typically deformed, fine-grained pyrite1 (Py1) occurs intergrown with white mica and biotite, whereas subhedral pyrite2 (Py2) replaces Py1 and contains local inclusions of galena. Siderite and ankerite show bleb-like texture and are replaced by Py2. In proximal alteration zones of the biotite schist, tourmaline overprints quartz-white mica-biotite- pyrite±pyrrhotite. Free gold occurs in quartz of the V2 veins. In the Sertão deposit, the proximal alteration zone of carbonaceous schist also displays a fine- grained, deformed Py1 and coarse-grained Py2. Arsenopyrite overprints Py1 and Py2, as well as chalcopyrite and pyrrhotite, and is therefore interpreted to have formed paragenetically late. The typical white mica-siderite/ankerite-Py2-chalcopyrite-arsenopyrite±pyrrhotite assemblage of V2 veins in carbonaceous schist is overprinted by pyrite3-stibnite±galena indicating their late paragenetic stage. Gold occurs within the lattice and as inclusions in arsenopyrite. Locally, gold fills fractures in arsenopyrite, therefore it is interpreted as the paragenetically latest mineral. Pyrite and chalcopyrite from distal BIF are replaced by hydrothermal Fe-rich carbonates. In the proximal zone poikiloblastic Py2 overprints Py1 and local chalcopyrite, whereas arsenopyrite overprints Py1- chalcopyrite±pyrrhotite. Gold occurs within arsenopyrite and as ‘free’ gold disseminated in white mica- rich zones.

7.2. Redox and pH conditions Cascavel: The stability of alkali-feldspar and white mica in proximal alteration zones are indicative of neutral to slightly alkaline pH conditions for the hydrothermal fluids at the Cascavel deposit (Fig. 13A; Mickuki and Ridley, 1993). The presence of pyrrhotite in equilibrium with chalcopyrite suggest reduced fluid conditions (Fig. 13B; Mickuki and Ridley, 1993). Sertão: The stability of alkali-feldspar over K-rich mica and quartz suggest hydrothermal fluids at the Sertão deposit were likely neutral to slightly alkaline (Fig. 13A). The equilibrium assemblage

involving pyrite-arsenopyrite±pyrrhotite in hydrothermal assemblages is indicative of reduced conditions (Fig. 13B). Inclusions of pyrrhotite in arsenopyrite evidence the replacement by pyrite- arsenopyrite±chalcopyrite assemblages, which is consistent with a decrease in oxygen fugacity towards the proximal alteration (Fig. 13B). Thus, the oxygen fugacity tends to vary from strongly reducing to slightly more oxidised conditions (e.g., Mickuki and Ridley, 1993). This is accompanied by an increase in the temperature of ore-forming fluids (Fig. 13C). Lower log aS2 inferred for Asp III is consistent with the presence of stibnite-pyrite (Thorne et al., 2008; Fig. 13C). Chlorite and arsenopyrite geothermometry suggest a broad temperature range of 310-480°C for the hydrothermal fluids at the Sertão deposit. There appears to be a temperature gradient from the distal alteration zone (~ 330°C) to the proximal alteration zone (~ 400°C). Moreover, the presence of hydrothermal minerals such as ankerite, white mica, and the assemblage pyrite-arsenopyrite is compatible with mid to upper greenschist facies conditions at and below 500°C (Anovits and Essene, 1987; Witt and Vanderhor, 1998).

7.3. Transport and precipitation mechanisms of gold

In the Cascavel deposit ‘free’ gold is in V2 veins and pyrite-chalcopyrite proximal alteration zones (≤ 2 m around V2 veins). ‘Free’ gold in veins was likely transported by sulfur complexes (Seward,

1989). The deposition of gold in V2 veins is possibly caused by fluid immiscibility (or boiling) and subsequent lowering of the gold solubility (Brown, 1986; Seward, 1989). The cause of fluid immiscibility in V2 veins within the D2 thrust faults is likely due to cyclic decompression of the hydrothermal fluid caused by seismic movement along the thrust faults and V2 veins (Sibson et al., 1988; Robert et al. 1995; Dugdale and Hagemann, 2001). Disseminated gold associated with pyrite-chalcopyrite in proximal alteration zones, increase of Fe/(Fe+Mg) ratios in chlorite and carbonate towards the proximal alteration zones, neutral pH and high gold fineness values, suggest that gold was transported as a gold bisulfide complex (Seward, 1973; Loucks and Mavrogenes, 1999). Sulfidation reactions in the proximal alteration zone favour gold precipitation (Neall and Phillips, 1987) but K and CO2 metasomatism (Fyfe and Kerrich, 1984; Kishida and Kerrich, 1987) may have also been a gold precipitation mechanism. The infiltration of sulfur- bearing ore fluid resulted in a reduction of aH2S and concomitant deposition of Au (Mickuki, 1998).

Potassium, Ca and CO2 metasomatism of the wall rocks in the proximal alteration zones could have caused acidification of the ore fluid, thereby also contributed to the destabilization of gold bi-sulfide complexes resulting in the precipitation of gold (Kishida and Kerrich, 1987). In the Sertão deposit disseminated mineralization in the pyrite-arsenopyrite±pyrrhotite

- proximal alteration zone reflects reducing conditions and suggests that H2S and/ or HS were potential ligands that favoured the transportation of gold (Phillips and Groves, 1993; Shenberger and Barnes, 1989; Benning and Seward, 1996; Williams-Jones et al., 2009). Under these reducing conditions, gold precipitation is mainly controlled by fluid redox state (Roberts, 1987). Reactions

between S-rich fluids and Fe-rich wallrock reduces the ore fluids (cf. Mickuki, 1998; Fougerouse et al,

- 2017) and causes destabilization of Au (HS )2 complex (Phillips and Groves, 1993). Gold in the Sertão deposit occurs as “invisible gold” in arsenopyrite or as inclusions in arsenopyrite grains. This suggests that gold may have been held in solid solution within arsenopyrite, and exsolved during cooling or recrystallization (Pokrovski et al., 2002).

TABLE 4. Summary of characteristics related to gold mineralization in the Cascavel and Sertão deposits

7.4. Descriptive model for gold mineralization in the Cascavel and Sertão deposits within the Faina greenstone belt The characterization of structural, mineralogical and textural features in the Cascavel and Sertão deposits are used to constrain the relative timing between metamorphism, deformation and hydrothermal alteration in order to propose a descriptive model for gold mineralization in the Faina greenstone belt. This model is defined by five (1 – 5) stages (Fig. 14A-E; Table 4). 1- Deposition of the sedimentary host rocks in the Faina paleobasin: Provenance ages on detrital zircon grains indicate that these are derived from Archean sedimentary protoliths (Brant et al., 2015).

Clasts of volcanic rocks in basal conglomerate (Resende et al., 1998) suggest a period of quiescence after the end of volcanism. Sedimentation is interpreted to post-date regional crustal extension recorded by mafic dike swarms with Sm-Nd ages between 2.3-2.5 Ga (Corrêa da Costa, 2003, 2006).

2- Regional SW-NE shortening during the D1 deformation stage: This stage led to the inversion of the volcano-sedimentary basin into a synform, and development of a penetrative regional S1 foliation defined by metamorphic the quartz-chlorite±alkali-feldspar±biotite assemblage. Locally, barren quartz±calcite V1 veins are developed.

3- Regional NNW-SSE shortening during the D2 deformation stage: This stage is characterized by the development of ~E-W-striking, tight to isoclinal F2 folds and subtly S-dipping axial planar S2 cleavage, L2 intersection lineation, and ~E-W-trending, gently S-dipping thrust faults. Hydrothermal alteration and gold mineralization are associated with the D2 event. In both deposits, the gold-related hydrothermal alteration minerals consistently replace and locally overprint metamorphic minerals indicating that hydrothermal alteration is post-peak metamorphism. Chlorite compositions tend to increase in Fe/(Fe+Mg) towards proximal alteration assemblages, preferentially in Fe-rich rocks, which suggest dependency on host rock composition.

At the Cascavel deposit fluid flux along F2 fold hinges was responsible for the formation of gold- bearing V2 veins and associated hydrothermal alteration zones. Folded, shallow-dipping V2 veins, typically developed between host rocks with contrasting rheology, e.g., quartzite and mica schist, are locally offset by reverse thrust faults. Geometrically, gold mineralization is oriented parallel to F2 fold hinges, which is subparallel to the L2 intersection lineation. Near neutral, reducing, 300-400°C hydrothermal fluids circulated in the V2 veins and also reacted, to a limited extent, with the adjacent wallrocks. Vein-hosted mineralization displays low sulfide content (≤ 1 vol. %) and ‘free’ gold in quartz. Fluid-rock reactions in the quartzite caused the formation of up to 30m wide hydrothermal alteration zones. In quartzite, hydrothermal fluids caused the formation of distinct hydrothermal alteration zones, defined by distal quartz-alkali feldspar±muscovite, intermediate quartz-alkali feldspar- fuchsitic mica±biotite, and proximal quartz-alkali feldspar-fuchsitic mica±biotite assemblages. In biotite schist, hydrothermal alteration zones are defined by distal quartz-chlorite-calcite and proximal white mica-biotite-ankerite±tourmaline assemblages. Chlorite geothermometry in biotite schist displays a small but distinct temperature gradient between intermediate and proximal alteration zones and V2 veins from 385°±6°C to 406°±4°C and 417°±11°C, respectively.

At Sertão, circulation of hydrothermal fluids along D2 fold hinges lead to the formation of local V2 veins and associated hydrothermal alteration zones. Gold mineralization is preferentially disseminated and hosted in proximal alteration zones within carbonaceous schist and BIF, characterized by pyrite-arsenopyrite-chalcopyrite±pyrrhotite assemblages. Gold is refractory in arsenopyrite, but also occurs as inclusions in arsenopyrite, and rarely as ‘free’ gold in quartz. Hydrothermal fluids are characterized by 320-430°C (based on chlorite and arsenopyrite geothermometry), near neutral pH and reducing redox state. Interaction of these hydrothermal

fluids with carbonaceous schist and BIF caused distinct fluid-rock reactions characterized by distal quartz-ankerite-pyrite-chlorite and proximal quartz-white mica-siderite-pyrite- arsenopyrite±chalcopyrite assemblages.

Fig. 14. Schematic block diagrams illustrating the geological evolution of the Faina greenstone belt (A- E), with emphasis on the structural evolution and hydrothermal alteration zones including equilibrium assemblages associated with gold mineralization in the Cascavel (F) and Sertão (G) deposits.

4- Regional NE-SW shortening during the D3 deformation stage: Structures related to this stage include NW-striking shear zones, moderately S-dipping S3 foliation and local NNW-striking F3 open folds. These structures are typically best preserved in the Sertão deposit area. The existing hydrothermal minerals in both deposits, such as biotite and tourmaline in biotite schist, fuchsitic and white mica in quartzite, arsenopyrite and white mica in carbonaceous schist and BIF, are mostly oriented parallel to the dominant S2 foliation but, in many places reoriented during D3.

At Cascavel, ‘free’ gold in D2 structures is locally oriented parallel to NW-striking D3 shear zones, evidence that gold-bearing D2 structures were locally re-oriented. In addition, minor tourmaline, galena and trace stibnite overprint D2-related proximal alteration assemblage. These newly formed minerals are evidence of local hydrothermal fluid flow during D3.

At Sertão, gold-arsenopyrite in D2 fold hinges are reoriented parallel to NW-trending D3 shear zones. Although there is a lack of hydrothermal alteration and mineralization during the D3 event, at a centimetre scale limited hydrothermal fluid flow and recrystallization of V2 veins is evidenced by: (i) comminution of quartz-siderite±ankerite grains near the borders of the vein; (ii) dissolution of arsenopyrite, represented by core to rim chemical zoning, with arsenopyrite in many places displaying pyrrhotite inclusions; (iii) common fractures in arsenopyrite grains; and (iv) presence of paragenetically late pyrite±stibnite assemblage and galena in V2 quartz-(Mg-siderite) veins.

5- Regional NE-SW to E-W shortening during the D4 deformation stage: The last stage comprises moderate S-dipping thrust faults and associated sub-horizontal S-plunging slickenlines, steeply W- dipping reverse faults and associated steeply S-plunging slickenlines. Common V4 fault-fill veins and breccias are also developed during this event. Additional minor structures may include NNW-striking shear zones. These cross-cut and locally offset (up to m-wide) D1 to D3 structures. At present, only limited whole rock and trace element data are available (internal report Orinoco, 2016), which suggest enrichment of Ag, W, Mo, Pb, Mn, U, Zn and no apparent gold concentration. Paleoplacer gold in the upper conglomerate of the Faina greenstone belt was reported by Carvalho et al. (2013). According to these authors, gold is associated with a fine-grained quartz matrix and within deformed clasts of mineralized quartz veins, BIF, meta-volcanic and -sedimentary rocks. There are no geochronological constraints on the deformation and hydrothermal events in the Faina greenstone belt. However, geochronological data from the northern greenstone belts in the Goiás Archean Block (Queiroz, 2000; Queiroz et al., 2008; Marques et al., 2013; Jost et al., 2014), approximately 150 km north of the Faina greenstone belt, suggest a Paleoproterozoic (Transamazonian) age for the gold-related, dominantly ductile, deformation events (D1-D3). The late brittle (-ductile) structures comprising the D4 deformation event are suggested to be associated with the Neoproterozoic Brasiliano (-Pan-African) orogeny.

8. Comparisons with other orogenic gold deposits Gold mineralization at the Faina greenstone belt has similar characteristics with other typical Archean and Paleoproterozoic orogenic gold deposits worldwide (Groves et al. 1998, 2003; McCuaig and Kerrich, 1998; Goldfarb et al. 1986, 2001, 2005). Some of these characteristics include: (i) greenstone belt setting with regional greenschist to upper amphibolite facies metamorphic host rocks; (ii) spatial association with major regional shear zones and strike- slip faults (Sibson et al., 1988; Robert and Poulsen, 2001; Goldfarb et al., 2005; Robert et al., 2005), such as the Faina Fault (≤5 km from Sertão deposit); (iii) hydrothermal alteration and gold

mineralization are structurally controlled; (iv) gold located in quartz veins and proximal alteration zones within pyrite and arsenopyrite; and (v) temperature estimates for hydrothermal alteration minerals between 300-450°C and pressures broadly < 3 kbars, which suggests a mid-crustal (mesozonal) level of ore formation. In such deposits, orebodies vary according to ductile-brittle nature dependent on host rock composition (Robert et al., 2007).

TABLE. 5. Comparison of geological characteristics of gold deposits in the Faina greenstone belt with gold deposits hosted in other greenstone belts within the Goiás Archean Block

Both Cascavel and Sertão deposits are hosted in metasedimentary rocks adjacent to tectonic unconformities marking the limits between lower volcanic and overlying sedimentary package. The development of orogenic gold deposits in close spatial relation with discontinuities marking the limits between volcanic and sedimentary rocks is suggested by several studies (e.g., Robert et al., 2005, 2007). Reactivation of regional high angle normal and reverse faults controlling sediment deposition through time is widely accepted as a key targeting feature for exploration of orogenic deposits. Early formed structures reworked during later deformation events were likely transformed into potential corridors for fluid flow and may represent a strong control on the distribution of sedimentary sequences hosting gold mineralization. Therefore, paleogeographic settings during the sedimentation of the Faina paleobasin,

or its early crustal architecture, may have played an important role in developing structural pathways potentially reactivated during protracted deformation. According to the classification of orogenic gold deposits by Groves et al. (1998), the gold-related hydrothermal assemblages at Cascavel and Sertão, combined with temperature estimates, suggest the deposits are compatible with mesozonal Au-As (6-12 km, 300° – 475°C) conditions. In addition, gold is likely to be mostly transported by reduced S-rich fluids (Phillips and Groves, 1993; Table 4). At the Sertão deposit, the host rocks are carbonaceous schist and BIF with the bulk of gold disseminated in shear zones and located in or close to fold hinges. Similar deposits include examples in the Quadrilátero Ferrífero (Cuiabá, São Bento, Lamego, Raposos), in the Archean Goiás Archean Block (Crixás, Pilar de Goiás), and in the São Francisco Craton (Fazenda Brasileiro, at the Rio Itapicuru greenstone belt). Worldwide comparisons can be drawn with Lupin (Northwestern Territories, Canada), Vubachicke (Zimbabwe), and Mt. Morgans (Yilgarn, Western Australia). At the Cascavel deposit, the quartzite host rock is less common when compared to host rocks in other orogenic gold deposit. Analogous metasedimentary host rocks for gold mineralization are encountered in the New Holland gold deposit in the Yilgarn Craton of Western Australia (e.g. Ackroyd et al., 2001). The structural control of gold mineralization via shallow dipping thrust faults is common in the Goiás Archean Block (e.g., in the Crixás gold deposit; Table 5), but in contrast to many other orogenic gold deposits in Canada and Western Australia where gold ore bodies are located in, and controlled by, steep dipping reverse fault zones and associate ‘steep’ and ‘flat’ vein systems (e.g., Sigma mine in Quebec) or strike-slip fault systems with associated veins and breccias (e.g., Wiluna and Golden Mile deposits in Western Australia).

9. Conclusions and Implications for exploration The Faina greenstone belt, in the southern Goiás Archean Block, is characterized by orogenic gold mineralization, including the Cascavel and Sertão deposits. These deposits display similarities in structural settings, hydrothermal alteration assemblages, zoning, and temperatures of hydrothermal alteration and mineralization when compared to other orogenic gold systems in Brazil and worldwide (Hagemann and Cassidy, 2000). These include: (i) structural control of gold mineralization via shear zones and folds, (ii) hydrothermal minerals and zones that surround the gold-bearing structures and veins, e.g., white mica±fuchsitic mica-pyrite ±chalcopyrite, are ubiquitous in proximal alteration zones of both deposits, (iii) overall increase in the Fe content towards proximal alteration zones, and (iv) temperature estimates of gold-bearing hydrothermal fluids between 330-420° at Cascavel and 320- 430°C at Sertão. At Cascavel, ‘free’ gold in deformed quartz veins hosted in quartzite is mostly controlled fold hinges, with minor disseminated mineralization in hydrothermally altered wallrock. The trending D3 shear zones locally deform D2-related orebodies, and are associated with minor fluid represented by the precipitation of paragenetically late minerals (e.g., tourmaline, galena) in the

related hydrothermal alteration zones. At Sertão, refractory gold is locally present within V2 veins and disseminated sulfide-rich quartz-siderite-white mica proximal alteration zones that surround the D2 shear zones. The gold-bearing D2 structures such as fold hinges are reoriented parallel to NNW-trending

D3 shear zones. Both the Cascavel and Sertão deposits represent epigenetic, mesozonal orogenic gold systems. A Paleoproterozoic age for gold mineralization is proposed based on the deformation history and alteration characteristics of similar gold systems in other greenstone belts of the Goiás Archean Block (Tassinari et al., 2006; Jost and Queiroz, 2008; Marques et al., 2013). There are presently no absolute age dates available for gold mineralization in the Faina greenstone belt, therefore further geochronological investigations are required to constrain the age of the hydrothermal alteration and mineralizing systems in the Faina greenstone belt. The characteristic alteration pattern including white mica, chlorite and carbonate can be applied as an exploration tool using hand-held or airborne hyperspectral analyses as demonstrated, for example, in the Kanowna Belle and Sunrise Dam orogenic gold deposits in Western Australia (Wang et al., 2017). In the case of the Sertão deposit, the appearance of idioblastic arsenopyrite is an excellent indicator for gold-bearing zones. Chemical changes depicted by mineral chemistry support the definition of an alkali- rich proximal alteration zone in the Cascavel and Sertão deposits, therefore using the alkali index K/Al (Kishida and Kerrich, 1987) in whole-rock geochemistry data may be useful to track potential ore zones.

Due to the widespread carbonate alteration around ore zones the carbonation index (CO2/ (Fe+Mg+Ca) can be used in whole-rock geochemistry data.

Acknowledgements The database presented at this research is a result of a PhD research project completed by Jessica Bogossian, developed at the CET-UWA (Centre for Exploration Targeting, University of Western Australia), and funded by CNPq (“Conselho Nacional de Desenvolvimento Científico e Tecnológico”; Project No. 207220/ 2014-0). The author is grateful to CNPq for the research grant. The local mining company, Orinoco Gold Limited, is thanked for their permission to publish information here and for the support given during the fieldwork campaigns. To Grant Alan ‘Rocky’ Osborne, for his predisposition to discuss ideas and for the constructive suggestions with geology in the Sertão deposit. To Quentin Masurel, Eliza Smith, Wally Witt and Dan Wiemer for the valuable and patient revisions.

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Introduction to Chapter 3: “U-Pb and Hf-O evidence for Mesoarchean crustal growth in central Brazil”

Chapter three of the thesis presents the results of an integrated study involving whole-rock, trace element and REE geochemistry, geochronology and isotopic analyses carried out on regional intrusive rocks within and along margins of the southern Goiás Archean Block (GAB). This chapter will be submitted to the journal “Precambrian Research” with the candidate J. Bogossian as the first author and A.I.S. Kemp and S. G. Hagemann (Centre for Exploration Targeting, University of Western Australia) as co-authors. Previous studies have assigned a complex and multiphase magmatic history to the Goiás Archean Block (Jost et al., 1993; Queiroz, 2000; Pimentel et al., 2003; Jost et al., 2005, 2013; Queiroz et al., 2008). However, the still unclear tectono-magmatic framework precludes the establishment of a comprehensive model that reflect sources and processes active during the formation of the Goiás Archean Block. Moreover, since the Archean TTGs have been considered devoid of economic potential (Jost et al., 2014), most studies in the region have focused on the greenstone belts, particularly in northern Goiás Archean Block. In order to address these issues, combined whole-rock geochemistry, in-situ zircon U-Pb SHRIMP geochronology and Hf-O isotopic analyses were conducted to help decipher the crustal evolution of the southern Goiás Archean Block.

The increasing use of zircon Hf isotopes to assess growth and differentiation of the continental crust (Patchett, 1983; Kemp and Hawkesworth, 2013; Vervoort and Kemp, 2016), makes the application of this method attractive to unravel the long-contended and complex crust-mantle evolution in the study area. The Hf isotopic ratios are used to determine whether magmas are derived from juvenile mantle sources or reworked continental crust. As oxygen isotopes are fractioned during near-surface water- rock interactions they can be used to evaluate the contribution of supracrustal rocks in magma genesis. The U-Pb zircon crystallization ages, associated inheritance and isotopic disturbance were used to constrain the relative timing of regional magmatism, co-parental relationships and deformation events. The framework provided by this study is also applied to investigate the role of geodynamic shifts associated with the formation and distribution of gold systems in the southern Goiás Archean Block.

The sampling of eight regional igneous rocks included Archean TTGs (n=3), Pink Syenite intrusion and granites enclosed along the east (n=2) and western (n=2) margins of the southern Goiás Archean Block. The Pink Syenite consists of the first felsic intrusive identified in the Faina greenstone belt. Until this time, no absolute isotopic age is available for gold mineralization in the region. However, the record of Au-related foliation in Pink Syenite intrusion allows estimating the relative timing of gold mineralization in the region. This enables comparisons with available mineralization ages from other gold deposits in the Goiás Archean Block and similar deposits worldwide.

The following analytical techniques were applied in this chapter: (i) optical petrography (transmitted and reflected light) on polished thin-sections of granitic rocks, which was supplemented by semi-quantitative secondary electron microscopy (SEM-EDS), (ii) whole-rock, trace element and REE geochemistry, and in-situ zircon (iii) U-Pb (SHRIMP) geochronology, (iv) oxygen isotopic analyses by SIMS (Cameca 1280), and (v) Lu-Hf (LA-ICP-MS) isotopes.

The collection of samples and laboratory work, e.g., crushing, grinding, mineral separation, mounting and imaging, of granitic rocks were conducted by the candidate. Polished thin-sections were prepared by Vancouver Petrographics in Vancouver, Canada. The zircon U-Pb SHRIMP analyses were conducted by the candidate at the John de Laeter Centre, Curtin University. Oxygen isotopes via SIMS analyses were performed under the supervision of Dr. H. Jeon at the Centre for Microscopy, Characterisation and Analysis at the University of Western Australia. The LA-ICP-MS Hf analyses were acquired with the assistance of Dr. A.I.S. Kemp who also handled the data reduction. Interpretation of the SHRIMP, SIMS and LA-ICP-MS data was carried out by the candidate with the assistance of the Dr. A.I.S. Kemp. All written work was completed by the candidate with discussions and revisions provided by Dr. A.I.S. Kemp and Prof. S.G. Hagemann.

Chapter 3

U-Pb and Hf-O evidence for Mesoarchean crustal growth in central Brazil

Abstract Evidence for early Earth geodynamics can be provided by methods sensitive to magmatic inputs such as Lu-Hf isotopes. Age, composition and relative contribution of magma sources can be inferred from isotope-time trends. This investigation presents a combined study of whole-rock geochemistry and in-situ zircon U-Pb and Hf-O isotopic compositions of eight granitic rocks spanning over 1500 m.y. in the southern Goiás Archean Block (GAB), central Brazil. Crustal growth models for the GAB are controversial, but current interpretation predicts assembly after consecutive accretion of exotic terranes. Disparate evolutionary paths expected for each terrane favour the allochthonous nature proposed for the GAB. However, coherence in Hf and O isotopes across the southern GAB supports a mutual evolution of these terranes. Zircon Hf isotope compositions at (i) 2870-2820 Ma are mostly near chondritic, with more radiogenic values with decreasing age (2820 Ma Paus de Choro granite), (ii) 2060 Ma show a steep shift to the unradiogenic Pink Syenite, (iii) 630-610 Ma, indicate wide range of unradiogenic compositions in K-rich granites, and (iv) 530 Ma, define a steep trend to the strongly unradiogenic Serra Negra granite. Oxygen isotopes in zircons are (i) largely mantle-like until 2820 Ma, after which they rise to heavier values marking the onset of subduction, (ii) overall mantle-like at 2060 Ma, (iii) abruptly 18O-rich at ~ 600 Ma due to widespread supracrustal reworking that sharply switches to (iv) mantle-like values at 530 Ma. The near chondritic Hf isotope evolution across Archean terranes designate ≤ 70 m.y. of relatively continuous mantle-derived TTG magmatism consistent with early rifting and greenstone belt formation. The shift to more radiogenic signatures at ~ 2820 Ma reflects the onset of subduction near the Meso- Neoarchean transition prior to collisional orogeny. Early crustal extension, sedimentation and subsequent collision estimated for the Paleoproterozoic favoured reworking and minor recycling indicated by the Pink Syenite in the Faina greenstone belt. Analogous syn-tectonic intrusions correlated with the Atlantica assembly enhance the targeting exploration potential of this intrusive event, particularly for gold. Nearly coeval emplacement of 630-610 Ma K-rich granites with regional peak metamorphism in the Brasiliano/Pan-African orogeny support reworking of previous crustal material and pervasive sedimentary and/or supracrustal recycling. The progression to the 530 Ma Serra Negra granite involves a sharp isotopic shift. These Neoproterozoic magmatic events are consonant with the assembly and stabilization of western Gondwana. Similar chemistry and isotopes of the Archean Uvá

orthogneiss and the Cambrian Serra Negra granite suggest a shared evolution proposed for terranes in the southern GAB. Reworking of an old protolith is elucidated by the Archean and Paleoproterozoic inherited zircon ages obtained for the Pink Syenite and Serra Negra granite. This implies that crustal reservoirs in the GAB remained fairly constant through time. Geological and isotopic constraints on crustal evolution in the southern GAB reveal continental growth dominantly through in-situ additions as opposed to the accretion of allochthonous terranes. Identical isotopic patterns documented in other Archean cratons worldwide argue for a global process.

Keywords: U-Pb geochronology, Hf-O isotopes, crustal growth, Goiás Archean Block, central Brazil

1. Introduction Geodynamic and petrogenetic processes involving Archean crustal growth are a long-standing controversy (Kemp and Hawkesworth, 2003; Smithies et al., 2009; Condie and Aster, 2010; Kamber, 2015). Lower viscosities and higher degree of partial melting, inferred from Archean high geothermal gradient, allowed the formation of thicker basaltic crust (Herzberg et al., 2010; Nisbet et al., 1993) and lower crustal rheology (Sizova et al., 2010) that support deformation and delamination instead of subduction (Johnson et al., 2014). Nonetheless, chemical and temperature shifts throughout the Archean may project major changes in geodynamic regimes (Brown, 2006; Korenaga, 2008; Dhuime et al., 2012; Shirey and Richardson, 2011). It is uncertain if conditions during this time led to the development of subduction zones currently accepted as key sites to generate crust (Rudnick, 1995). Archean cratons record prevailing “arc-like” tonalite-trondhjemite-granodiorite (TTG; Drummond and Defant, 1990; Moyen, 2011), and “plume-like” komatiite-basalt (Hollings et. al., 1999; Van Kranendonk et al., 2015). Consequently, Archean crustal growth models contemplate a contradiction between subduction and plume settings, as well as a dilemma about the onset of plate tectonics (Rey et al., 2003; Stern, 2005; Bédard, 2006; Bédard et al., 2013; Van Kranendonk, 2011; Wyman, 2013; Stern et al., 2016). Evidence for continental crust formation is provided to reconcile Archean geodynamic models. Crust formation settings predicting the availability and depth at which magmas stall and fractionate (Chiaradia, 2015; Loucks, 2014; Lu et al., 2015) are associated with characteristic life cycles manifested in their isotopic and chemical records (Collins, 2002; Ducea et al., 2015; Nash et al., 2006). This is illustrated by cyclic shifts in Phanerozoic accretionary margins that reveal unradiogenic Hf/Nd isotopic signatures during compression, whereas radiogenic signatures typify extension (DeCelles et al., 2009; Kemp et al., 2009a). Nevertheless, juvenile input associated with intraplate (Nash et al., 2006) or convergent margin magmatism (Kemp et al., 2009a) can cause isotopic variety (Champion and Huston, 2016). In central Brazil, a fragment of Archean crust of uncertain origin known as GAB was amalgamated into the mid-western margin of the Brasília fold belt during the Brasiliano/Pan-African orogeny (Pimentel et al., 2000a). Accretionary growth and stabilization of cratons can trigger processes

such as slab break-off, with associated magmatism and hydrothermal fluids, emplacement of granitic rocks and potentially gold mineralization (Halla et al., 2017). The collage of the GAB into the Brasilia fold belt resulted in sparsely distributed granitic intrusions, partial anatexis of Archean TTGs, and gold- related hydrothermal alteration (Fortes et al., 2003; Jost and Fortes, 2001; Jost et al., 2014; Tassinari et al., 2006; Queiroz et al., 2008). In addition to the still uncertain origin of this cratonic fragment, contrasting Nd/Hf isotopic signatures argue for the exotic nature of the GAB (Jost and Fortes, 2001; Jost and Queiroz, 2008). Regardless the complex evolution recorded by the southern GAB (Beguelli Jr., 2012; Pimentel et al., 2003a; Jost et al., 2005, 2013), the tectono-magmatic history delineated by this terrane remains under-constrained. Therefore, this area makes an attractive case study to track magma sources and to evaluate the nature of magmatism reflecting major geodynamic changes throughout time. In this contribution, we report combined whole-rock geochemistry, in-situ zircon U-Pb geochronology and Hf-O isotopes of eight granitic rocks of the southern Goiás Archean Block. The results of this study are used to pinpoint the setting, nature and magma sources involved in the petrogenesis of the southern GAB. Distinct geochemical trends, ages and isotopic compositions elucidate a complex geological evolution reflected by major crustal growth during the Archean with reworking in the Paleoproterozoic and Neoproterozoic. Field, petrological and isotopic constraints are used to reconstruct the magmatic and tectonic history of this relict Archean continental crust. The investigation of terranes in the southern GAB are supported by the new Hf-O isotope dataset but proposed allochthone nature (Jost and Fortes, 2001; Pimentel et al., 2011; Jost et al., 2013) is contested. These findings challenge the allochthonous nature proposed for the GAB and provide evidence for Paleoproterozoic to Neoproterozoic Wilson cycles.

2. Regional Geology The collision between the Amazonian, São Francisco-Congo and Paranapanema cratons during the amalgamation of West Gondwana in the Brasiliano/Pan African orogeny resulted in the formation of the Tocantins Province (Almeida et al., 1981; Marini et al., 1984; Brito Neves et al., 1999). The province (Fig. 1A) is composed of three mobile fold belts known as Araguaia, Paraguai and Brasília (Pimentel et al., 2000a). Bordering the western margin of the São Francisco craton, the Brasília fold belt (1000km N-S length and ~ 250 km wide) forms one of the most complete and well-preserved orogens in the South American Platform (Fuck et al., 2014). It includes: (1) sedimentary sequences typical of continental margins that are progressively more deformed towards the west (Fuck et al., 2014), (2) mafic-ultramafic layered complexes (Ferreira Filho et al., 1994), (3) volcano-sedimentary sequences and tonalite/granodiorite gneisses from the Goiás magmatic arc (Pimentel and Fuck, 1992; Pimentel and Charnley, 1991; Pimentel et al., 1997), (4) granulite metamorphic core (Anápolis-Itauçu Complex; Piuzana et al., 2003), ophiolite slices (Pimentel et al., 2000a), and (5) fragment of Archean crust known as GAB (Fuck et al., 2014).

Fig. 1. Geologic setting of Brasília fold belt (modified after Pimentel et al., 2004), within the Tocantins Province (Inset A). Sketch displaying the location of greenstone belts and Archean TTG complexes of the GAB (Inset C), with the location of Fig. 2.

The GAB was assembled into the Brasília fold belt during the Brasiliano/Pan-African orogeny (Pimentel and Fuck, 1994; Pimentel et al., 2000a). It consists of voluminous Archean TTG complexes and narrow Archean-Paleoproterozoic greenstone belts (Jost and Queiroz, 2008; Fig. 2) and rare Neoproterozoic intrusions (Rodrigues, 2011; Pimentel et al., 2003b). The GAB is mainly composed of Archean to Paleoproterozoic rocks, and it is bordered by the Goiás magmatic arc in the north (Mara Rosa Arc) and in the south (Arenópolis Arc). The northwest and southwest margins are delimited by the Transbrasiliano Lineament and the Moiporá-Novo Brasil shear zone, respectively. The western margin comprises sedimentary rocks of the Araguaia basin and the eastern margin consists mostly of metasedimentary rocks of the Serra Dourada and Araxá groups (Fig. 2). The TTG complexes of the GAB comprise six intensely deformed orthogneisses of variable composition and age that were best delineated by gamma-spectrometry (Blum et al., 2003). Regional studies indicate Archean stabilization at ~ 2700 Ma (Jost and Queiroz, 2008), Paleoproterozoic magmatism after crustal extension in the Siderian and closure of the orogen in the Rhyacian (Danni et al., 1986; Jost et al., 1992, 1993, 2010, 2014; Queiroz, 2000; Corrêa da Costa, 2003). The polycyclic evolution of the GAB resulted in its subdivision into northern and southern

portions (Montalvão, 1986; Pulz, 1995; Fortes et al., 1997; Pimentel et al., 2003a; Queiroz et al., 2008; Jost et al., 2010; Beguelli Jr., 2012). TTG magmatism in the northern GAB encompass: (i) juvenile (εNd = +2.4 to -1.0), 2840-2780 Ma (zircon SHRIMP U-Pb), batholith-sized tonalite, granodiorite and granite orthogneisses, and (ii) crustal-derived (εNd = -2.2), 2700-2790 Ma (zircon SHRIMP U-Pb) dike- like granodiorite to granite orthogneisses (Queiroz et al., 2008). The 2790 and 2710 Ma main metamorphic events are coeval with younger TTG intrusions in the northern GAB (Queiroz et al., 2008; Jost et al., 2013). TTG terranes in the southern GAB are represented by the Uvá and Caiçara complexes, separated by the Faina and Goiás greenstone belts (inset Fig. 2). The study area covers approximately 700km2 of the southern GAB (Fig. 2). Samples include intrusions from the Archean TTGs (Caiçara, Uvá and Paus de Choro), the eastern (Itapuranga I and II) and western (Caiapó and Serra Negra) margins of the southern GAB and an intrusive rock in the Faina greenstone belt (Pink Syenite). The geological background of the studied samples is detailed below and provided in Table 1. Table 1. Compilation of available geochronological data of main geological units in the southern GAB

2.1. Geology of the Archean rocks - Caiçara, Uvá and Paus de Choro Initial Rb-Sr and Sm-Nd geochronology data of the Caiçara and Uvá complexes by Tassinari and Montalvão (1980), Montalvão (1986), Vargas (1992), Tomazzoli (1992), and Pimentel et al. (1996) showed a Neoarchean time-span but were unable to define the genesis of these complexes. The first U- Pb SHRIMP zircon ages of diorite gneiss from the Uvá TTG obtained by Pimentel et al. (2003a) yielded ~ 2930 Ma and 3090 Ma for crystallization and inherited zircon ages, respectively, with whole-rock model ages of ~ 3000 Ma and εNd between -4.6 and +0.4, indicating a contribution from an older crust. The Uvá TTG complex was later characterized as two groups of intrusions (Jost et al., 2005, 2013). The first and major group, known as Rio do Índio domain, is composed of intensely deformed batholith-like tonalite to granodiorite orthogneisses and minor diorite stocks, with zircon LA-ICP-MS U-Pb crystallization ages between 3040-2930 Ma, and positive εNd values for > 3000 Ma samples, suggesting

a juvenile crustal affinity (Jost et al., 2013). The second group, known as the Rio Vermelho domain, is composed of tabular tonalite, monzogranite and granodiorite with zircon LA-ICP-MS U-Pb ages of approximately 2880 and 2850 Ma for the monzogranite and tonalite, respectively, with slightly negative εNd values (Jost et al., 2013) supporting minor reworking of older source. The same authors suggest a metamorphic event recorded by those rocks at ~ 2830 Ma.

Fig. 2. Geological map of the southern GAB (modified after CPRM - Brazilian Geological Survey).

The Paus de Choro consists of coarse-grained, pink to white, granular to proto-mylonitic muscovite granite with four distinct magmatic facies, suggesting a multiple intrusion (Jost et al., 2005). The tabular bodies of the Paus de Choro granite share fault-related tectonic contacts with the Uvá orthogneiss, being defined as a late pulse of the latter (Jost et al., 2005). The Caiçara complex is composed of tonalite and granodiorite orthogneisses with minor granite and charnockite (Beguelli Jr., 2012). Zircon LA-ICP-MS U-Pb crystallization age of ~ 3140 Ma and ~ 2820 Ma, with Sm-Nd model ages of ~ 3100 Ma and ~ 2900 Ma are reported for the tonalite and intrusive granodiorite orthogneiss, respectively (Beguelli Jr., 2012).

2.2. Geology of the eastern border - Itapuranga I and II The Itapuranga suite (Oliveira, 1994) comprises a cluster of subalkaline potassic intrusive rocks, including pink syenite, alkali-feldspar granite, granite, quartz diorite, granodiorite and tonalite (Souza et al., 1993; Lacerda Filho and Oliveira, 1994, 1995; Oliveira, 1997). They form elongate intrusions along E-W trending lineaments (aka Pirineus Syntax) that support their contemporaneous emplacement

with the development of these regional faults during peak metamorphic conditions in the central portion of the Brasilia Fold Belt (Pimentel et al., 2003b). Contacts with the intrusive Itapuranga suite comprise shallow dipping, oblique contraction structures (Baêta Jr. et al., 1999). Available geochronology includes zircon U-Pb SHRIMP 624 ± 10 Ma crystallization age, 1790-1490 Ma inherited cores, Sm-Nd 1440 Ma model ages and εNd values between -5.1 and -5.7, suggesting older sialic crust as part of parental magmas (Pimentel et al., 2003b).

2.3. Geology of the western border - Rio Caiapó and Serra Negra The Rio Caiapó suite comprises syn- to post-tectonic plutons, e.g., quartz monzodiorite, quartz monzonite, granodiorite and granite, with calk-alkaline affinity (Pimentel and Fuck, 1986, 1987, Seer, 1985, Pimentel et al., 1985). These granites have whole-rock Rb-Sr ages between 585 and 759 Ma (Pimentel and Fuck, 1994, 1996), ~ 1100 Ma model ages and +3.5 εNd values (Pimentel and Charneley, 1991). The granites from the Serra Negra suite were emplaced after the last deformation and metamorphic events that affected the Goiás Magmatic Arc (Pimentel et al., 1996). Whole-rock Rb-Sr geochronology of this igneous suite yield ages between 524-508 Ma (Pimentel and Fuck, 1987). Ages between 500-570 Ma, reported for zircon LA-ICP-MS U-Pb analyses, include a sample with a 506 ± 4 Ma igneous age and 2680 ± 20 Ma concordant inherited core (Marques, 2017).

2.4. Geology of the Pink Syenite The Pink Syenite is the first record of a felsic intrusion within the Faina greenstone belt (FGB). The stratigraphy of the FGB consists of basal komatiites followed by tholeiitic basalts unconformably overlain by platform-like sedimentary sequence culminating with dolomite, chert and banded (-iron) formation (Resende et al., 1999). Volcanic and plutonic rocks have zircon LA-ICP-MS U-Pb ages of ~ 2960-2920 Ma (Borges et al., 2017). An Archean protolith for the detrital rocks in the Faina greenstone belt is inferred from Sm-Nd model ages and zircon LA-ICP-MS U-Pb provenance ages that suggest maximum age for sedimentation between ~ 3200-2800 Ma (Resende et al., 1998; Brant et al., 2015). The minimum age of sedimentation is constrained by stable isotopes on upper carbonates from all the GAB greenstone belts. The positive δ13C values (+10.0 to +14.0 ‰) of upper marble of northern belts and lower marble of southern belts suggest their sedimentation was completed at the Huronian glaciation at the beginning of the Rhyacian (Fortes, 1996; Resende et al., 1998; Jost and Queiroz, 2008; Santos et al., 2008; Jost et al., 2014). The δ13C values (-0.7 and +0.7 ‰) from upper marble of southern greenstone belts suggest minimum age of deposition near the transition from Rhyacian to Orosirian times (Resende et al., 1999; Jost et al., 2014).

3. Methodology Eight representative samples of igneous rocks were collected from outcrops in the southern GAB. Polished thin-sections from each of these samples were studied using optical microscopy and scanning

electron microscope (SEM) microscopy at the Centre for Microscopy, Characterization and Analyses (CMCA – UWA). Whole-rock major, trace and rare earth element analyses were conducted by Bureau Veritas (Perth, Australia). The samples were cast using a 66:34 flux with 4% Lithium nitrate added to form a glass bead. Major element oxides were analysed by XRF (X-ray fluorescence spectrometry). Fused bead laser ablation (ICP-MS) was utilized to analyse 51 trace elements (see Tables 2, 3). Gold, Pt and Pd were measured by Inductively Coupled Plasma (ICP) Optical Emission Spectrometry. Carbon and S were quantified by total combustion analysis, and FeO was determined volumetrically after acid digestion. The Geological Survey of Western Australia (GSWA) rock standard KG1 (Kerba Monzogranite; Morris, 2007) was analysed to monitor data quality. For the zircon work, the samples were crushed, milled, sieved (50 mesh), panned and magnetically separated (Frantz) to obtain heavy mineral concentrates. Heavy minerals within the non- magnetic portion were subsequently separated using LST (lithium heteropolytungstates; density of 2.95 g/mL at 25°C). The zircon grains were handpicked and mounted in epoxy discs and polished (JB 1 and JB2). Each mount included four samples and reference zircons (OGC, BR266, Penglai, Temora 2; see below). Prior to the analysis, optical (transmitted, reflected) and electronic (CL, BSE) microscopy was conducted to choose the location of spot analysis (avoiding cracks, inclusions or any other irregularities in their morphology/composition). After imaging, the mounts were cleaned and gold-coated to create the conductive surface required for the SHRIMP analysis. The two principal standards used for U-Pb dating were BR266 (206Pb/238U and 207Pb/206Pb ages of 559.0 ± 0.3 and 562.2 ± 0.5 Ma, 903 ppm U; Stern, 2001) and OGC (207Pb/206Pb age of 3465.4 ± 0.6 Ma; Stern et al., 2009), with the latter being used to monitor the 207Pb/206Pb ratio. The U-Pb analyses were conducted by using the SHRIMP II – Sensitive High-mass Resolution Ion MicroProbe (De Laeter and Kennedy, 1998), at the John de Laeter Centre (Curtin University), based on the methodology of Compston et al. (1992). Each zircon analysis consisted of six scans with five

196 204 206 207 208 238 248 254 channels each ( Zr2O, Pb, background, Pb, Pb, Pb, U, ThO and UO). Corrections for common lead were made using the measured 204Pb. For each spot, a raster of 60-180 s was used for pre- sputtering the mount surface and to minimise contamination from surface 204Pb. Zircon data were reduced using the Microsoft Excel add-in Squid (Ludwig, 2002). The U-Pb concordia diagrams, weighted average and probability plots were made using Isoplot v. 3.71 (Ludwig, 2003), with uncertainties plotted at the 95% confidence level (2σ). Neoproterozoic samples had ages calculated using 206Pb/238U ratio, whereas ages for the Archean (-Paleoproterozoic) samples are based on 207Pb/206Pb ratios. The accuracy of the data was monitored by repeated analysis of standards, including OGC (207Pb/206Pb age of 3463 ± 6 Ma, n = 18) and BR266 (206Pb/238U age of 559 ± 3 Ma, n = 62). The analyses were preferentially conducted in clear prismatic grains rather than in dark stubby zircons displaying patchy core and rim zoning. A summary of the acquired data is presented in Tables 4 and 5.

The mounts were repolished to enable oxygen and hafnium isotope analyses (Table 6) in the same area of previous SHRIMP U-Pb dating. The in-situ Hf-O isotopic measurements were preferentially undertaken on polished zircon grains with less than 10% U-Pb age discordance, i.e., the percentage difference between the 238U/206Pb* and the 207Pb/206Pb* ages. However, considering ancient lead loss can result in vectors that are subparallel to concordia over millions of years (e.g., Whitehouse and Kemp, 2010), concordance seems to be a less reliable index of disturbance in Archean zircons (Kemp and Hawkesworth, 2014). The oxygen isotope composition (18O/16O ratio) of zircons was measured by secondary ion mass spectrometry (SIMS) using a Cameca IMS 1280 multi-collector ion microprobe at the CMCA – UWA. The methodology applied for oxygen analyses follows that of Nemchin et al. (2006) and Whitehouse and Nemchin (2009). For the hafnium isotope work, analytical protocols used in the laboratory follow Kemp et al. (2009b). The data reproducibility was evaluated by quality control analysis using the reference zircons Penglai, OGC, FC1 and Mud Tank zircon. The εHf values for analysed sample zircons were calculated using a 176Lu decay constant of 1.865 x 10-11 yr-1 (Scherer et al., 2001), and the chondritic values of Bouvier et al. (2008). The detailed methodology of oxygen and hafnium isotope analysis in zircon is available as Supplementary data.

4. Results 4.1. Petrology of granitic rocks The investigated rocks of the southern GAB show varying degrees of deformation, expressed by homogeneous to augen textures, as for the Pink Syenite and Uvá orthogneiss, respectively. The igneous rocks are mostly medium- to coarse-grained and typically present granoblastic to nematoblastic textures. Their characteristic assemblage is quartz-K-feldspar-plagioclase±biotite. Apatite, titanite, epidote, allanite and zircon are common accessory minerals.

Fig. 3. Hand sample and respective photomicrograph (cross-polarized) of Archean and Paleoproterozoic samples in the southern GAB. (A) Porphyroclastic texture of Pink Syenite; (B) K-feldspar porphyroclasts within a fine- grained matrix; (C) Coarse-grained Caiçara orthogneiss; (D) Saturate intergrain contacts of latter; (E) Mica-rich Paus de Choro granite; (F) Cross-polarized image of “E”; (G) Uvá orthogneiss shows intense deformation, and (H) typical augen texture.

Fig. 4. Representative hand sample and respective photomicrographs of Neoproterozoic rocks from the southern GAB. (A) Itapuranga I augen granite with rounded feldspar phenocrysts in green hornblende groundmass; (B) Nematoblastic fabric outlined by hornblende and biotite (plane-polarized light); (C) Subtle foliation marked by biotite grains and translucid portions of anhedral quartz and feldspar; (D) Allanite phenocrysts surrounded by epidote and biotite lath-like grains; (E) Hand sample of Serra Negra monzonite; (F) Photomicrograph under cross- polarized light displaying rhombic titanite in equilibrium with epidote; (G) Outcrop of the Rio Caiapó monzonite (-tonalite) showing common mafic enclaves, and (H) Photomicrograph under plane parallel polarizers of “G” with epidote corona surrounding allanite grain within groundmass composed of quartz, feldspar and biotite.

The TTG samples comprise medium- to coarse-grained and leucocratic orthogneisses composed of xenomorphic aggregates of biotite and rare hornblende as the main mafic minerals (Fig. 3C-H). These rocks display the effects of intense deformation, evidenced by irregular intergrain contacts, e.g., bulging, and augen texture. Typical accessory minerals include epidote, titanite, apatite and zircon. The medium- to coarse-grained Pink Syenite has bimodal grain size consisting of alkali-feldspar phenocrysts (≤ 1cm), with sub-phaneritic to fine-grained (~ 50 µm) matrix of xenomorphic plagioclase and quartz. Other constituents include trace carbonate, muscovite and zircon. Rare intergrowths of

quartz and feldspars form micrographic to granophyric textures (Fig. 3A). Although the homogeneous texture suggests the sample is undeformed, a regional foliation is locally defined by quartz and mica oriented sub-parallel to S2 (Chapter 3). The Itapuranga granites are medium- to coarse-grained with hornblende, and minor biotite, as the major mafic minerals. The Rio Caiapó and Serra Negra granites are fine- to medium-grained and have biotite as the main ferromagnesian mineral. Mafic enclaves are often present in the Rio Caiapó granite (Fig. 4G). Accessory minerals include epidote, titanite, allanite, apatite and zircon.

4.2. Whole-rock geochemistry 4.2.1. Major element geochemistry

Fig. 5. (A) Aluminium saturation index (Shand, 1943) classify Pink Syenite and Itapuranga as metaluminous and others as peraluminous; (B) TAS (Total alkali vs. silica; Middlemost, 1994) plots Archean rocks in granite and others in quartz-monzonite field; (C) (Na2O + K2O - CaO) vs. SiO2 (Frost et al., 2001) diagram categorize Archean samples as calc-alkalic, Rio Caiapó/ Serra Negra as alkali-calcic, and Itapuranga granites/ Pink Syenite as alkalic;

(D) All samples plot as magnesian in XFe vs. silica diagram, and (E) K2O vs. SiO2 diagram (Peccerillo and Taylor, 1976) shows tholeiite (Uvá), calc-alkaline (Caiçara, Paus de Choro and Serra Negra), high-K calc-alkaline (Rio Caiapó), and shoshonite (Itapuranga and Pink Syenite) compositions.

The aluminium saturation index (Shand, 1943) classify Pink Syenite and Itapuranga as metaluminous, with other granites as peraluminous (Fig. 5A). Based on their total alkali versus silica contents (Middlemost, 1994), the intrusions are categorized as granites and quartz monzonite (Fig. 5B). According to their alkaline elements, TTG orthogneisses are calc-alkalic, Rio Caiapó and Serra Negra granites are alkali-calcic, whereas Itapuranga granites and Pink Syenite are alkalic (Fig. 5C). The

K2O/Na2O ratios increase from TTG orthogneisses (0.33-0.60) to Paus de Choro granite (1.07), Pink Syenite (1.46), and culminate with overall high ratios of Neoproterozoic granites (1.27-1.72), except the Serra Negra granite that has relatively lower K2O/Na2O ratio (0.44).

In Harker diagrams (Fig. 6), a negative correlation is observed for MgO, CaO, TiO2, FeO and

P2O5 with increasing SiO2, whereas poor correlation is depicted for Al2O3, Na2O and K2O (Fig. 6). High

LaN/YbN ratios are shown by Neoproterozoic granites and the Caiçara orthogneiss (Table 3).

Fig. 6. Harker diagrams show a negative correlation of SiO2 with Al2O3, CaO, K2O, MgO, Na2O, P2O5, TiO2 and FeO2. Neoproterozoic and Archean samples display two major clusters, which is different from the distribution illustrated by the Pink Syenite.

Table 2. Geochemical data of granitic samples from the southern Goiás Archean Block

4.2.2. Trace and rare earth element geochemistry The trace element geochemistry of the plutonic rocks is presented as primitive mantle normalized

REE plots (Sun and McDonough, 1989; Fig. 7A, C), and primordial mantle normalized multi-element diagrams (Taylor and McLennan, 1985; Fig. 7B, D). All samples show relatively strong negative anomaly of high field strength elements (HFSE), such as Ta, Nb and Ti. The LILEs (Cs, Ba, Rb) are enriched in the Neoproterozoic granites when compared to the TTG orthogneisses.

Fig. 7. Diagrams illustrating primitive mantle normalized REE patterns (Sun and McDonough, 1989) from (A) Neoproterozoic, and (C) Archean-Paleoproterozoic samples. Primordial mantle normalized patterns (Taylor and McLennan, 1985) of (B) Neoproterozoic compared with (D) Archean-Paleoproterozoic samples.

Table 3. Geochemical data of granitic rocks of the southern GAB continued (Taylor and McLennan, 1985).

The total lower REE concentration from Pink Syenite (29.5 ppm) increases through TTG orthogneisses (~ 85.6 ppm), towards Neoproterozoic granites (~ 258.2 ppm). Similarly, the (La/Yb)N ratio increases from Pink Syenite (4.3), through TTG orthogneisses (7.2-27.7), to Neoproterozoic granites (29.1-45.5; Table 3). The Pink Syenite and Rio Caiapó granites display a negative Eu anomaly (Fig. 7B), whereas the Serra Negra and Itapuranga granites show anomalous high Ba (1700-2300 ppm) and Sr (1000-1600 ppm) contents (Table 2).

4.3. In-situ SHRIMP U-Pb zircon geochronology Results of the SHRIMP U-Pb zircon geochronology are summarised in Tables 4 and 5. Cathodoluminescence images of representative zircon grains and concordia diagrams are provided in Figs. 8 and 9.

4.3.1. Serra Negra granite The Serra Negra granite has euhedral to sub-euhedral, yellow to smoky zircon grains with prismatic to bipyramidal shapes. The crystals have variable sizes (100-500 µm), with an average length/width ratio of ~ 3:1. Oscillatory core to rim zoning is commonly revealed in CL images, with rare unzoned to weakly zoned cores. The grains often contain fractures and/or inclusions. Uranium contents range between 197-593 ppm, whereas thorium content varies between 24-130 ppm (Th/U average = 0.17). Grains with oscillatory zoning yield a weighted mean 206Pb/238U age of 527 ± 5 Ma (n =13; MSWD = 1.2), whereas a single patchy core yields a 207Pb/206Pb age of 2088 ± 10 Ma (Fig. 9D).

4.3.2. Rio Caiapó granite The Rio Caiapó granite is characterized by euhedral to sub-euhedral, prismatic to bipyramidal zircon grains. The crystals are 100-400 µm in size, with a length/width ratio of ~ 4:1. Zircons are characterized by weak oscillatory core and rim zoning. Uranium contents range between 253-5590 ppm, whereas thorium varies between 47-3216 ppm (average Th/U = 0.19). The grains are typically dark coloured and commonly display inclusions. From a total of 20 analyses, 6 were rejected both for high uranium content (> 1000 ppm; Table 4) and for highly discordant ages. This sample yield a weighted mean 206Pb/238U age of 627 ± 5 Ma (n=13; MSWD=1.3).

4.3.3. Itapuranga Suite The Itapuranga suite includes Itapuranga I hornblende granite and Itapuranga II biotite granite. Zircons from Itapuranga I are sub-euhedral to anhedral, stubby to elongate, with shapes varying from prismatic to bipyramidal. Grains are typically about 150 µm long (60-400 µm), with an average length/width ratio of ~ 3:1. Crystals are mostly colourless and contain fractures and inclusions. The zircons display a poorly developed core to rim oscillatory zoning revealed in CL images. Uranium contents vary between 242-1030 ppm, whereas thorium content varies between 183-549 ppm (average Th/U = 0.56). The core and rim analysed in two different grains (IT1-13 and 14; IT1-15 and 16), returned age differences of 30 and 10 Ma, respectively. Based on 15 concordant analyses, the Itapuranga I yield a weighted mean 206Pb/238U age of 612 ± 4 Ma (MSWD=1.4).

Table 4a. Results of SHRIMP U-Pb analyses of Neoproterozoic samples

The zircons from Itapuranga II granite are sub-euhedral to anhedral, prismatic to bipyramidal and mostly colourless to light brown. The grain size varies between 90-370 µm, with a length/width ratio of about 3:1. Fractures and inclusions are common. Poorly defined core and rim oscillatory zoning is revealed in CL images, with some zircons characterized by patchy cores. Uranium contents vary between 255-856 ppm, whereas thorium content range between 149-473 ppm (average Th/U = 0.51). Based on 10 concordant spot analyses, the Itapuranga II yield a weighted mean 206Pb/238U age of 613 ± 4 Ma (MSWD=0.9).

Table 4b. Results of SHRIMP U-Pb analyses of Neoproterozoic samples continued

Fig. 8. Cathodoluminescence (CL) images of representative zircon grains from the southern GAB with the indication of spot locations as indicated: green circles – U-Pb; yellow squares - δ18O, and dashed circles – Lu-Hf.

4.3.4. Pink Syenite Zircons from the Pink Syenite include two major populations subdivided according to their morphology, size, Th/U ratio, and features evident in CL images. The first group consists of dark, subhedral to prismatic grains (~ 150 µm) with faint to convolute oscillatory core to rim zoning and dark rims revealed in CL images. These grains commonly have inclusions and/or fractures, and an average Th/U ratio of 0.47. The second group is characterized by cloudy, stubby/round to subprismatic elongate grains with average sizes of 200µm and well-developed core to rim oscillatory zoning. Fractures and local inclusions are common and the average Th/U ratio of these grains is 0.57. Overall, both groups present a length/width ratio of ~ 2:1. In order to characterize each population, a large number of analyses were conducted during different sessions, which were compared by repeated analysis in two grains. Analyses in areas of the zircon associated with high common lead (204Pb) contents or patchy zoning due to mixed compositions were not considered in weighted mean ages. The first group of stubby subhedral zircon grains yield a mostly discordant 207Pb/206Pb weighted mean age of 2882 ± 42 Ma (MSWD=21, n=7), whereas the elongate prismatic grains from the second group yield a more precise 207Pb/206Pb weighted mean age of 2061 ± 14 Ma (MSWD=2.6, n=11).

Table 5a. Results of SHRIMP U-Pb analyses of Archean (-Paleoproterozoic) samples

Table 5b. Results of SHRIMP U-Pb analyses of Archean (-Paleoproterozoic) samples continued

Fig. 9. Concordia diagrams for all samples with inheritance indicated. (A) Itapuranga I yielded 612±4 Ma and Itapuranga II yielded 613±4 Ma weighted mean ages (B); (C) Rio Caiapó displays a weighted mean age of 627±5 Ma; (D) The emplacement of Serra Negra yielded a weight mean age of 527±5 Ma, with inheritance of 2088±10 Ma; (E) Paus de Choro returns a weighted mean age of 2821±3 Ma; (F) Pink Syenite emplacement at 2061±14 Ma and inheritance at 2882±42 Ma, based on mean weighted average and upper intercept, respectively; Crystallization of Uvá (G) and Caiçara (H) orthogneisses based on an upper intercept at 2870±59 Ma and 2844±38 Ma, respectively. Red colour represents spot analysis used for calculation of crystallization ages, blue represent rejected analysis and yellow shows inheritance.

4.3.5. Caiçara orthogneiss Zircon grains from the Caiçara orthogneiss are dark brown, subhedral to anhedral, with mostly elongate prismatic shapes. The grain size is about 150µm, and the length/width ratio is ~ 2.5:1. Uranium contents range between 202-3154 ppm, whereas thorium varies between 100-2402 ppm (average Th/U = 0.70). Fractures and inclusions are commonly associated with these zircons, which display core to rim oscillatory zoning in CL images. Overall high thorium and uranium contents associated with pervasive spongy texture indicate the grains were susceptible to radiogenic lead loss during younger thermal events (Pidgeon et al., 1966; Silver and Deutsch, 1963). Thus, zircons with high discordant U- Pb ages that favour isotopic disturbance are avoided for Hf analyses (cf. Patchett, 1983). A normal Concordia diagram, based on 10 mostly discordant analyses, display upper and lower intercepts at 2844 ± 38 Ma and 645 ± 89 Ma, respectively (MSWD=2.6). The same 10 analyses yield a mean 207Pb/206Pb age of 2814 ± 20 Ma (MSWD=1.2).

4.3.6. Uvá orthogneiss Zircons from the Uvá orthogneiss are anhedral, stubby to prismatic, and commonly contain fractures and inclusions. The average size of the grains is 100µm, with a length/width ratio of 2.5:1. Fine oscillatory core to rim zoning is revealed in CL images. Uranium contents vary from 104- 639 ppm, whereas thorium content range between 54-332 ppm (average Th/U = 0.46). An array on a concordia diagram based on five analyses (in four grains) shows upper and lower intercepts at 2870 ± 59 Ma and 636 ± 280 Ma, respectively (MSWD=7.9), with a weighted mean 207Pb/206Pb age of 2814 ± 20 Ma (MSWD=2.7).

4.3.7. Paus de Choro granite Zircon grains from the Paus de Choro granite are clear to cloudy, euhedral to subhedral, with elongate prismatic to bipyramidal shapes. The grains are typically yellow to pink and commonly contain fractures and/or inclusions. The zircon grains are about 200µm, with a length/width ratio of approximately 3:1. Uranium contents vary between 93-457 ppm, whereas thorium contents vary between 27-134 ppm (average Th/U = 0.42). Strongly discordant data was not considered for age calculation purposes. Based on 16 analyses, a mean 207Pb/206Pb age of 2821 ± 3 Ma (MSWD=1.7) was obtained for the Paus de Choro granite (Fig. 9E).

4.4. Hf and O isotopes through time The hafnium isotope compositions of zircon grains from granitic rocks in the southern GAB are plotted as a function of U-Pb ages in Fig. 10. The data define 4 major groups: (i) 2890-2820 Ma grains with εHf(t) between +2.4 and -0.7; (ii) ~ 2100 Ma grains with εHf(t) between -8.3 and -10.6; (iii) 630- 612 Ma grains with Hf(t) between -3.0 and -10.5, and (iv) ~ 530 Ma grains with εHf(t) between -18.6

and -20.6. In the first group, 11 grains with 207Pb/206Pb ages > 2840 Ma (out of 15 from Uvá and Caiçara orthogneisses) have nearly chondritic εHf(t) values from -0.7 to +1.1. Despite the slightly higher εHf(t) values (≤ +2.4) attributed to the 2821 Ma Paus de Choro granite, eight out of 13 grains have εHf(t) values between 0.0 and +1.1. The second group of 2060 Ma Pink Syenite have relatively unradiogenic εHf(t) values that are similar to those of the 627-612 Ma K-rich granites. The extremely unradiogenic εHf(t) values of zircons from the ~ 530 Ma Serra Negra granite plot on the extension of a trend roughly delineated by the first group.

Fig. 10. A plot of εHf versus crystallization and inheritance ages of magmatic zircons from southern GAB (error bar indicated at 2σ). The top right inset shows a comparison with available εNd data in the GAB. “New crust” and “CHUR” lines according to indicated references. The dashed grey line linking compositions of the Serra Negra granite and Uvá orthogneiss corresponds to the Hf isotope evolution of protoliths extracted from the mantle at ~ 2900-2800 Ma. The zircons that plot between this line may reflect a mix of ages or a contribution from sedimentary sources during reworking, e.g., Pink Syenite. Inset: εNd(t) diagram representing, from top to bottom, rocks from: Neoproterozoic rocks from the Goiás Magmatic Arc (A: Pimentel and Fuck, 1992), Paleoproterozoic rocks from the Crixás greenstone belt (B: Fortes et al., 2003), and Archean rocks from the Uvá TTG complex (C: Pimentel et al., 1996).

The zircon oxygen compositions are mostly heterogeneous within samples > 2060 Ma, with standard deviations comparable to that of the reference materials (0.2 < 2σ < 0.5 ‰; Supplementary Table A1). Oxygen isotopic compositions between 2890-2820 Ma form a scattered range from 2.1 to 7.2 ‰, with higher values that correlate with decreasing age, i.e. Paus de Choro (2.8-7.2 ‰). Abrupt rise to heavier oxygen isotopic compositions after 2821 Ma is comparable to the values measured for

2060 Ma Pink Syenite (5.8-7.2 ‰). Following considerable shift to high-δ18O zircons of 630-612 Ma K-rich granites (10.7-8.4 ‰), δ18O values rapidly decrease at 530 Ma Serra Negra granite (5.6-6.0 ‰).

18 18 Fig. 11. A plot of δ O versus emplacement age of granitic rocks in the southern GAB. The “δ Omantle” field of 5.3 ± 0.6 ‰ as defined by Valley et al. (1998). Inset shows comparison with available oxygen data in the São Francisco Craton - SFC (Albert, 2017), and worldwide (Valley et al., 2005; Kemp et al., 2006; Wang et al., 2008, 2011; Pietranik et al., 2008; Dhuime et al., 2012; Lancaster et al., 2014; Rapaela et al., 2014).

Table 6. Geochronology and isotopic dataset of granitic samples from the southern GAB

5. Discussion

This section presents an interpretation of the geochemistry, oxygen and radiogenic isotopes obtained for the granitic rocks and explores their significance in the context of the geological evolution of the southern GAB.

5.1. Geochemistry and igneous petrogenesis Archean samples of the southern GAB share similar geochemical signatures with typical TTGs

(Martin, 1994; Martin and Moyen, 2002), i.e. low K2O and HREE contents, high Al2O3, Na2O, and SiO2 contents. However, they have lower CaO and Rb than most TTGs (Moyen, 2011). Older TTG melts show tholeiitic affinity (i.e. Uvá) followed by major calc-alkaline intrusions (Caiçara, Paus de Choro and Serra Negra), revealing a conspicuous connection between TTG intrusions and the Serra Negra granite. A shoshonitic affinity discloses Pink Syenite and Itapuranga granites, whereas data for the Rio Caiapó plot in the high-K calc-alkaline field. The peraluminous character of Archean samples, Rio Caiapó and Serra Negra granites support the formation of other Al-minerals besides feldspar e.g. muscovite/ biotite (Frost et al., 2001).

The negative correlation of SiO2 with P2O5 indicate fractionation of apatite, whereas the decrease of SiO2 with increasing MgO and TiO2 suggests fractionation of biotite, amphibole and/or magnetite. Negative Ta, Nb and Ti anomalies in primordial mantle-normalized diagrams characteristic of all samples indicate the fractionation of rutile and/or titanite. Despite relative scatter of incompatible versus compatible elements, the sub-horizontal trend of TTGs, Pink Syenite and Serra Negra granite suggests partial melting was an important mechanism of their differentiation (Cocherie, 1986; Martin, 1986, 1994). The high content of incompatible over compatible elements suggests K-rich granites originated predominantly by fractional crystallization (Fig. A3.1; Supplementary material). The involvement of TTG-derived melts to produce Archean K-rich magmas is envisaged by previous studies (Friend and Nutman, 2005; Champion and Smithies, 2007). Geochemical similarities of the Paus de Choro granite and the Uvá orthogneiss favour a cogenetic relation between both samples.

Low YbN of TTG samples and Pink Syenite indicate residual garnet and/or hornblende in their source (Moyen and Martin, 2012). High Sr/Y ratios and low Y contents of Caiçara, Pink Syenite and Serra Negra support a garnet-bearing residuum associated with deep melting at convergent margins

(Condie, 2008; Martin, 1993). Despite the low Sr content (332 ppm), fractionated REE patterns (La/YbN = 27), low Y (2.5 ppm), and Yb (1 ppm) contents of Caiçara orthogneiss correspond to geochemical signatures characteristic of adakites (Drummond and Defant, 1990). However, since the term reflects Cenozoic magmatism associated with subduction of young (≤ 25Ma) oceanic crust, this nomenclature is considered misleading for Archean magmatism. Late K-rich magmas derive from low-degree melts of mantle previously metasomatized by hydrous fluids (Eklund et al., 1998; Scarrow et al., 2009; Lu et al., 2015). LILE enrichment relative to HFSE depletion of Pink Syenite supports the contribution of an enriched source. The low REE contents and (La/Yb)N ratio of this sample indicate a low degree of partial melting involving garnet-bearing

residue (Arth and Hanson, 1975; Arth and Barker, 1976). The Eu anomaly illustrated by Pink Syenite

(Fig. 7C) is consistent with fractionation of plagioclase. Moreover, very high K2O contents supported by K-feldspar phenocrysts with zircon inclusions and porphyritic texture, suggest K-feldspar accumulation was likely an additional process for Pink Syenite genesis. Equivalent post-collision high- K plutonism is produced through extensional relaxation incited by slab break off and/or delamination of the mafic lower crust during subduction in the Phanerozoic (e.g. Tibet, Lu et al., 2015). Delamination takes place after significant thickening (e.g. Fowler and Rollinson, 2012), whereas extension promotes extensive melting and input of mantle-derived fluids into a thick crust. K-rich granites are characterized by strong LILE enrichment (Cs, Rb, Th, Sr) indicative of crustal influence (Brown et al., 1984; Hawkesworth and Kemp, 2006), HFSE depletion (Ta, Nb, Hf, Y) and negative Ti anomaly typical of arc-related continental margins (Hawkesworth et al., 1993). Compared to other Neoproterozoic samples, the low (La/Yb)N ratios of Itapuranga granites likely reflect the presence of LREE-rich minerals e.g. allanite that led to the depletion of these elements in the melt. The contribution of an ultramafic component via partial melting of an intermediate to lower mafic continental crust and/or an enriched source could explain the abundance of transition metals, e.g. Ni, Cr, Co in these rocks.

5.2. Geochronological framework and significance of inherited ages Geochronological results provide evidence for three major peaks of igneous activity in the southern GAB. Older magmatism is related to the formation of Archean TTGs. Emplacement of the 2870 ± 59 Ma Uvá orthogneiss is analogous to 2876 ± 11 Ma zircon U-Pb LA-ICP-MS age obtained for tonalite of the Rio do Vermelho domain (Pimentel et al., 2003a). Widespread evidence of isotopic disturbance e.g. spongy microstructures, high Th/U ratios and age discordance diagnostic of ancient lead loss (Nemchin et al., 2016) explain the variability of 207Pb/206Pb zircon ages reported for the Caiçara orthogneiss. The crystallization of the latter, denoted by 2844 ± 38 Ma upper intercept, is comparable to 2818 ± 14 Ma zircon LA-ICP-MS U-Pb age of granodiorite (Beguelli Jr., 2012). Regardless of the apparent lack of older sources recognized herein, ≤ 3140 Ma crystallization ages presented for Caiçara igneous rocks by Beguelli, Jr. (2012) support the contribution of an earlier Mesoarchean crust producing this TTG. The lower intercept at ~ 640 Ma acquired for Archean TTG samples is interpreted as a result of thermal effects during Neoproterozoic deformation. The paucity of detailed studies and reliable geochronology led to the still unclear characterization of magmatism related to the 2821 ± 3 Ma Paus de Choro granite. However, the spatial distribution, compatible geochemical and isotopic signatures with the Uvá orthogneiss support a cogenetic relationship between both magmatic events. Emplacement of 2061 ± 14 Ma Pink Syenite intrusion in the FGB is contemporaneous with Au- related Paleoproterozoic deformation event in the GAB. Analogous magmatism is illustrated in the northern GAB by: (i) 2146 ± 2 Ma Posselândia diorite intrusive in the Hidrolina TTG (Jost et al., 1993); (ii) 2145 ± 12 Ma syn-tectonic granite intrusive in the Pilar de Goiás greenstone belt (Queiroz, 2000),

and (iii) 2170 ± 12 Ma mafic dikes crosscutting gold mineralization in the Crixás greenstone belt (Jost et al., 2010). The 2880 Ma inheritance obtained by Pink Syenite is comparable with Archean Uvá and Caiçara TTGs. Ages yielded by the Rio Caiapó granite (627 ± 5 Ma) concur with 587 ± 17 Ma whole-rock Rb- Sr age (Pimentel et al., 1996), and are at least 100 m.y older than the 522 Ma biotite K-Ar cooling age (Hasui and Almeida, 1970). The ~ 40 Ma time gap between the zircon U-Pb SHRIMP and Rb-Sr isochron from Pimentel et al. (1996) can be attributed to the use of distinct methodologies and/or sample location. Crystallization ages (612-613 ± 4 Ma) and zircon inheritance (1409 ± 13 Ma) obtained for the Itapuranga granites agree with previous 624 ± 10 Ma igneous and 1700-1400 Ma inherited zircon U-Pb SHRIMP ages (Pimentel et al., 2003b). Their synchronous emplacement with regional peak metamorphism and their spatial distribution along E-W trending thrust faults formed during the closure of the BFB endorse the syn-collisional setting argued by Pimentel et al. (2003b). This is compatible with other syn- to post-collisional magmatism recorded in the South American Platform during the Brasiliano/Pan-African orogeny, e.g. Dom Feliciano belt, in Uruguay (Lara et al., 2017). The age of younger intrusive event recorded by the 527 ± 5 Ma Serra Negra granite is consistent with the whole-rock Rb-Sr age of 508 ± 24 Ma (Pimentel et al., 1994, 1996), and the zircon U-Pb LA- ICP-MS age of 506 ± 4 Ma (Marques, 2017). The 2088 ± 10 Ma inheritance revealed by Serra Negra sample correlates with Pink Syenite emplacement. Previous negative εNd values and 2680 ± 20 Ma inheritance presented by Marques (2017) reconcile with 2880 Ma Pink Syenite inheritance and favour an Archean precursor for the genesis of the Serra Negra magmatism.

5.3. Zircon Hf-O isotopes and magma sources Hafnium and oxygen compositions in igneous zircons from terranes of the southern GAB define a remarkable progression that implies a partially linked crustal evolutionary history. The Hf-O isotope trends define four major groups at: (i) 2890-2820 Ma, (ii) ~ 2060 Ma, (iii) 630-610 Ma, and (iv) 530 Ma. The significance of these groups and the transitions between them are explored below. Mostly near chondritic εHf values of 2870-2840 Ma magmatic zircons from the Uvá and Caiçara orthogneisses highlight major mantle-derived input. The Paus de Choro initial Hf composition (εHf(t) > 0) indicate derivation by melting of juvenile crust recently extracted from the mantle (e.g., Amelin et al., 1999), which attests to diminishing crustal reworking over time. In contrast, unradiogenic εHf values of > 2880 Ma Pink Syenite inherited zircons suggest reworking of an older crust, but the limited dataset provides no insight into the nature of that crust. The transition from unradiogenic to more radiogenic Hf values can be promoted by remelting of an old crust by lower crustal delamination (e.g., Johnson et al., 2014), an increased mantle-derived input or progression of the mantle heat source into newly- formed juvenile crust. Recycling of ancient crust disagrees with the unradiogenic signatures of Pink

Syenite Archean inheritance and negligible record of synchronous magmatism imply this transition may be attributed to the progression of mantle heat source into juvenile crust. Superchondritic to near chondritic Hf isotopic evolution across 2890-2820 Ma reveals at least 70 m.y. of mantle-derived magmatism by extraction of melts from an analogous depleted reservoir. The nearly identical Hf isotope arrays of Archean rocks favour the consanguinity of these terranes. In turn, comparable U-Pb magmatic ages argue for a shared Mesoarchean evolution. Neodymium < 3280 Ma model ages for Uvá orthogneiss (Jost et al., 2013) and hafnium ≤ 3550 Ma model ages for Caiçara orthogneiss (Beguelli Jr., 2012) suggest a Paleoarchean heritage of these terranes. The older crust- forming event proposed by Pimentel et al. (2003a) based on ≤ 3500 Ma model ages of low Sm/Nd ratio samples is contentious. This is because Sm-Nd model ages are only applicable to estimate the timing of crust differentiation represented by single extraction (Arndt and Goldstein, 1987), which conflicts with dominantly negative εNd(T) values (Pimentel et al., 2003a) obtained for these rocks. In addition, alteration of accessory minerals controlling the Nd content (e.g. monazite, allanite, apatite) can result in isotopic re-equilibrium (Hammerli et al., 2014; Bauer et al., 2015) and disturbance of the Sm-Nd systematics is expected given the poly-metamorphic history of the GAB (e.g. Vervoort and Kemp, 2016). Accordingly, heterogeneous Hf isotope composition and discordance on concordia diagrams of the Caiçara sample may be due to zircon overgrowths formed during younger metamorphic episodes (e.g., Patchett, 1983; Amelin et al., 1999). Another point that emerges from the obtained Hf isotope dataset is the lack of evidence for collision pinpointed by the isotopic signature of Archean intrusions. The crustal thickening and progressive reworking of older components promoted during collisional settings are manifested by unradiogenic Hf signatures (Kemp et al., 2009b; DeCelles et al., 2009; Collins et al., 2011). Nonetheless, geochemical and isotopic signatures of the Paus de Choro granite are compatible with the trend defined by magmatic zircon population of the Uvá orthogneiss (Figs. A3.1 and A3.2; Supplementary material), which suggests similar sources and/or processes involved in their genesis. Accordingly, the Paus de Choro granite is developed on the margin of the Uvá terrane rather than as an exotic terrane. Also, considering the stratigraphic integrity between the Faina and Goiás greenstone belts that separate the Uvá and Caiçara TTG terranes in the south (Resende et al., 1998), the Paus de Choro granite is presumably linked with the Uvá terrane. Oxygen isotopic compositions of 2890-2820 Ma rocks are mostly in agreement with values obtained for worldwide Archean igneous zircons (5.0 < δ18O < 7.4 ‰; Valley et al., 2005) that are interpreted as typical of magmas in equilibrium with the mantle. The low δ18O values associated with high discordance and Th/U ratios largely accredited to the Caiçara orthogneiss suggest some degree of isotopic disturbance (Fig. A3.3; Supplementary material). Radiogenic lead loss due to isotopic disturbance can be promoted by hydrothermal alteration and/or metamorphism (Silver and Dautsch, 1963; Pidgeon et al., 1966; Timms et al., 2006; Kemp and Hawkesworth, 2014). The shift to unradiogenic Hf values recorded by 2064 Ma magmatism predicts the prevalence of

crustal reworking from initially more radiogenic Archean values and suggests continental collision (e.g., Kemp et al., 2009b; DeCelles et al., 2009) for the Pink Syenite crystallization. Similarly, early collision and minor reworking predict an unradiogenic signature of ~ 2880 Ma Pink Syenite inheritance. The low δ18O values of inherited grains (~ 4‰) may indicate high-temperature alteration of the source to the magmas from which these zircons crystallized. Higher oxygen ratios (6-7‰) of Pink Syenite igneous grains imply minor supracrustal and/or sedimentary contribution. The broad range of unradiogenic Hf signatures (-3.0 to -10.5) encapsulated by 630-610 Ma K- rich magmatism suggests melting of an old crust. This agrees with negative εNd (-5.1 to -5.7) and Mesoproterozoic model ages (~ 1440 Ma) of Itapuranga granites (Pimentel et al., 2003b), and with mixed εNd (-4.2 to +2.1) and Mesoproterozoic model ages (~ 1100 Ma) obtained for the Rio Caiapó granite (Pimentel and Fuck, 1996). All zircons with δ18O > 8.0 ‰ are confined to the 630-310 Ma age group. High δ18O ratios are diagnostic for the contribution of a component (likely sedimentary) formed by low-temperature processes (e.g. Kemp et al., 2006). Thus, anataxis of metasedimentary crust is inferred from the elevated oxygen composition of this group. Tectonic transport of sedimentary material into depth during granite generation is also registered in other collisional events along convergent margins, e.g. Lachlan and Ross-Delamerian orogens (Kemp et al., 2009). The strongly unradiogenic Hf values of the 530 Ma Serra Negra granite define the predominance of crustal reworking following a linear decrease from initial mostly near chondritic Archean signatures.

Derivation from an older sialic crust is reinforced by negative εNd (-3.0 to -4.0), 1100-1900 Ma TDM model ages (Pimentel and Fuck, 1996; Guimarães et al., 2012), and ~ 2680 Ma inheritance (Marques, 2017). Accordingly, comparable chemistry and isotopic signatures with the Uvá orthogneiss evoke shared sources and/or similar mechanisms forming the Serra Negra granite (Figs. A3.1, A3.2; Supplementary material). A regression line linking εHf isotope compositions of these samples highlights identical isotopic evolution from an Archean protolith. The linear decrease from ~ 2870 Ma mostly near chondritic signatures implies successively younger magmas likely exploited the same structural conduits as older batches. This challenges the heterogeneity attributed to disparate sources and growth paths across terranes of the GAB (Jost and Fortes, 2001; Jost et al., 2001, 2014).

5.4. Contentious Archean crustal growth A major dispute concerns the presence, or lack, of plate tectonic processes during the Archean. Longstanding debate on the formation of Archean crust has been polarized between: (i) subduction processes dominated by horizontal accretion analogous to modern convergent plate margins (Blewett, 2002; Condie, 2005), and (ii) non-uniformitarian models (e.g. sagduction; Goodwin, 1980) involving plume-related processes dominated by vertical accretion (Stern, 2008; Bédard, 2006, 2017). Crustal growth is mostly expressed by tectonic accretion and juvenile magmatism along subduction zones in the Phanerozoic (Şengör et al., 1993), but the geodynamic and petrogenetic processes involved in Archean crust generation are still unclear (Foley et al., 2002; Rapp et al., 2003; Hawkesworth et al.,

2010). Early mantle depletion characteristic for Nd/Hf isotopic compositions from rocks of distinct Archean cratons (Hawkesworth et al., 2010) is compatible with high Nb/Th and Nb/U ratios of Neoarchean komatiites and tholeiitic basalt (Sylvester et al., 1997; Kerrich and Xie, 2002). Archean continental crust is hence widely accepted to derive from reworking of early oceanic crust (Campbell and Taylor, 1983; Taylor and McLennan, 1995; Rollinson, 2010). Archean TTG generation may occur by partial melting of: (i) oceanic plateau crust, (ii) subducted oceanic crust derived from spreading centers, and (iii) basalt at the base of oceanic arcs (Polat, 2012). This implies in a non-uniformitarian model for the first option and Phanerozoic-like tectonic processes for the last two options. Depletion of incompatible elements in tholeiitic greenstone belt lavas (Tarduno et al. 1991) attests for significant mantle contribution characteristic of early plume-related oceanic plateau growth (Reymer and Schubert 1986; Hill, 1993; Choukroune et al., 1997). Despite episodes of plume-related activity being particularly pronounced in the Archean (Stein and Hoffmann, 1994) the lifespan of this magmatism is too short to account for ≤ 160Ma volcanism of the Faina/ Goiás basins. Spreading center basalts are characterized by depletion of incompatible elements that precludes these being the sources of TTG generation (Polat, 2012). Accordingly, ‘arc-like’ geochemistry of Archean greenstone belts is largely used as evidence for early Earth subduction and many workers suggest that subduction became the dominant mechanism of lithospheric recycling at some time in the late Archean (Hawkesworth and Brown, 2008; Korenaga, 2008). Arc-like chemistry denotes LILE-rich, HFSE-poor mantle derived from dehydration and partial melting of underlying slab and enrichment of overlying mantle wedge during subduction (e.g. Jenner et al., 2009). Yet, incorporation of hydrated and enriched supracrustal material like sediments into the upper mantle can also produce ‘arc-like’ signature (Van Hunen and Moyen, 2012). Therefore, uniformitarian plate models are difficult to reconcile with early Archean terranes (e.g. Van Kranendonk et al., 2004). Alternative scenarios to subduction invoke diapirism as an important mechanism in Hadean and early Archean tectonics (Robin and Bailey, 2009; Johnson et al., 2014; Sizova et al., 2015). Partial melting of mafic and ultramafic inputs to the crust after multiple cycles of plume activity would generate TTGs by dominant vertical tectonics or diapirism (Bédard, 2006). The dichotomy between these tectonic styles seems ultimately misleading as they are not mutually exclusive (e.g. Polat et al., 1998; Percival, 2007; Van Kranendonk et al., 2014). Evidence for coeval vertical and horizontal tectonics is recognized in the Archean record (Van Kranendonk, 2010; Bédard et al., 2013; Lin et al., 2013; Harris and Bedard, 2014). The next sections contextualize these two paradigms in the Archean record of the GAB.

5.4.1. Plume-related oceanic plateaux model Intercalation of tholeiitic basalts with komatiites of several Archean oceanic plateaux is identical to the typical plume-related tholeiite-komatiite association (e.g. Herzberg, 1992; Arndt, 1994; Condie,

1994; Polat, 2009). Consonant up-ward younging stratigraphy and geochemistry with Archean greenstone belts make oceanic plateaux an appealing model (Arndt, 1999; Barnes and Van Kanendonk, 2014; Condie, 2005). Unlike MORB, plateau-basalts are commonly mentioned as TTG sources (Rudnick, 1995; Kemp and Hawkesworth, 2003; Smithies et al., 2009; Martin et al., 2014). Generation of bimodal geochemistry due to infracrustal melting at the base of an over-thickened plateau can produce TTG magmas (e.g. Willbold et al., 2009). Models proposed for Mesoarchean granite- greenstone petrogenesis often argue for oceanic plateau (Hollings and Kerrich, 1999; Fralick et al., 2008), plume-driven continental rift (Tomlinson et al., 1998), and plume-arc (Hollings and Wyman, 1999). Still, plateau plume-related models have a fairly short lifespan. Plateau-derived crust is estimated to form in pulses of tens of m.y., e.g. ~30 m.y. Ontong Java plateau (Kerr, 2013) rather than enduring (Abbott and Isley, 2002). Juvenile sources are inferred from mostly chondritic Hf composition of Archean samples. Generation of TTG in oceanic plateaux can be promoted by: (i) remelting of pre-existing lower crust (Smithies et al., 2009), (ii) density inversion (Bedard, 2006), or (iii) partial melting of underplated basaltic melts (Moyen and Stevens, 2006). Plume-driven infracrustal differentiation (Smithies et al., 2009) would cause progressive reworking of older material alike melting of a thick basaltic pile during density inversion (Bedard, 2006). Increasing contrast between newly formed and reworked crust for the two first premises is expected to cause an expansion of isotopic compositions. This contradicts from the isotopic record of Archean intrusions. Instead, more super chondritic compositions with decreasing age may subscribe remelting of older material at the base of the stratigraphy (Moyen and Stevens, 2006), e.g. forming the Paus de Choro granite. Mesoarchean isotopic data in the GAB is not consistent with transient plume magmatism. Long- lived, extensive and incessant plume(s) activity would be required for at least 160 m.y. of near- continuous magmatism recorded in the southern GAB. Plume-related magmatism, however, could explain the initial melting of older crust at ~ 2880 Ma, illustrated by Pink Syenite inheritance. Subsequent crustal thinning would cause a progression to more radiogenic compositions representing 2820 Ma Paus de Choro magmatism except the 60 m.y. timescale is too long for a plume. At last, linear hot-spot trails are also inconsistent with the isotopic-time record.

5.4.2. Subduction-related model Subduction zones create enduring, but cyclic, linear successions of arc-back arc magmatism (Stern, 2002). Sporadic magmatism produced by melting of an older crust is enhanced during initiation of extension after compressive thickening (Collins and Richards, 2008) or compression (DeCelles et al., 2009). Correspondingly, the fluctuation of radiogenic to unradiogenic Hf/Nd signatures in convergent margin Phanerozoic magmatism reflects upper plate shifts from extension to compression (Ducea, 2001; Kemp et al., 2009b). Isotopic flare-ups ascribe crustal thickening and increased melting of back-arc sediments (Kemp et al., 2009b) during flat subduction caused by buoyant oceanic crust e.g.

plateau (Collins, 2002), or incorporation of lower continental crust in arc region by underthrusting during compression (DeCelles et al., 2009). Depending on their position relative to mantle convection cells, Phanerozoic orogens are proposed to be external (accretionary) or internal (collisional; Collins et al., 2011). The Hf/Nd isotopic distribution in external orogens, whether advancing (Andes; Bahlburg et al., 2009) or retreating (Tasmanides; Kemp et al., 2009b), converge to radiogenic values, whereas Hf isotopic trend in internal orogens (Himalayas) fan out with time and are ultimately considered part of Wilson cycles (Collins et al., 2011). Proterozoic Wilson cycles delineate inverted U shapes in Hf isotope trends (Smits et al., 2014; Partin and Sylvester, 2016) postulated as diagnostic of plate tectonic processes in the Archean e.g. West Greenland (Naeraa et al., 2012), Dharwar (Lancaster et al., 2015). This paradigm concurs with chemical similarities of ~ 2960 Ma Faina/Goiás ultramafic rocks with boninites (Borges et al., 2017) attesting for the involvement of subduction zones in the Archean. Additional piece of evidence in favour of Archean subduction in the southern GAB is provided by the recognition of Nb-rich basalts and adakite-like rocks, e.g. Caiçara orthogneiss, akin to modern subduction environments (Drummond and Defant, 1990; Martin et al., 2005) in the Faina/Goiás greenstone belts (Borges et al., 2017;). Thus, subduction processes can afford for major features in the southern GAB isotopic record such as enduring juvenile Archean magmatism and the drastic isotopic shifts denoted by Pink Syenite, Itapuranga and Serra Negra intrusions. The swing to unradiogenic isotope compositions expressed by these rocks coincides with a change to compression (DeCelles et al., 2009; Kemp et al., 2009b) and/or collision in an internal orogeny (Collin et al., 2011), whereas super chondritic values in the late Archean record would reflect initial extension.

5.4.3. Implications for the geological evolution of the southern GAB The secular evolution recorded by investigated granitic rocks is presented in the geological context of the southern GAB and illustrated in Figure 15. Previous models for the petrogenesis of juvenile crust in the GAB attribute the generation of early oceanic crust to plume-arc interaction (Jost et al., 2019) and intraoceanic fore- to back-arc accretion (Borges et al., 2017). The first option, recently proposed for the Crixás greenstone belt in the northern GAB, is difficult to reconcile due to the need for plume-arc interaction in each greenstone belt (e.g. Bédard, 2013). Alternatively, basalt can be generated in back arcs (Fryer, 1981). High Nb/U and Nb/Th ratios prescribed for tholeiitic magmas of the Faina/Goiás greenstone belts (Borges et al., 2017) are inconsistent with the old crustal contribution presumed for a continental rift or flood basalts. Thus, Archean greenstone belts can be accommodated by deep-seated, mantle-derived, fore- to back-arc magmatism. Bimodal Phanerozoic back-arc rifts in continental settings are dominated by continent- derived sediments (e.g. Collins, 2002; Menzies et al., 2002). Crustal thickening after the closure of oceanic back-arc potentially led to supracrustal interaction and transient Hf/Nd isotopic dataset. Most U-Pb zircon ages of TTG rocks of the southern GAB are between 2810-3040 Ma (Beguelli

Jr., 2012; Jost et al., 2013). Crystallization ages may be subdivided into two-time intervals 2810-2880 Ma and 2920-3040 Ma, with a 40 Ma gap between the two suggesting at least two crust-forming events. Mesoarchean crustal growth is represented by juvenile magmatism producing 2870 Ma Uvá and ~ 2840 Ma Caiçara orthogneisses followed by more differentiated magmas originating the 2820 Ma Paus de Choro granite. Distinct sources for Archean magmatism include tholeiitic metabasalts (Uvá) and calc- alkaline magmas (Caiçara and Paus de Choro). In terms of differentiation mechanisms, partial melting is interpreted to dictate the genesis of the Uvá and Caiçara orthogneisses, whereas fractional crystallization is proposed for the Paus de Choro granite. The spatial distribution, geochemical and isotopic evidence suggest a co-genetic relation between the Uvá orthogneiss and the Paus de Choro granite. Initial positive εNd values from Borges et al. (2017) preclude continental crust contribution for ≥ 2920 Ma mafic-ultramafic rocks of the Faina/ Goiás greenstone belts. This agrees with typical LREE- rich, depleted mantle reservoir of Mesoarchean magmatism (Shirley and Henson, 1984). Positive εNd values can be attributed to the transient storage of subducted mafic-ultramafic oceanic crust (Chase and Patchett, 1988). Thus, ultramafic rocks are interpreted to result from hydrous melting of a depleted mantle in forearc setting similar to boninites (Borges et al., 2017). Fore-arc crustal growth by seafloor spreading appears to be contemporaneous with the onset of subduction (Bloomer et al., 1995), which supports initiation of subduction during early extension instead of shortening (Stern and Bloomer, 1992). Fore-arc extension of ~ 2960 Ma juvenile komatiitic and tholeiitic basalt crust (Fig. 12A) progressed to rift volcanism, subduction, arc formation, and ≤ 2920 Ma oceanic slab melting, with slightly crustal contaminated rocks marking the ca. 2830 Ma collision (Fig. 12B; Jost et al., 2013; Borges et al., 2017). Based on Menzies et al. (2002), the volcanic-dominated Goiás basin could illustrate an active rift, whereas the sediment-dominated Faina basin would represent a passive rift. Intercalation of felsic volcanoclastic e.g. calc-alkaline andesite, diorite and tonalite with a relative crustal contribution (Borges et al., 2017) is typical of back-arc basins, and to a minor extent accretionary prisms (Cawood et al., 2009). Mesoarchean greenstone belts of the southern GAB denote preferential preservation of back-arc basins hence the relative paucity of arc basalts may assign for their preferential destruction. Selective preservation of plume rather than arc origin for the Faina/ Goiás basalts can be explained by the buoyant nature of Archean oceanic plateau (Condie, 1997; Kerrich et al., 1999; Polat and Kerrich, 2001). TTG plutonism synchronous to greenstone belt volcanism (Borges et al., 2017) closely correlates with timing of felsic intrusions relative to greenstone belt formation in other Archean cratons (e.g. Superior Province; Corfu and Davis, 1992). After Archean growth and stabilization, quiescence in the GAB is evidenced by unconformities in the contacts of basal Archean volcanic and Paleoproterozoic sedimentary sequences of the Faina/Goiás and Crixás greenstone belts. The only record of magmatic activity during this time is registered by stocks and mafic dyke swarm recording crustal extension at ~

2400-2300 Ma recorded in the Archean TTG complexes (Fig. 12C; Corrêa da Costa et. al., 2006). Collectively, the geological settings, isotopic signatures, and geochemical features provide arguments for an intraoceanic setting in the southern Faina/Goiás basins. Contrasting paleogeography and depositional regime promoted deep to progressively shallower marine deposition at Goiás and two transgressive cycles of crescent depth platform sedimentation at Faina (Resende et al., 1998). Sm-Nd model ages reveal decrease from lower (3000 to 2800 Ma) to upper (2700 to 2600 Ma) sources for the Faina/Goiás sedimentary sequences (Resende et al., 1999) that agrees with more recent LA-ICP-MS zircon U-Pb provenance ages and Sm-Nd data of detrital rocks from Brant et al. (2015).

Fig. 12. Context of studied rocks along the geological evolution in the southern GAB. A) Opening of early oceanic basin during forearc extension; B) Subsequent back arc compression at late Archean and formation of TTG- derived melts (i.e. Paus de Choro granite); C) Regional crustal extension recorded by mafic dike swarm; D) Successive deposition of sedimentary rocks in platform settings; E) Reworking (and minor recycling) during Paleoproterozoic orogeny characterized by greenschist metamorphism, ductile deformation and Au mineralization; F) Widespread crustal recycling for syn-collision magmatism of K-rich granites in the Neoproterozoic Brasiliano (Pan-African) orogeny and reworking for post-collisional Serra Negra granite.

The stratigraphic, temporal and spatial correspondence of the Faina and Goiás sequences is interpreted to reflect two fore- to back-arc systems developed in a pull-apart basin (Resende et al., 1998; Borges et al., 2017). Dominantly dextral strike-slip displacement along early basin-accommodating lithospheric structures across southern GAB reconciles the stratigraphic record and the inferred merging of the basins (e.g. Faina fault between the Faina and Goiás greenstone belts; Resende et al., 1998).

Archean deformation is bracketed in the southern GAB at ~ 2830 Ma (Jost et al., 2013), and in the northern GAB at ~ 2700 Ma (Queiroz et al., 2008). All greenstone belts in the GAB record < 2300 Ma sedimentation, but Archean sources are only registered at Crixás (Jost et al., 2010), Faina (Fig. 12D) and Goiás (Resende et al., 1999; Brant et al., 2015). Albeit not absolute, the timing of Au mineralization is constrained by late Pink Syenite intrusion in the Faina greenstone belt. If the intrusion crosscuts < 2300 Ma sedimentary rocks, deformation, metamorphism and mineralization likely occurred between 2064-2300 Ma (Fig. 12E). The uneven distribution of Paleoproterozoic magmatism derived from reworking in deep lithospheric boundaries e.g. accretionary orogens may be clarified by isotopic evidence. Significant crustal growth in accretionary orogens is promoted by the consecutive accretion of oceanic basins to active continental margins (Cawood et al., 2009; Kemp et al., 2009b). Yet, recycling of continental material can exceed juvenile crust formation (Stern and Scholl, 2010), causing a mix of juvenile arc magmas with subducted continental crust. As a result, the isotopic and geochemical variation of subduction-related magmatism in intra-oceanic arcs is widely accepted to reflect input from such mixed sources (Kuzmichev et al., 2005; Carpentier et al., 2009; Nebel et al., 2011). Accordingly, restricted crustal recycling during subduction support mixed sources with negligible supracrustal contribution and near-surface interaction for Pink Syenite intrusion. The trend to heavier δ18O composition in 2064 Ma Pink Syenite can be explained by reworking of extensive sedimentary record in the FGB. Isotope and geochemical features from this rock suggest the incorporation of sedimentary material at depth. Incorporation of sedimentary material during tectonic burial can occur by underthrusting (Collins and Richards, 2008) or subduction of ocean basin/trench material (Scholl and von Huene, 2009). Thus, subduction of oceanic crust in accretionary orogen confined between the Uvá and Caiçara continental blocks is proposed to generate the Pink Syenite intrusion during the Paleoproterozoic orogeny. Evidence supporting the genesis of Pink Syenite after reactivation of Archean structures (-sources) is corroborated by contrasting juvenile signature of 2880 Ma inherited and unradiogenic εHf signature of 2060 Ma magmatic zircons. Comparable features are recognized for arc-related granitic rocks associated with subduction processes in other areas affected by the Eburnean orogeny in the West African craton, e.g. the Baoulé-Mossi domain (Petersson et al., 2017). Geochronological spectrum of TTG crust (2700-3000 Ma) and greenstone belt formation (ca. 2900-2700) of the GAB show similarities with plutonic rocks of the Jequié block (2950-2700 Ma) and the Umburanas greenstone belt (2700 Ma), in the northern São Francisco craton (Teixeira et al., 2000; Barbosa and Sabaté, 2002). In the southern São Francisco craton, development of TTG crust (3500- 3000 Ma, 2950-2820 Ma, 2750-2680 Ma and 2600 Ma) and volcanism of the Rio das Velhas greenstone belt (2800-2750 Ma) are at least 100 Ma younger when compared to the Faina, Goiás and Crixás greenstone belts (Teixeira et al., 2000; Jost and Fortes, 2001; Fortes et al., 2003; Noce et al., 2005; Borges et al., 2017).

The unradiogenic Hf isotopic values and heavy oxygen compositions of K-rich granites favour a strong continental affinity and involvement of older continental material for their genesis. A greater crustal contribution is ascertained by high K and Th contents (Condie, 2005) of these intrusions. K-rich granites crystallization during peak of orogenic processes in the Brasiliano/Pan-African orogeny (Pimentel et al., 2003b), involved continental arc (Rio Caiapó) and continental subduction magmatism (Itapuranga granites). Relatively low temperatures and high-pressure magmatism in continental collisional setting, e.g. for the Itapuranga granites, designate (paleo-) subduction zones and/or high mountains (Araújo et al., 2014). The insurgence of such rocks in the Neoproterozoic is proposed to mark the onset of modern plate tectonics dominated by ‘cold’ subduction (Stern et al., 2005; Brown, 2008). This agrees with a sharp change to heavier 18O compositions at ~ 600 Ma, which corresponds to the largest known absolute rise in atmospheric pO2 (Canfield et al., 2007). Equivalent collisional zones with deep roots, less dense than mantle rocks, are formed due to deep subduction of felsic continental crust (Austrheim, 1991; Gilotti, 2013). If high Rb/Sr ratios point towards increased granitic contribution in the melt (Dhuime et al, 2012), contrasting Rb/Sr ratios documented for the 612 Ma Itapuranga I and 613 Ma Itapuranga II granites (~ 1.0 and ~ 0.2, respectively) endorse their emplaced within suture zone along the eastern margin of the GAB. However, the exact timing and spatial distribution marking the progression of collision- to subduction-dominated tectonics have not yet been constrained. Emplacement into sedimentary rocks and volcano-sedimentary sequences favoured contamination and/or remelting of sedimentary and volcanic rocks (Fig. 12F). Mesoproterozoic inheritance and collisional settings agree with their unradiogenic εHf signatures (Kemp et al., 2009a; DeCelles et al., 2009; Collins et al., 2011). General consensus assumes the ultimate creation of topography during continental collision that may have facilitated the contribution of supracrustal/sedimentary material to form K-rich granites, as inferred by their high δ18O values. Felsic magmatism synchronous to sedimentary sequences, high-P magmatism and protracted crustal-mantle evolution are trademarks of Phanerozoic accretionary orogens (Cawood, 2005; Kemp et al., 2009b). The intrusion of K-rich granites into sedimentary rocks (Rio Caiapó) and volcano-sedimentary sequences (Itapuranga) favour contamination and/or remelting of sediments (± volcanic rocks). Coeval thermal resetting evidenced by lower intercepts at ~ 640 Ma for the Uvá/ Caiçara orthogneisses is compatible with 625 ± 6 Ma leucogranite in the Caiçara TTG (Pimentel et al., 2003a). The link between supercontinent formation and crystallization age peaks (Hawkesworth et al., 2013) endorse the emplacement of K-rich granites during the Neoproterozoic amalgamation of Western Gondwana. Similar syn-collision magmatism in other areas affected by the Brasiliano/Pan-African orogeny includes the Dom Feliciano and Ribeira fold belts (Hasui, 2010; Bley et al., 2014). Strongly unradiogenic εHf and mantle-like δ18O values of subduction-related 530 Ma Serra Negra granite bear evidence for the contribution of an old felsic crust, as attested by Paleoproterozoic and Archean inheritances (Marques, 2017). The abrupt shift of Hf isotopic compositions from syn-

collisional K-rich granites to late Serra Negra magmatism likely echo geodynamic changes. Switch to strongly unradiogenic Hf isotope composition implies continental compression predating final stages of the Gondwana amalgamation. The evolution of isotopic trends in the southern GAB defines distinct stages of the Wilson cycle, including: (i) Archean craton formation by fore-arc continental extension, subduction, back-arc formation and late compression; (ii) Paleoproterozoic ocean opening, passive margin development, ocean closure and accretion of back-arc system, and (iii) Neoproterozoic continental collision, high-T magmatism, ‘cold’ subduction and arc accretion. The allochthonous terrane accretion model previously proposed for the geological evolution of the GAB is challenged based on evidence for an in situ crustal growth. This is illustrated by: (i) tectonic and stratigraphic correlation of Archean and Paleoproterozoic geological records of greenstone belts in the southern GAB, and (ii) isotopic testimony of similar deep-seated magma sources across the different terranes. The partially autochthonous history of the GAB is constrained by the southern Uvá (2870 Ma), Paus de Choro (2821 Ma) and Caiçara (2840 Ma) TTGs, the Faina/Goiás greenstone belts (2960-2100 Ma), post-collisional Pink Syenite (2064 Ma) and Serra Negra (527 Ma) intrusions. Disparate evolution of marginal syn K-rich (640-610 Ma) granites is coherent with the Goiás Magmatic Arc (e.g. Pimentel and Fuck, 1992). A key implication of this work predicts that a major component of the GAB formed via in-situ processes instead of by accretion of exotic terranes (e.g., Jost et al., 2010, 2014). The allochthonous nature of the GAB postulated by the former authors conflicts with the fairly consistent sources established for terranes across the southern GAB, which is analogous to that recorded in worldwide Archean cratons. Isotopic transition near the Meso- to Neoarchean boundary in different Archean cratons reflects the onset of subduction associated with mantle wedge melting (e.g. Dhuime et al., 2012; Hawkesworth et al., 2013). In view of the strong testimony for Neoarchean subduction (Polat and Kerrich, 2001; Percival et al., 2006), both arc and back-arc were likely important sites of crust generation, as claimed for the Phanerozoic (Cawood, 2005; Kemp et al., 2009b). The more mafic composition assumed for Hadean to early Archean crust (Dhuime et al., 2015; Kemp et al., 2010) induced recycling into the mantle (e.g. Hawkesworth et al., 2016). However, most > 3000 Ma crust assumed to be of felsic composition (e.g. Pimentel et al., 1996) was likely reworked and overprinted by Mesoarchean magmatism rather than destroyed. This is corroborated by the lack of > 3000 Ma exposure and unradiogenic inheritance. Considering Mesoarchean samples in this study comprise TTG, a basalt source is inferred (e.g. Barker and Arth, 1976; Foley et al., 2002). The vertical trend of these samples in the Hf-time plot (Fig. 10) suggests increasing depletion due to successive extraction from mantle source.

6. Conclusions Isotopic and elemental measurements in granitic rocks reflect a remarkable evolution across

terranes of the southern GAB. The tectonic evolution of the Archean-Paleoproterozoic southern GAB and Neoproterozoic western and eastern margins comprises opening and closing of ocean basins, accretion of autochthonous and allochthonous terranes, deformation and magmatism. Results of this investigation highlight the still uncertain origin of the GAB and allow reassessment of existing geodynamic models proposed for the area. Major outcomes of this work include: (1) Secular chemistry and zircon Hf-O isotopes highlight shifts in magma sources and depth of partial melting and differentiation. Four major domains are depicted by hafnium isotopes in zircons: (i) near chondritic 2870-2840 Ma Uvá and Caiçara TTGs, with minor superchondritic compositions at ~ 2820 Ma for the Paus de Choro granite; (ii) unradiogenic 2060 Ma Pink Syenite and ~630-610 Ma K- rich granites, and (iii) the strongly unradiogenic 530 Ma Serra Negra granite. Oxygen isotopes in zircons are: (i) overall mantle-like up to ~ 2820 Ma, when they rise to heavier compositions related to the onset of ‘hot’ subduction (and plate tectonics?), (ii) largely mantle-like at ~ 2060 Ma, (iii) δ18O-rich at ~ 600 Ma due to the onset of ‘cold’ subduction typical of modern-tectonics, until steep shift to mantle-like 530 Ma subduction-driven magmatism. (2) Disparate evolutionary paths favour a collage of terranes amalgamated by younger tectonic processes and imply allochthonous nature of the craton (Jost et al., 2010, 2014). However, isotopic consonance across terranes suggest generation of crust by successive magmatic additions, which predicts an autochthonous evolution for the southern GAB. (3) Long-lived magmatism supports crustal growth in the GAB persisted for over 70 m.y. The near-continuous Archean juvenile magmatism recorded in the southern GAB is inconsistent with models that invoke mantle plume driven magmatism. Instead, the most cogent model for Mesoarchean protracted magmatism agrees with relatively prolonged subduction processes. Therefore, model to reconcile Archean elemental and isotopic records include early intraplate rifting, greenstone belt formation, subduction and late collision. (4) The transition from Na-rich Archean TTGs to K-rich Neoproterozoic arc magmatism is represented by the Pink Syenite. Relative LILE enrichment, HFSE depletion and low REE content expressed by the latter favour contribution of an enriched source originated by a low degree of partial melting or by metasomatized slab-derived melts. Moreover, Pink Syenite generation by partial melting of such source suggest the presence of metasomatized mantle beneath the GAB in the Paleoproterozoic orogeny. Zircon sector zoned cores and oscillatory zoned rims of Pink Syenite are presumed to reflect high P differentiation and low P crystallization. This is reinforced by K-feldspar megacrysts with zircon inclusions and the porphyritic texture that typifies shallow crystallization. (5) After a > 1000 Ma hiatus, magmatic activity in the GAB resumed with the emplacement of the K-rich granites during the peak of collisional processes in West Gondwana (~ 660-600 Ma). Heavy O compositions and unradiogenic Hf signatures implicate widespread reworking of older crust and sedimentary/supracrustal material for their genesis. (6) Chemical and isotopic shifts from K-rich granites to the Serra Negra granite is associated with

a geodynamic change at the initial stages of Gondwana breakup. Based on chemical, geochronological and isotopic signatures, the 2870 Ma Uvá orthogneiss emerges as a precursor of the 527 Ma Serra Negra granite. (7) Pink Syenite is critical to understand the role of prominent Paleoproterozoic subduction- driven accretionary belts in the GAB. The syn- to late-collision nature of this sample assist to bracket the timing of orogenic Au mineralization in the Faina greenstone belt. Pink Syenite is analogous to Au- related syn-collisional magmatism in the northern GAB and other areas affected by the Paleoproterozoic orogeny during assembly of Atlantica, e.g. West African, Amazon and São Francisco cratons and Guiana Shield (Teixeira et al., 2007; Eglinger et al., 2017; Tedeschi, 2018). (8) Isotope-time similarities with highly metal endowed São Francisco and Amazon cratons have important implications in terms of the exploration of mineral systems. However, additional research is required to further investigate these links. (9) Complexities to unravel the geodynamic evolution of the GAB are partially attributed to the limited exposure. Additionally, the pervasive obliteration of early primary textural and structural relationships after intense polyphase ductile-brittle deformation in the GAB hampers geological interpretations. Distinct peak T and P conditions of TTG intrusions and greenstone volcano-sedimentary sequences are also still poorly investigated. Advances may lie in additional in-situ U-Pb and Hf-O isotopic data of accessory minerals from regional intrusions in order to better constrain the absolute timing of deformation, metamorphism and mineralization events.

Acknowledgements This research is part of PhD project completed by Jessica Bogossian and developed at the CET- UWA (Centre for Exploration Targeting, University of Western Australia). The authors would like to acknowledge the Australian Microscopy & Microanalysis Research Facility, AuScope, the Science and Industry Endowment Fund, and the State Government of Western Australian for contributing to the Ion Probe Facility at the Centre for Microscopy, Characterisation and Analysis at the University of Western Australia. This work was supported by the “Conselho Nacional de Desenvolvimento Científico e Tecnológico” - CNPq (Project No. 207220/ 2014-0) and the Society of Economic Geologists – SEG.

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Introduction to Chapter 4: “Gold systems in the Goiás Archean Block, central Brazil”

Chapter four provides an integration of the major outcomes obtained during this investigation into the context of the GAB. A summary of available data connecting the settings, distribution and characteristics of major gold deposits of the GAB allows a more comprehensive and current geological framework of the area. The dataset presented herein is used to evaluate paradigms regarding the contentious nature of this Archean crust and further assist exploration efforts and targeting strategies in the GAB. This chapter will be submitted to the “Journal of South American Earth Sciences” with the candidate J. Bogossian as the first author and S.G. Hagemann (Centre for Exploration Targeting, University of Western Australia) as co-author. Prevailing studies have focused in the characterization of gold occurrences and deposits enclosed in the three northern greenstone belts hosted in the Goiás Archean Block (Fortes et al., 2003; Petersen Jr., 2003; Tassinari et al., 2006; Marques et al., 2013). From the two southern greenstone belts, Faina has been extensively explored for gold, whereas only local mineralization is associated with the Goiás greenstone belt. Results presented refer to gold mineralization in the Faina greenstone belt, based on data collected from the Cascavel and Sertão deposits (Chapter 2). According to Hf studies on detrital zircons, 60-70% of the actual continental crust is estimated to form by the end of the Archean (Belusova et al., 2010; Dhuime et al., 2012). Yet, enduring controversy reflect the settings and geodynamics of Archean crustal growth (e.g. Condie and Aster, 2010; Kemp and Hawkesworth, 2003; Smithies et al., 2009). Thus, a better understanding of the petrogenesis of felsic to intermediate intrusive rocks that dominate exposed Archean crust (i.e. TTGs) is critical to help decipher the cratonic evolution of these terranes. Isotopes and chemistry of felsic magmatism across the Goiás Archean Block provide a window into crustal growth and geodynamics (Chapter 3). Results suggest the southern portion of the craton evolved mainly by magmatic additions rather than by accretion of exotic terranes (Jost and Fortes, 2001; Jost et al., 2014). The secular chemical and isotopic trends recorded in the GAB echo major shifts in Earth’s history. The overlap of ~ 2700 Ma, ~ 2100 Ma and ~ 600 Ma peaks with global magmatic and mineralization events enhances the importance of unravelling processes and geodynamics during the geological evolution of the GAB. Increasingly applied empirical paradigms link mineralization to tectonic settings in similar terranes worldwide (Kerrich et al., 2000; Mole et al., 2014; Champion and Huston, 2016). Regional isotopic zonation allows the definition of proxies for the distribution and formation of certain mineral systems, particularly gold, in the GAB. Identical isotopic patterns documented in other Archean cratons worldwide points to a global process.

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Chapter 4

Linking Au systems to crust-mantle evolution of Archean crust in central Brazil

Abstract The first discovery of gold in central Brazil dates to the early 18th century and since then it has been important to the socio-economic development of the region. One of the most important gold districts encompasses the Goiás Archean Block (GAB), which since the early 1970s has been the site of several world-class gold discoveries including the 7 Moz Crixás deposits. This contribution presents an overview of the regional geology related to the gold systems by combining previously published data from deposits in the northern Goiás Archean Block with new data from deposits in the south, particularly in the Faina greenstone belt. A detailed discussion on the nature and major controls of gold mineralization, with emphasis on orogenic greenstone-belt hosted deposits, is also presented. Moreover, this contribution provides new highlights into the crustal evolution of the area linking the distribution and formation of gold systems with the isotopic signatures of regional magmatism.

1. Introduction

* Goiás Archean Block hosts several economically important gold deposits. These include the world-class Serra Grande mine, in the Crixás greenstone belt (≤ 7 Moz Au; AngloGold), and deposits hosted in the Pilar de Goiás and Guarinos greenstone belts (6.5 Moz, Yamana Gold Inc.). Gold systems are hosted predominantly in Archean-Paleoproterozoic greenstone belts and rare Neoproterozoic intrusions (Fig. 2; Rodrigues, 2011; Jost et al., 2014). The deposits occur in a variety of mineralization styles, host rocks and structural settings (Jost et al., 2001, 2014), however, they all show strong structural control, CO2, K, Fe, S enrichment, similar fluid conditions and hydrothermal alteration characteristics typical of orogenic deposits (Goldfarb et al., 2001; Robert et al., 2005). Field relationship and geochronology evidence suggest these deposits formed during a regional hydrothermal event related to a Paleoproterozoic orogeny at ca. 2100 Ma (Jost and Queiroz, 2008).

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Fig. 1. Geologic setting of Brasília fold belt (modified after Pimentel et al., 2004) with the location of Fig. 2.

The majority of studies in the Goiás Archean Block have focused on the greenstone belts as they are the most important host for gold deposits in the region (Jost and Oliveira, 1991; Vargas, 1992; Resende and Jost, 1995; Jost et al., 2001, 2014). Limited work in the TTG complexes, including U-Pb geochronology Sm-Nd isotopic studies, and whole-rock geochemistry, have shown that the crustal architecture of the Goiás Archean Block is composed of discrete lithospheric blocks of varying age, composition and origin (Queiroz et al., 2000, 2008; Pimentel et al., 1996a, 2003a; Jost et al., 2005, 2013; Beguelli Jr., 2013). However, none of these studies has linked the isotopic characteristics of these blocks to the localization of mineral systems, as proposed for other cratonic areas such as the Yilgarn craton, in Western Australia (Mole et al., 2015). Available data throughout the Goiás Archean Block is highly variable, with most previous studies focused in the north (Queiroz et al., 2000, 2008; Beguelli Jr., 2012) with comparatively little work carried out in the south. The allochthonous nature and unknown origin of the Goiás Archean Block make the area an exciting case study to investigate relations between crustal evolution and distribution of mineral systems.

The purpose of this work is to provide an updated appraisal of existing literature on the regional geology of the main gold deposits and a qualitative and reasonably comprehensive tectono-magmatic data set for the GAB. The dataset is used to develop a more robust understanding of the crustal evolution

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of this cratonic terrane. The first part of the paper presents a synthesis of geological characteristics of orogenic gold mineralization in the GAB. The second part focuses on available chemical, geochronological and isotopic records in order to offer a framework of the crust-mantle evolution for the area. The long-term controversy concerning growth of Archean crust is assessed in an attempt to reconcile available models. The spatial framework of geological data is used to propose feasible links between the crustal evolution and mineral systems, particularly gold, in the GAB.

2. Tectonic Setting During the amalgamation of Neoproterozoic West Gondwana, the collision between Amazonian, São Francisco and Paranapanema cratons in the Brasiliano/Pan-African orogeny led to the formation of the Tocantins Province (Almeida et al., 1981; Pimentel and Fuck, 1992; Strieder and Suita, 1999). The province comprises the Brasília fold belt in the east and the Araguaia and Paraguai fold belts in the northeast and west, respectively (Pimentel et al., 2000a). The Brasília fold belt (BFB), which extends for more than 1000 km in the N-S direction, represents one of the most complete Neoproterozoic orogens of western Gondwana (Pimentel et al., 2016). The BFB is subdivided into NNE-SSW trending and NNW-SSE trending branches, which are separated by the WNW-ESE trending Pirineus syntaxis (Araújo Filho, 2000; Pimentel et al., 2000a, 2004). Several tectonic domains are distinguished in the Brasília fold belt (Fuck et al., 1994; Pimentel et al., 2000a, 2004; Dardenne, 2000; Valeriano et al., 2004, 2008). Its eastern part consists of thick passive margin sedimentary sequences deformed under low greenschist facies metamorphism (Fig. 1). In the northeast, the BFB is made up of a Paleoproterozoic granite-gneiss terrane containing minor volcano-sedimentary sequences known as the Natividade greenstone belt (Costa et al., 1976; Padilha, 1984), and Meso- to Neoproterozoic layered mafic-ultramafic complexes (Ferreira Filho et al., 2010). Its western core is composed of up by: (i) Neoproterozoic Anápolis-Itauçu granulite facies metamorphic core (Piuzana et al., 2003) and metasedimentary rocks from the Araxá Group (Pimentel et al., 2016); (ii) Archean granite-greenstone terrains of the Goiás Archean Block (Jost et al., 2014); (iii) Neoproterozoic Goiás Magmatic Arc, composed of calc-alkaline orthogneisses and volcano- sedimentary sequences (Pimentel and Fuck 1992; Pimentel et al., 2000b), as well as several syn- to post-orogenic granite intrusions (Pimentel and Fuck, 1994; Pimentel et al., 1996a).

3. Geology of the Goiás Archean Block (GAB)

Several names have been attributed to the area encompassing the Goiás Archean Block (GAB), e.g., ‘Crixás granite-greenstone belt terrane’ (Queiroz et al., 2000), ‘Archean terranes of Crixás-Goiás’ (Pimentel et al., 2000a), ‘Goiás Archean Nuclei’ (Jost et al., 2001), ‘Goiás-Crixás Archean Block’ (Pimentel et al., 2003a), ‘Goiás-Crixás Block’ (Delgado et al., 2003), ‘Archean terrain of central Brazil’ (Jost et al., 2010), ‘Goiás Archean Block’ (Jost et al., 2008, 2012), and ‘Archean-Paleoproterozoic

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terrane of central Brazil’ (Jost et al., 2014; Borges et al., 2017). The nomenclature adopted herein (e.g. Jost et al., 2008, 2012) refers to the terrane that consists of the only example of exposed Archean crust in the Tocantins Province (Almeida et al., 1981; Fuck et al., 1994; Pimentel et al., 2003a), which was amalgamated into the Brasília fold belt during the Brasiliano/Pan-African orogeny (Pimentel and Fuck, 1994; Pimentel et al., 2000a).

The GAB, formed by approximately 70% granite-gneiss with TTG association and 30% greenstone belts, records a chronological range from Archean to Paleoproterozoic rocks (Jost et al., 2008), with rare Neoproterozoic intrusions (Rodrigues, 2011; Pimentel et al., 2003a). It is bordered by the Goiás magmatic arc at north (Mara Rosa Arc) and south (Arenópolis Arc), whereas its northwestern and southwestern limits are defined by the Transbrasiliano Lineament and the Moiporá-Novo Brasil shear zone, respectively. The western margin comprises sedimentary rocks of the Araguaia basin, whereas the eastern margin consists of syn-tectonic granitic intrusions (e.g. Itapuranga granites; Pimentel et al., 2003b) and metasedimentary rocks of the Serra Dourada and Araxá groups (Fig. 1).

The polycyclic evolution of the GAB, inferred from studies concerning Archean TTG magmatism, led to its subdivision into northern and southern portions (Vargas, 1992; Pimentel et al., 2003; Queiroz et al., 2000, 2008). The northern portion was described by Queiroz et al. (2008) as consisting of two major magmatic stages: (i) juvenile, batholith-sized and polydeformed orthogneisses with compositions of tonalite, granodiorite and granite with SHRIMP U-Pb zircon ages between 2840- 2780 Ma and εNd +2.4 to -1.0, and (ii) crustal-derived, dike-like intrusion of granodiorite to granite gneisses with SHRIMP U-Pb zircon ages between ca. 2790-2700 Ma and a negative εNd (-2.2).

The southern GAB consists of the Uvá and Caiçara TTG complexes, separated by the Faina and Goiás greenstone belts (Fig. 2). Regional studies indicate the stabilization of the GAB occurred at ca. 2700 Ma, with Paleoproterozoic magmatism after crustal extension in the Siderian (i.e. 2500-2300 Ma; Corrêa da Costa, 2003) and closure of accretionary orogenic margins in the Rhyacian (i.e. 2300-2050 Ma; Jost and Queiroz, 2008; Jost et al., 2014).

3.1. TTG terranes

The TTG terranes of the GAB consist of six intensely deformed orthogneisses of variable composition and age that were first delineated by gamma-spectrometry (Blum et al., 2003). Compositionally, TTG orthogneisses vary from tonalite to granodiorite, with minor granite, charnockite, monzogranite and adakite (Jost et al., 2001; Beguelli Jr., 2012). Available geochronology shows that TTG terranes of the GAB were emplaced during several discrete episodes dating back to ~ 3140 Ma (Beguelli Jr., 2012), with the majority of preserved intrusions emplaced at ~ 2960-2840 Ma, 2845-2785 Ma, and 2790-2700 Ma (Queiroz et al., 2000, 2008; Pimentel et al., 2003a; Jost et al., 2005, 2013; Borges et al., 2017). The 2845-2785 Ma and 2960-2840 Ma episodes are the most widespread in

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the northern and southern GAB, respectively, and were broadly synchronous with Archean volcanism (Queiroz et al., 2000, 2008). The ~ 2790-2700 Ma episode is more common in the north and appears to be the last magmatic event in the GAB.

3.2. Greenstone belts There are five greenstone belts in the GAB (Fig. 4). In the north, three NNW-trending, subparallel inliers (ca. 40km length and 6km width) comprise, from west to east, the Crixás, Guarinos, and Pilar de Goiás greenstone belts. These greenstone belts were defined by Danni and Ribeiro (1978), and Sabóia (1979), with further description by Jost and Oliveira (1991). In the south, two NW-trending belts (ca. 60km length and 5km width) are represented, from west to east, by the Faina and Goiás greenstone belts. They both form a synclinorium offset by the NE-trending, dextral Faina fault (Resende et al., 1998; Jost et al., 2014). The preservation of primary structures e.g. upward younging in pelites indicates that the southwestern limb of the Faina synclinorium is inverted whereas the northeastern limb is upright (Resende et al., 1998). Based on the distribution of sedimentary sequences in the Faina and Goiás greenstone belts, their apparent right-lateral displacement, coupled with the contrasting sedimentation recorded in both sequences, the Faina Fault is inferred to represent a syn-sedimentary growth fault or a rift-related transform fault (Resende et al., 1998). Based on their geometry and lithostratigraphy, all five greenstone belts of the GAB were initially interpreted as synformal keels (Jost et al., 2001). However, this model is currently accepted only for the two southern Faina and Goiás greenstone belts (Resende et al., 1998), whereas the northern greenstone belts are interpreted as fold-thrust belts (Jost et al., 2001). Overall, the five greenstone belts display a similar sequence of lower komatiite (400-900 m) of peridotite to pyroxenite composition followed by tholeiitic basaltic flows (300-500 m; Jost et al. 2001). Primary volcanic features, such as spinifex and/or cumulatic textures are locally preserved in ultramafic rocks (Salles et al., 2014; Tomazzoli, 1985; Profumo, 1993). Basaltic flows are characterized by pillow to variolitic structures with local dolerite and gabbro dikes/sills (Jost et al., 2001). Minor BIF, gondite and/or chert lenses occur as thin layers intercalated with volcanic rocks (Resende et al., 1998; Jost et al., 2001). The transition from lower volcanic to upper sedimentary rocks is typically marked by a tectonic unconformity. In the Faina and Goiás greenstone belts, the volcanic sequences have a maximum depositional age between approximately 2960 and 2920 Ma (U-Pb LA-ICP-MS on zircon; Borges et al., 2017), whereas in Crixás, the maximum depositional age is indicated by Sm-Nd whole-rock isochron age of ca. 3000 Ma (Fortes et al., 2003). In the Guarinos greenstone belt, lower volcanic rocks are dated at 2180 Ma (Jost et al., 2012), whereas in Pilar de Goiás greenstone belt, volcanism is dated at 2165 Ma (U-Pb LA-ICP-MS on zircon, Jost et al., 2014).

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Fig. 2. Stratigraphic columns for the greenstone belts of the Goiás Archean Block organized by northern and southern terranes. Age compilation for greenstone belts according to Jost and Oliveira (1991), Resende et al. (1998), Jost et al. (2001, 2014), Brant et al. (2015), and Borges et al. (2017). Age references for the TTGs based on Queiroz (2000), Queiroz et al. (2008), Jost et al. (2005, 2013, 2014), and this work.

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Fig. 3. Geology of northern (A, B, C) and southern (D) greenstone belts from the Goiás Archean Block (GAB), with corresponding stratigraphy (E).

Unlike the lower mafic-ultramafic sequences, the overlying sedimentary units of the five greenstone belts indicate diverse depositional environments (Jost and Oliveira, 1991; Resende et al.,

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1998). At Crixás, initial sedimentation is represented by carbonaceous schists typical of euxinic environments, with oolitic to stromatolitic marbles of dolomitic composition that are unconformably overlain by pelites and greywackes (Theodoro, 1995). At Guarinos, lower chlorite-rich pelites containing clasts of basalt are laterally interlayered with basalts that are overlain by chemical sedimentary rocks (BIF, chert), conglomerate lenses and pelites (Resende and Jost, 1994; Jost, 1995; Jost et al., 2001). At Pilar de Goiás, calcsilicastic rocks, sandstone and dolomite are followed by greywackes (Resende and Jost, 1995).

The stratigraphy of the two southern greenstone belts was initially proposed by Danni et al. (1981) as a single greenstone belt composed of lower volcano-sedimentary sequence unconformably overlain by upper sedimentary rocks. Due to their contrasting characteristics, Teixeira (1981) subdivided this belt into the Faina and Goiás greenstone belts with lower komatiitic, basaltic and felsic volcanic rocks and upper siliciclastic, pelitic and chemical sedimentary rocks. The Goiás greenstone belt (previously known as Serra de Santa Rita) is characterized by lower carbonaceous-rich pelites followed by chert, BIF and dolomite unconformably overlain by turbidites (Resende et al., 1999). The Faina greenstone belt is characterized by two sequences consisting of lower conglomerate, quartzite and pelite overlain by chemically precipitated rocks (Resende et al., 1999).

The maximum deposition age of sedimentary rocks for all greenstone belts is constrained by microgabbroic dike swarm emplaced between 2500-2300 Ma (Corrêa da Costa, 2006) that crosscut Archean TTG complexes and volcanic rocks from the Goiás, Faina and Goiás belts but not the upper sedimentary rocks. This field relationship implies the sedimentary rocks must have been deposited following or during regional crustal extension after 2300 Ma. The end of sedimentation in all greenstone belts is indirectly constrained by carbon isotopes on upper dolomitic rocks (Fortes, 1996; Resende et al., 1998; Jost et al., 2008; Santos et al., 2008) based on chemostratigraphic markers that reflect the oxidation of the atmosphere in the Paleoproterozoic (2220-2060 Ma) known as the Jatuli event (Melezhik et al., 2007). Upper dolomites hosted in three northern greenstone belts and the lower sedimentary sequence in the southern greenstone belts show positive δ13C values (+10 to +14‰) that coupled with available geochronology are indicative the deposition of these dolomites occurred between 2220-2060 Ma (Santos et al., 2008; Jost and Queiroz, 2008), marking the first carbon anomaly of the Jatuli event following glaciation in the Huronian. Moreover, upper dolomites from the second sedimentary cycle of Faina and Goiás greenstone belts record mixed δ13C values (-0.66 to +0.66‰), which suggest their deposition at the end of Jatuli event, late in Rhyacian times and likely associated with extension in early Orosian (Jost et al., 2001).

3.3. Deformation History The GAB shows a polyphase evolution recorded by complex deformation patterns (Jost and Fortes, 2001). This complexity, together with the lack of absolute dating of regional structures has

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hindered a better understanding of the deformation history of the GAB, as indicated by previous studies published in the region (Jost and Ferreira Filho, 1987; Magalhães, 1991; Queiroz, 1995, 2000). Early structural interpretations are here reconsidered utilizing the most recent geochronology and isotopic geochemistry studies. At least three deformation events in the GAB were developed in the Archean, Paleoproterozoic and Neoproterozoic (Table A4.1. of appendices). The distinct deformation periods are recorded by titanite U-Pb SHRIMP dating of the Crixás Açu gneiss, which revealed distinct metamorphic episodes at 2711 Ma, 2011 Ma and 590 Ma (Queiroz et al., 1999, 2000). Following crustal extension represented by mafic dike swarms in the early Paleoproterozoic (Corrêa da Costa, 2003), additional deformation is recorded by faint crosscutting fabric in titanite with U-Pb SHRIMP age of 2011 ± 15Ma (Queiroz et al., 2000).

The Archean Dn event was responsible for regional amphibolite facies metamorphism and deformation of all granite-gneiss TTG complexes and lower volcanic sequences of the Crixás, Faina and Goiás greenstone belts (Fortes et al., 2003; Borges et al., 2017). Structures formed during Dn include tight to isoclinal folds, thrust faults and metamorphic foliation (Sn). The amphibolite facies metamorphism is expressed by talc-serpentine-chlorite±titanite and quartz-hornblende-biotite-epidote- titanite assemblages oriented parallel to Sn in the volcanic rocks and TTG granite-gneisses (Magalhães, 1991; Thomson, 1991). The contacts with adjacent granite-gneiss TTG complexes are commonly marked by deformed xenoliths of volcanic rocks in the southern belts which suggest the greenstone belts are allochthonous (Resende et al., 1999; Borges et al., 2017). In the northern GAB, an Archean metamorphic event, registered at ~ 2700, is coeval with the emplacement of younger orthogneisses (Queiroz et al., 2000), whereas in the southern GAB metamorphism is dated at ca. 2840 Ma (Jost et al.,

2013). The minimum age of the Dn event is provided by NE- and NW-trending mafic dikes dated by Rb-Sr (whole rock) and 40Ar/39Ar (hornblende) to 2490 ± 40 Ma which crosscut the TTG complexes and lower volcanic rocks in the Faina, Goiás and Crixás greenstone belts (Corrêa da Costa et al., 2006).

In the Paleoproterozoic, D1, D2 and D3 brittle-ductile events comprise regional deformation, metamorphism and gold mineralization hosted by greenstone belts in the GAB. Early developed structures include subtly S-dipping S1 foliation developed parallel to subparallel to the bedding (S0), axial planar to tight to isoclinal folds and the inversion of the stratigraphy in the greenstone belt sequences (Resende et al., 1999; Carvalho, 2005; Sobiesiak, 2011). Regional metamorphism comprises quartz-chlorite-muscovite-biotite±epidote assemblages oriented parallel to S1 foliation in the volcano-sedimentary rocks of the greenstone belts. This assemblage is compatible with greenschist to lower amphibolite facies and temperature ranges that roughly correspond to the brittle-ductile transition (Sibson, 1975).

Gold mineralization is associated with D2 thrust faults and, to a minor extent, D3 shear zones. The Au-related hydrothermal assemblages in all deposits overprint greenschist facies assemblages and therefore post-date peak metamorphism (Jost and Fortes, 2001; Chapter 2).

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Contrasting structural evolution is recorded in the northern and southern greenstone belts for D1,

D2 and D3 events during the Paleoproterozoic (Table 1). In the northern GAB, structures developed throughout these events include: (i) subtly ~ W-dipping S1 foliation (S0 transposed into S1), subtly W- dipping, NS-trending thrust faults, greenschist facies metamorphism and stratigraphic inversion; (ii) subtly S-dipping S2 foliation, planar axial to isoclinal F2 folds, local mm- to cm-wide shear zones, subtly

N-dipping, EW-trending thrust faults and stacking of the stratigraphy; (iii) subtly E-dipping S3 foliation, subtly E-dipping, NS-trending F3 folds and lateral thrust ramps (Magalhães, 1991; Petersen Jr., 2003; Massucato, 2004; Carvalho, 2005). In the southern GAB, Paleoproterozoic deformation is largely described based on observations in the Faina greenstone belt (Chapter 2). D1 is expressed by moderately

N-dipping S1 foliation (S0 transposed into S1), subtly W-dipping, NNW-trending F1 folds, stratigraphic inversion and greenschist facies metamorphism. D2 is characterized by subtly S-dipping S2 foliation, planar axial to tight to isoclinal F2 folds, gently W-plunging mineral lineation, gently S-dipping, EW- trending thrust faults that caused stacking of the stratigraphy. D3 is represented by moderately N-dipping

S3 foliation, NW-trending shear zones and subtly S-dipping thrust faults. Significant hydrothermal fluid flow along D2-related structures resulted in the development of Au-bearing quartz±carbonate veins with only localized hydrothermal fluid flow produced during D3 event.

The timing of Paleoproterozoic deformation encompassing D1 to D3 events can be stipulated by a series of intrusions between ~ 2170 and 2061 Ma. Early deformation is constrained by the Posselândia diorite, controlled by a NW-trending D2 shear zone that crosscuts the Hidrolina complex and has a zircon U-Pb SHRIMP age of 2146 ± 2 Ma (Jost et al., 1993). Jost et al. (2010) report mafic dikes crosscutting Au-related hydrothermal alteration in metasedimentary rocks from Crixás greenstone belt dated at 2170 ± 17 Ma (U-Pb LA-ICP-MS on magmatic zircons. Syntectonic granites within N-verging thrust faults in the Pilar de Goiás greenstone belt yield a zircon U-Pb SHRIMP age of 2145 ± 12 Ma (Jost et al., 1992, Queiroz, 2000). Syn- to late-tectonic Syenite intrusive in the Faina greenstone belt with zircon U-Pb SHRIMP age of ca. 2061 ± 14 Ma (Chapter 3).

The Neoproterozoic D4 event overprints, offsets and/or reactivates previous structures. Structural features developed in D4 include subtly S- to W-plunging crenulation lineations associated with thin- skinned, E-verging thrust to reverse faults and local NS-trending open folds (Araújo Filho and Kuyumjian, 2000). This event is attributed to the Brasiliano/Pan-African orogeny at 650-480 Ma

(Alkmin and Marshak, 1998). The timing of D4 is indirectly delimited by isotopic disturbance, e.g., Pb loss and lower concordia intercepts, at approximately 600 Ma in zircon U-Pb geochronology (Pimentel et al., 2003a; Queiroz et al., 2000, 2008; Chapter 3). The maximum age range of the D4 event is constrained by leucogranite intrusion in the Guarinos greenstone belt with U-Pb LA-ICP-MS age on hydrothermal zircon of 729 ± 15 Ma (Rodrigues, 2011). The minimum age range of this event can be ascertained by the late Serra Negra intrusion in the western border of the GAB with zircon U-Pb SHRIMP age of 527 ± 5 Ma (Chapter 3).

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Table 1. Deformation events proposed for the GAB

4. Gold in the Goiás Archean Block The mineral systems of the GAB are dominated by greenstone-hosted gold deposits (e.g., Mina III, Maria Lázara, Cachoeira do Ogó, Cascavel, Jordino), with poorly explored paleoplacer and intrusion-related deposits. Additional deposit types include minor Algoma and Superior-type Fe formations, sedimentary-hosted Mn, and Au±Cu VMS deposits are recognized but are not currently economic (Fig. 2; Jost et al., 2014). The Greenstone-hosted gold mineralization share similarities, e.g., strong structural controls, metal associations, hydrothermal alteration and mineralogy, which are characteristic of orogenic gold deposits worldwide (Goldfarb et al., 2005; Groves et al., 2000). This section presents a description of the structural, mineralogical, hydrothermal alteration, ore assemblages and fluid conditions of orogenic gold deposits and occurrences of the Goiás Archean Block.

4.1. Regional distribution In the 18th century, gold was first recognized in central Brazil in the region between the Faina and Goiás greenstone belts (Jost and Fortes, 2001). Since the 1970s, active mining has focused in the Crixás greenstone belt, which hosts the 7th largest Au-only deposit in Brazil, the Serra Grande Mine (ca. 80t Au, average grade 5g/t, Anglo Gold; Jost et al., 2014). Most gold deposits in the GAB are classified as orogenic and they include the Serra Grande Mine (e.g., Mina Inglesa, Mina III, Pompex deposits) in the Crixás greenstone belt, the Pilar Mine (e.g., Cacheira do Ogó, Jordino deposits) in the Pilar de Goiás greenstone belt, the Maria Lázara deposit in the Guarinos greenstone belt, and the Sertão and Cascavel deposits in the Faina greenstone belt. Presently, there is no evidence for gold mineralization within the Archean granite-gneiss terrains (Jost et al., 2014).

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4.2. Deposit Structural Controls The distribution of the gold deposits and the geometry of the ore shoots display strong structural control, with mineralization typically hosted in intensely deformed zones associated with shear zones (Fortes, 1996; Jost and Fortes, 2001). At the regional and camp scales, high- grade gold deposits (> 2g/t Au) are spatially associated with thrust and/ or strike-slip faults, ductile-brittle shear zones and folds developed during D2 (-D3) deformation events. These structures are commonly located parallel to lithological contacts such as those between the volcano-sedimentary sequences and the surrounding TTG complexes (Jost and Fortes, 2001). In the northern portion of the GAB, D2 thrust faults are consistently verging to the east, whereas strike-slip faults include dextral faults e.g. Engenho Velho and Faina, in Guarinos and Faina greenstone belts, respectively, or sinistral faults e.g. Cachoeira do Ogó (in the Pilar de Goiás greenstone belt; Jost and Fortes, 2001). At a deposit scale, ore shoots typically comprise elongated (up to 1500m; Fortes, 1996; Jost and Fortes, 2001) bodies with down-plunge continuity, particularly developed in dilatational sites associated with high-strain shear zones. Other high-grade ore shoots are controlled by the intersections of these structures with chemically favourable units, e.g., carbonaceous schist and BIF (Mina Nova and Sertao deposits in the Crixás and Faina greenstone belts, respectively), and are typically stratabound (Jost and Fortes, 2001).

4.3. Hydrothermal Alteration and Mineralization Styles Hydrothermal assemblages associated with Au mineralization in the GAB indicate enrichment in SiO2, CO2, K2O and S (Jost et al., 2001). Major controls on hydrothermal alteration include host rock chemistry, structural setting, the fluid/rock ratio, and permeability during deformation. The gold-related hydrothermal alteration typically overprints metamorphic assemblages, but the distinction between regional metamorphic host rocks and distal alteration zone can be subtle. Three major mineralization styles are observed in the GAB: disseminated sulfide, vein-hosted and massive sulfide (Table 2). Most deposits are associated with a single style of mineralization, although several contain more than one style. Disseminated sulfide: the most common of mineralization style observed in the Crixás, Guarinos, Pilar de Goiás and Faina greenstone belts. It consists of disseminated sulfides including pyrrhotite- arsenopyrite-pyrite±chalcopyrite in 0.1 to 5m-wide proximal alteration zones surrounding quartz±carbonate veins. It consists of 1500 m wide, 200 m long orebodies that extend up to 1000 m down plunge. This mineralization style is developed in hydrothermally altered wall rock associated with ductile shear zones, typically hosted by carbonaceous schists, metavolcanic and metasedimentary Fe- rich rocks. The hydrothermal alteration consists of white mica-quartz-ankerite/siderite±chlorite assemblages, with minor albite±K-feldspar and associated quartz±ankerite/siderite veins. At Sertão deposit, proximal white mica-K-feldspar-pyrite-chalcopyrite alteration zone shows poikiloblastic

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textures that suggest replacement of Fe-rich carbonates by sulfides (Fig. 3). Gold is in equilibrium with pyrrhotite (Crixás) and arsenopyrite (Sertão, Jordino, Ogó), or as inclusions and fracture-fills within these sulfides (≤ 50 µm), with rare free gold in both quartz from veins and proximal alteration assemblage (≤ 30 µm). The approximated gold fineness associated with this style of mineralization is 310 (Table 3). At the Crixás deposits, estimated production of 3 Mt of ore is reported with an average of 6 g/t Au (Jost and Fortes, 2001), whereas at Sertão deposit, oxidized sulfide-rich ore mined 256 Koz at 24.95 g/t (Troy Resources).

Massive sulfide: Consists of 0.5-2.5 m wide, 50-200 m long foliation-parallel sulfide-rich orebodies. This mineralization style is developed adjacent to the contact between metabasalts and carbonaceous schists, e.g., Palmeiras and Zona Superior orebodies, in the Crixás greenstone belt (Fortes, 1996). At the Zona Superior orebody (Serra Grande Mine), massive sulfide lenses extend up to 200 m down-plunge, with an estimated production of 2 Mt with an average of 12 g/t Au for deposits in the Crixás greenstone belt (Jost and Fortes, 2001). Sulfides (up to 95 vol. %) include massive pyrrhotite and/or arsenopyrite, with subordinated magnetite, bornite, chalcopyrite, and ilmenite (Jost et al., 2014). Gold (0.1-2 mm) is distributed as irregular grains with an approximated fineness of 900 (Fortes, 1996; Fortes et al., 2000). In the Faina greenstone belt, minor stratabound massive sulfide mineralization is associated with chemically reactive host rocks as BIF e.g., in the Sertão. In this deposit, pervasive replacement of Fe-rich carbonates (Fe-dolomite, ankerite, siderite) by Fe-sulfides is observed (Chapter 2). Vein-hosted: Consists of deformed quartz±carbonate veins locally hosted in shear zones. It forms 0.5-2m wide, 500 m long, 1500 m down-plunge orebodies that record an estimated production of 3 Mt with an average of 8 g/t Au for deposits in the Crixás greenstone belt (Jost and Fortes, 2001). Despite the lack of available reserve associated with this mineralization style in the Faina greenstone belt, V2 veins in the Cascavel deposit record grades of 4 g/t (Orinoco Gold Ltd. 2014 internal report). Mineralization is hosted by carbonaceous schist and quartzite in the Crixás Faina greenstone belts, respectively. At the Crixás deposits (Mina III), typical white mica-carbonate-chlorite-albite-pyrrhotite- arsenopyrite assemblages enveloping D2-related veins contain gold (0.1-2 mm) as free grains or within fractures associated with fineness approximately 910 (Jost and Fortes, 2001). In Cascavel, distal chlorite-quartz±K-feldspar, intermediate quartz-fuchsitic mica-biotite-pyrite and proximal white mica- quartz±K-feldspar±pyrite±chalcopyrite alteration zones developed around V2 veins display free gold (≤ 250 mm) with a fineness of 992 (Chapter 2).

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Table 2. Mineralization styles of orogenic gold deposits in the GAB

Fig. 4. Hand samples and photomicrographs showing styles of mineralization in orogenic gold deposits in the Goiás Archean Block (GAB). A. Vein-hosted mineralization next to carbonaceous schist of Inferior Zone orebody, Mina III deposit (Crixás; from Coelho, 1999); B. Massive sulfide mineralization in the Superior Zone orebody, Mina III deposit (Crixás) consisting mostly of arsenopyrite, pyrrhotite, pyrite, with minor ankerite/siderite (from Fortes, 1991); C. Ore sample showing gold mineralization associated with disseminated sulfides in the Maria Lázara deposit (Guarinos). Note textural relation of typical idiomorphic arsenopyrite overprinting white-mica-tourmaline hydrothermal alteration (from Pulz, 1990); D. Laminated quartz veins and gold disseminated in pyrite-arsenopyrite-pyrrhotite of Ogó

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deposit (Pilar de Goiás; from Souza, 2012), and E. Isoclinal F2 folds outlined by vein-hosted mineralization in the Cascavel deposit (Faina).

Table 3. Summary of characteristics of orogenic gold deposits in the GAB

4.4. Fluid conditions Few available studies have focused on the conditions of ore-related fluids (Pulz, 1990, 1995; Fortes, 1996; Petersen Jr., 2003). Fluid inclusions in quartz from the Mina III disseminated massive sulfide in the Crixás greenstone belt (Fortes, 1996) show highly saline fluids in the H2O-CO2-NaCl-

KCl-CH4-N2 system. However, vein-hosted quartz also from the Mina III has low salinity (< 10 wt. %

NaCl eq.), H2O-CO2-rich fluids, and minor saline (> 8 wt. % NaCl eq.), H2O-CO2-rich fluids enriched in CH4, N2 with rare methane- and nitrogen-rich inclusions (Fortes, 1996). Estimated fluid temperatures at the Mina III deposits range from 350° to 475°C, at pressures of 2-3 kbar (Fortes, 1996). Garnet- biotite geothermobarometry and fluid inclusions in quartz from disseminated and vein-hosted mineralization styles in the Crixás greenstone belt indicate a temperature range of 428° to 580°C, and pressures of 5.7 to 8.3 kbar (Petersen Jr., 2003). These temperatures are comparable with greenschist to lower amphibolite facies, although the higher pressure conditions may reflect increased crustal depth at the time of mineralization (Petersen Jr., 2003). Estimates of fluid conditions at deposits hosted in other northern greenstone belts include studies by Pulz (1990, 1995) at the Maria Lázara and Cachoeira do

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Ogó deposits, in the Guarinos and Pilar de Goiás greenstone belts, respectively. According to these studies, 335° to 450°C temperature ranges and 2 to 3 kbar pressures conditions are associated with hydrothermal fluids in these deposits. In the Faina greenstone belt, chlorite and arsenopyrite geothermometry suggest temperatures of 330°-400°C and 320°-430°C for the Cascavel and Sertão deposits, respectively. Thus, the deposits of the GAB have estimated fluid temperatures comparable with data from similar greenstone-hosted orogenic gold deposits (Table 3).

4.4.1. Gold transport and precipitation mechanisms The transport and precipitation of gold in orogenic Au deposits of the GAB is interpreted to have occurred by several different mechanisms, as evidenced by contrasting occurrence and composition of gold. Major controls include destabilization of sulfide and chloride complexes, phase immiscibility and variations in pH and fO2. Mineralization during episodic flux of hydrothermal fluids (e.g. Sibson, 1990) is widely attributed to gold deposits in the GAB. Evidence for the multistage precipitation of ore-related minerals is illustrated by (i) distinct paragenetic generation of ore-related mineral phases, (ii) contrasting Ag contents of gold (e.g., Maria Lázara; Michel et al., 1994), (iii) free and refractory gold (e.g., Cascavel and Sertão deposits), (iv) crack-and-seal textures in vein-hosted mineralization (e.g. Pulz, 1990), and (v) dissolution of ore-related minerals such as arsenopyrite (e.g. Pulz, 1995). Textural relationships indicate most gold deposition occurred synchronously with sulfide precipitation (Fortes and Giuliani, 1993; Pulz, 1990, 1995; Petersen Jr., 2003; Chapter 2). The spatial and temporal relationships with sulfide suggest bisulfide complexes likely constitutes an important mechanism for the transport of gold. Sulfidation reactions imply a decrease in sulfur fugacity, which induces instability in the gold complexes and promotes gold precipitation along with sulfide minerals (Seward, 1973; Loucks and Mavrogenes, 1999).

Previous studies argue gold deposition in Fe-rich rocks is enhanced via pH and fO2 variations associated with wall rock sulfidation (Roberts, 1987; Phillips and Groves, 1993). H2S loss from the fluid involving Fe-sulfides, as illustrated by FeOrock + H2S(aq) -> FeS + H2O, implies in the destabilization of trio-complexes and further Au deposition (Mikucki, 1998). Similarities of the deposits presenting disseminated and massive sulfide mineralization styles with deposits hosted in Proterozoic Fe-rich sedimentary rocks, e.g. Tanami in Northern Territory Australia (Lambeck et al., 2011) and Lamego, in the Quadrilátero Ferrífero, Minas Gerais, Brazil (Lobato and Vieira, 1998; Martins et al., 2016), suggest analogous depositional processes. Although ore depositional mechanisms were likely similar, local physico-chemical conditions, e.g., pressure control on fluid immiscibility, may play an important role at the deposit scale (Mickuki et al., 1993). Disparate mechanisms of metal transport/precipitation and pressure/temperature conditions are inferred to dictate late Au-Te-Bi mineralization at the Maria Lázara deposit (Pulz, 1990).

Precipitation as native gold or maldonite (Au2Bi) presumably replaced gold-telluride (e.g. Paddefet, 1982) upon cooling and exsolution from disseminate sulphide in proximal white-mica and intermediate

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K-rich alteration zones. Stable between 113-371°C (Barton and Skinner, 1976), at lower temperatures maldonite tends to be substituted by Au-Bi intergrowths (Anthony et al., 1990). At the Cascavel deposit, vein-hosted mineralization contains mainly ‘free’ gold and minor sulphides, therefore, suggesting that phase immiscibility and subsequent precipitation of ‘free’ gold from the hydrothermal solution was the main gold forming process (cf. Mikucki et al., 1993).

Fig. 5. Photomicrographs showing ore assemblage, mode of occurrence and distribution of gold in the GAB deposits. A. Gold and chalcopyrite filling microfracture of arsenopyrite from massive sulfide in Mina III deposit, Crixás (from Fortes, 1991). B. Gold (≤ 200 µm) with pyrrhotite-chalcopyrite- arsenopyrite of massive sulfide from Mina III deposit, Crixás (from Fortes, 1991). C. Native gold bordering arsenopyrite grain in disseminated mineralization in the Maria Lázara deposit, Guarinos. Dissolution microcavities in gold grains suggest remobilization (from Pulz, 1990). D. Idiomorphic arsenopyrite overprinting sulfides (Py1 and Py2) in proximal BIF in the Sertão deposit, Faina (Chapter 2). Rarely, the latter is observed superimposed by white mica±sulfide. E. Textural contrast between deformed, fine-grained Py1 and coarse-grained Py2 of proximal BIF in the Sertão deposit, Faina. Note the poikiloblastic texture developed after replacement of ankerite/siderite by Py2 (Chapter 2). F. Siderite overprinting Py1 and being replaced by Py2 of BIF in the Sertão deposit, Faina. (Chapter 2). G. Gold (≤ 50 µm) in equilibrium with arsenopyrite showing microfractures and dissolution features of proximal carbonaceous schist in the Sertão deposit, Faina. H. BSE image showing elongated gold grain (ca. 30 µm) in equilibrium with arsenopyrite showing pressure shadows in between white mica of proximal carbonaceous schist in the Sertão deposit, Faina (Chapter 2). I. Free gold (ca. 20 µm) adjacent to subhedral arsenopyrite grain in equilibrium with pyrite and chalcopyrite of BIF in the Sertão deposit, Faina (Chapter 2). Abbreviations correspond to Asp: arsenopyrite, Cpy: chalcopyrite, Py: pyrite, Po: pyrrhotite, Sid: siderite, Au: gold, Wm: white mica.

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In several deposits, including Cachoeira do Ogó and Cascavel in the Pilar de Goiás and Faina greenstone belt, gold occurs as free grains in quartz veins, indicating the involvement of contrasting ore-forming processes. The deposition of free gold in D2-related veins at the Cachoeira do Ogó and Cascavel deposits is possibly caused by fluid immiscibility (or boiling) and subsequent lowering of the gold solubility (Brown, 1986; Seward, 1989). The cause of fluid immiscibility in Au-bearing veins associated with D2 thrust faults is likely due to cyclic decompression of the hydrothermal fluid caused by seismic movement along the thrust faults and veins (Sibson et al., 1988; Robert et al. 1995, Dugdale and Hagemann, 2001). In this scenario, gold transport involving vein-hosted mineralization style would be facilitated via chloride complexes (e.g. Chapter 2).

4.5. Timing of Au mineralization

A longstanding scientific dispute concerns the timing of the gold mineralization event in the GAB (e.g. Thomson and Fyfe, 1990; Fortes, 1996; Marques et al., 2013). Based on early interpretations of greenstone belt sequences, gold mineralization was initially interpreted to be Archean in age (Danni and Ribeiro, 1978). Subsequent studies proposed that massive sulphide orebodies in the CGB formed post-peak metamorphism associated with the Brasiliano orogeny (Thomson and Fyfe, 1990). Neoproterozoic ages for gold mineralization were reinforced by K-Ar and Ar-Ar ages from amphibole at 660-730 Ma, biotite and chloritoid at 520-580 Ma, and biotite and muscovite at ~ 500 Ma (Fortes et al., 1997), together with a whole-rock Rb-Sr isochron of Au-bearing chlorite-garnet schist at 505 ± 7 Ma (Fortes et al., 2003). Irrespective of the geochronological evidence pointing to a Neoproterozoic gold event, these authors do not disregard the possibility of older ages for the gold mineralization event.

The use of geochronology methods easily disturbed by subsequent thermal events (e.g. K-Ar, Ar- Ar dating) is unreliable to constrain ages within polymetamorphic terranes such as the GAB and, therefore, are not taken in consideration. Additional Neoproterozoic ages including intrusion-related 726 Ma mineralization (Rodrigues, 2011) is at odds with the main orogenic gold mineralization in the GAB and thus is not discussed herein.

Most available geochronology points to a Paleoproterozoic age for the gold mineralization event. For instance, 2025 Ma galena Pb-Pb model age is presented for orebodies in the Pilar de Goiás greenstone belt (Pulz, 1995). Considering the distinct analytical techniques, the initially proposed model age are comparable with zircon U-Pb SHRIMP age of 2145 ± 12 Ma for albite granite intrusion in the same belt (Queiroz, 2000). This is consistent with U-Pb SHRIMP age of 2165 ± 47 Ma obtained for hydrothermal zircon hosted in mineralized metagreywacke reported by Tassinari et al. (2006). These are supported by arsenopyrite Re-Os from massive sulphide of Mina III dated at 2126 ± 16 Ma (Marques et al., 2013).

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The maximum age of gold mineralization in the Faina greenstone belt is indirectly constrained by the 2061 Ma Pink Syenite intrusion, which is interpreted as synchronous to late in relation to Au- related deformation event (Chapter 3). The latest Paleoproterozoic deformation event resulted in the development of the Arraial Dantas paleoplacer (Carvalho et al., 2013) in basal meta-conglomerate of the Faina greenstone belt. Evidence to support that the paleoplacer post-dates the main Au-related hydrothermal event include the presence of deformed clasts of Au-bearing quartz-veins, disseminated mineralization in carbonaceous schist and gold in a fine-grained matrix (Carvalho et al., 2013). This is correlated with the formation of paleoplacer gold deposits at Orosian times in the Jacobina greenstone belt, in the São Francisco craton (Teixeira et al., 2010).

4.6. Exploration fingerprints

Overall, Fe-rich rocks, e.g., BIF, comprise important hosts to Au mineralization. Similar gold deposits hosted in Paleoproterozoic sedimentary sequences have applied the use of element ratios to track mineralization (Lambeck et al., 2011). According to these authors, Paleoproterozoic Fe-rich rocks deposited along active continental margin settings associated with low Th/Sc and high Cr/Th ratios are interpreted to be associated with gold mineralization. The geochemical signature of gold mineralization includes Au-As at deposits hosted in the Crixás greenstone belt (Fortes, 1996), Au-Ag-Sb-Te-Bi in Guarinos (Pulz, 1990), Au-Ag-Bi-Mo-Pb-Sb-W in Pilar de Goiás (Carvalho, 1999). Syngenetic deposits represented by Neoproterozoic leucogranite intrusive in the Guarinos greenstone belt display major Au- S-Se-Sb-Ag-Te-As and minor Ba-Zn-Pb-Mo geochemical signatures (Jost and Fortes, 2001). The hyperspectral band range profile of typical hydrothermal alteration minerals, e.g., white mica, chlorite and carbonate, can be applied as an exploration tool using hand-held or airborne hyperspectral analyses as demonstrated, for example, in the Kanowna Belle and Sunrise Dam orogenic gold deposits in Western Australia (Wang et al., 2017). In the case of the Sertão and Jordino deposits, the appearance of idioblastic arsenopyrite can be correlated with elevated gold content, therefore is an excellent indicator for gold-bearing zones. Chemical changes depicted by mineral chemistry support the definition of an alkali-rich proximal alteration zone, e.g., in the Cascavel and Sertão deposits (Chapter 2), therefore the use of an alkali index such as K/Al (Kishida and Kerrich, 1987) in whole- rock geochemistry data may be useful to track potential ore zones. Similarly, the widespread carbonate alteration around ore zones can be more easily discriminated via the carbonation index (CO2/ (Fe+Mg+Ca) in whole-rock geochemistry data.

5. Geological evolution of the GAB

The shortage of more detailed geochemical and geochronological dataset available for the granite-greenstone terranes of the GAB precludes a better understanding of the petrogenetic

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mechanisms and geodynamics involved throughout its formation. Albeit the still unclear geological framework, the integration of previous data with findings presented in this research is used in an attempt to provide the tectono-magmatic context throughout the evolution of the GAB. This is summarized in Fig. 6 and 7 for the northern and southern GAB, respectively. Isotopic evidence for crust generation is illustrated in Fig. 8 and detailed by topics as follows.

In the northern GAB, TTG magmatism occurs in two main stages spanning 70 m.y. (Queiroz et al., 2008). An early-stage from 2845 to 2785 Ma consists of tonalite to granodiorite (with minor granite) orthogneisses with juvenile εNd signature (Anta, Caiamar and Hidrolina terranes). The recognition of up to 3300 Ma inherited ages for these rocks attest the involvement with older crust, which is devoid of any exposure until the moment. Another stage from 2711 to 2707 Ma comprises granodiorite to granite orthogneisses with crustal εNd signature (Moquém terrane). The isotopic zonation for this magmatism denotes an eastward migration of crustal growth (Queiroz et al., 2008). The diachronous evolution of TTG granite-gneiss terranes in the northern GAB is augmented by whole-rock geochemistry for these rocks (Vargas, 1992). According to the former, pre- to syn-collisional intrusions from tonalite and granodiorite early stage show sub-alkaline to calc-alkaline affinities, whereas syn-collisional to post- tectonic granodiorite and granite intrusions from the second stage are calc-alkaline to metaluminous.

The Caiamar terrane comprises roughly coeval Águas Claras and Tocambira intrusions (~ 2840 Ma), which are intruded by the 2820 Ma Crixás-Açu gneiss (Queiroz et al., 2008). The Hidrolina terrane is formed by 2785 Ma granodiorite that is later intruded by 2146 Ma Posselândia diorite (Jost et al., 1993; Queiroz et al., 2008). The Anta terrane includes 2840-2820 Ma granodiorite orthogneisses followed by 2790 Ma granitic stocks of the Chapada suite (Lacerda Filho et al., 2000; Queiroz et al., 2008; Beguelli Jr., 2012). Fractionated REE, particularly HREE, and slightly positive εNd (0.74) support contribution of old crust for the latter (Begueli Jr., 2012). Despite radiometric signature that assigns its limit with the Caiçara terrane (Blum et al., 2003), isotopic, geochemical and petrological similarities of the Anta with the latter led Beguelli Jr. (2012) envisage their probable cogenetic relation. However, 2818 Ma charnockite intrusions in the Caiçara terrane (Begueli Jr., 2012) reinforce the initial distinction of these terranes (Fig. 8). In the southern GAB, two major pulses of magmatism spanning 40 m.y. in the Uvá and Caiçara terranes are recorded by mostly juvenile 3040-2930 Ma and slightly crustal-derived 2890-2820 Ma orthogneisses (Pimentel et al., 1996b; Beguelli Jr., 2012; Jost et al., 2005, 2013; Borges et al., 2017; Chapter 3). Inheritance registered at 3090-3050 Ma predicts the involvement of up to 3100 Ma crust. Older model ages up to 3500 Ma from Pimentel et al. (1996b) for high 147Sm/144Nd ratio samples might correspond to the fractionation of the Sm-Nd system during metamorphism (Pimentel et al., 2003a). Amphibolite facies metamorphism in the Archean is depicted by 2772 Ma U-Pb zircon age and 2711 Ma titanite U-Pb age for Crixás-Açu gneiss, in the Caiamar terrane (Queiroz, 2000), whereas in

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the southern GAB, metamorphism is dated at ca. 2840 Ma (Jost et al., 2013). This supports the stabilization of the GAB occurred at ca. 2700Ma.

Fig. 6. Schematic block diagram evolution of northern GAB during Archean (A), Paleoproterozoic (B) and Neoproterozoic (C) times (not to scale).Abbreviations in A for the Caiamar TTG terrane refer to Crixás-Açu (CA), Tocambira (T) and Águas Claras (AC).

In the northern GAB, volcanic sequences of the CGB reveal spinifex textures (Teixeira, 1981; Teixeira et al., 1981; Kuyumjian and Teixeira, 1982) typically associated with komatiite flows (Arndt, 1994). Ultramafic flows are overall characterized by smoothly-sloping, LREE-rich patterns, whereas basalts show flat REE patterns and slight depletion in both LREE and HREE (Arndt et al., 1989). The metabasalts of the Guarinos greenstone belt have a tholeiitic affinity, slightly fractionated REE patterns and negligible negative Eu anomaly proposed to be formed in a backarc environment (Jost et al., 1999). Volcanic rocks in the Pilar de Goiás greenstone belt show flat to slightly fractionated REE patterns interpreted as similar to tholeiitic lavas (Santos, 2008). In the southern GAB, preserved pillowed structures in lower volcanic sequences attest to the subaqueous nature of these sequences that allowed their correlation with komatiites (Borges et al., 2017). According to the same authors, the geochemistry of lower ultramafic rocks is characterized by tholeiitic, sub-alkaline to calc-alkaline affinities and flat REE patterns. Mafic sequences of both greenstone belts are LREE-rich, HREE-poor and show subtle negative to positive Eu anomalies (Borges et al., 2017).

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Two main events are recorded in the GAB during Paleoproterozoic times (Queiroz et al., 2008). An early crustal extension recorded by ~ 2300 Ma epicratonic mafic dike swarm in TTG terranes (Tomazzoli, 1997; Corrêa da Costa, 2003) and compression expressed by ~ 2150 Ma metamorphic age obtained from titanite (Queiroz, 2000). Contrary to the Archean age of Crixás, Faina and Goiás lower sequences, volcanism at Guarinos and Pilar de Goiás greenstone belts dates approximately 2200 Ma (Jost et al., 2008, 2012, 2014). Yet, sedimentary rocks in all greenstone belt sequences were deposited in the Paleoproterozoic (Jost et al., 2014). Uneven magmatism at this is represented by the 2146 Ma Posselândia diorite stock hosted in NW-trending shear that crosscuts the Hidrolina TTG (Jost et al., 1993), 2145 Ma albite granite along N-verging thrust faults intrusive in metasedimentary rocks from the northern greenstone belts (Queiroz, 2000) and 2170 Ma mafic dikes that crosscut Au-mineralized orebodies at the CGB (Jost et al., 2010).

Fig. 7. Schematic diagrams showing the evolution of southern GAB during Archean (A), Paleoproterozoic (B) and Neoproterozoic (C) times (not to scale; Chapter 3).

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Influence of the Neoproterozoic has been pervasively credited to thermal disturbance along granite-gneisses in the GAB. Additional features at this time include 625 Ma age interpreted as resulting from the anataxis of the Caiçara terrane (Pimentel et al., 2003a) and 729 Ma leucogranite intrusive in the Guarinos greenstone belt (Rodrigues, 2011). However, the interpretation of these data is contested. First, thermal disturbance commonly displayed in polydeformed terranes does not necessarily denote an intrusive event (e.g. Pimentel et al., 2003a). Secondly, the shortage of analyses, method applied and lack of relevant synchronous magmatism in the GAB brings into question the reliability of the data reported for leucogranite in the Guarinos greenstone belt. Even so, more research needs to be conducted to address this interpretation. In the southern GAB, the significance of marginal Rio Caiapó, Itapuranga and Serra Negra intrusions is explored in Chapter 3.

5.1. Towards a model The GAB records U-Pb crystallization ages and inheritance peaks at 3000, 2800 and 2700 Ma that overlap with fairly continuous magmatism (Fig. 9A-D). Magmatic quiescence occurs between 2930-2890 Ma (in the south) and 2840-2710 Ma (in the north). Temporal and spatial relationship between TTG plutonism and greenstone belt development require a tectonic process that allows the coeval formation of both komatiite-tholeiite basalt sequences and calc-alkaline magmatism. Longstanding controversy ascertains the settings and geodynamic processes of Archean crustal growth (e.g. Kemp and Hawkesworth, 2003; Smithies et al., 2009; Condie and Aster, 2010; Kamber, 2015). Archean cratons are dominated by ‘arc-like’ tonalite-trondhjemite-granite i.e. TTG association (Drummond and Defant, 1990; Moyen, 2011) and ‘plume-like’ komatiite-tholeiite basalt association (Hollings and Wyman, 1999; Van Kranendonk et al., 2015). As a result, Archean crustal growth models involve a dichotomy between subduction and plume-related processes, and disagreements over the timing of onset of plate tectonics (Rey et al., 2003; Stern, 2005; Bédard, 2006; Shirey and Richardson, 2011; Van Kranendonk, 2011; Wyman, 2013). Typical upward younging, tholeiitic basalt-komatiite association and geochemistry similar to Archean greenstone belts make oceanic plateaux an appealing model (Arndt, 1994; Condie, 2005; Barnes and Van Kranendonk, 2014). In this context, the production of contemporaneous bimodal geochemistry and TTG-like magmas can occur by infracrustal melting at the base of a thick plateau (Willbold et al., 2009). As a result, oceanic plateau (Hollings and Kerrich, 1999; Fralick et al., 2008), sagduction (François et al., 2014), plume-derived continental drift (Tomlinson et al., 1998) and plume- arc (Hollings and Wyman, 1999) are often offered to explain the generation of granite-greenstone terranes (Percival et al., 2012). An argument against this theory, however, is that plume-derived basalts, commonly associated with negligible water contents, tend to form a scanty source for partial melt (Arndt, 2013). The melting of an anhydrous source contrasts with the ubiquitous hydrous mineralogy associated with that of Archean TTGs (Moyen and Stevens, 2006). Apart from that, the refractory and cumulus nature typical

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of lower mafic-ultramafic sequences is likely a very unfertile source for evolved magmas (Arndt, 2013). Yet, investigations have shown that plume-derived basalts provide an appropriate geochemical (Martin et al., 2014) and isotopic (Guitreau et al., 2012) source for Archean TTGs.

Fig. 8. Map showing available εNd and εHf in the Goiás Archean Block. The diachronous stabilization of TTGs is evidenced by crustal growth migration towards the east in the northern TTGs, suggested by increasingly crustal input eastward and dominant juvenile TTG crust in the south. References: Pimentel et al. (2003a, 2003b), Tassinari et al. (2006), Queiroz et al. (2008), Jost et al. (2005, 2008, and 2013), Beguelli Jr. (2012), and this work.

Recent interpretations have proposed a sagduction plume-related scenario for the early evolution of terranes in the GAB (Jost et al., 2019). Several studies advocate for similar models to explain Archean geodynamics (Chardon et al., 1996; Bédard et al., 2003; Van Kranendonk et al., 2004; Bédard, 2006; Robin and Bailey, 2009). According to this paradigm, granitic rocks are formed via partial melting of a thick basaltic pile in an oceanic plateau (Bédard, 2006), which implies mantle plume magmatism. However, this hypothesis is challenged by several arguments. Based on Jost et al. (2019), after the formation of ≤ 3300 Ma sialic crust (so far devoid of exposure), rift induced by a hotspot or mantle plume led to the production of ca. 3000 Ma komatiite and pillowed basalts in an oceanic island, stratovolcano or plateau environment. According to the same authors, the diapiric emplacement of felsic magmatism is attributed to originate the Anta and Caiamar terranes at least 200 m.y. after crystallization of their igneous protoliths.

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Nevertheless, crust formation in oceanic plateaux, assumed to occur during extension (after a compressional event), would result in upward younging and low metamorphic grade that preclude tectonic stacking via horizontal tectonics (Van Kranendonk et al., 2007). This disagrees with the scenario previously proposed for the CGB, which advocates for high-temperature, amphibolite facies metamorphism in the Archean and stacking of the stratigraphy during early deformation events (e.g. Queiroz, 2000; Jost and Fortes, 2001; Jost et al., 2014). The Nd isotopic record for the Anta (0.01 to 0.74; Queiroz et al., 2008; Beguelli Jr., 2012) and Caiamar terranes (-0.63 to 2.41; Queiroz et al., 2008; Fig 9A) implies dominantly juvenile sources. Plume-related generation of TTG in an oceanic plateau could be either by: (i) density inversion (Bédard, 2006), (ii) melting of previous crust (Smithies et al., 2009), or (iii) differentiation of underplated coeval basaltic melts. The two first options imply in the reworking of older sources, which conflicts with the isotopic record of these terranes. In turn, the last option would reflect higher εNd values that are also inconsistent with the juvenile signature presented by previous studies (Queiroz et al., 2008; Beguelli Jr., 2012). Another point of content debated is whether the triangular, narrow, cusp-shaped and kilometre- deep keel shape of the CGB (Blum et al., 1996) can be used to support plume-related model invoked for the early generation of continental crust in the GAB (Jost et al., 2019). Besides the very limited evidence provided for vertical tectonics, the parameters used to demonstrate this hypothesis are fairly vague. For example, there is no conspicuous record of steep regional structures in the CGB. This can be partially explained by the intense deformation and low preservation of any earlier formed structures in the area. Yet, low-pressure, contact-style metamorphism characterized by isograds concentrically distributed around TTG domes varying from high-temperature (i.e. amphibolite facies) adjacent to the intrusions to low-temperature (i.e. prehnite-pumpellyite facies) away from those (e.g. Hickman, 1984), which is also not consistent with previous structural-metamorphic investigations conducted in the CGB (Thomson and Fyfe, 1990; Jost and Oliveira, 1991; Thomson, 1991; Jost et al., 2001, 2014). An additional piece of evidence against a plume-related hypothesis is that, despite being proposed to explain the formation of other granite-greenstone associations e.g. in the Kaapval Craton in South Africa and the Pilbara Craton in South Australia (Condie, 1994), the extremely thick mafic-ultramafic magmatism attributed to these areas (in the order of several kms) is much more pronounced than the volcanism associated with terranes in the GAB (up to 900m in the Pilar de Goiás greenstone belt; Jost and Oliveira, 1991; Jost et al., 2011, 2014). However, the genesis of lower supracrustal sequences in the GAB is still poorly explored and a plume-related origin is not completely disregarded. Hybrid examples of the two end-members of models proposed for crust formation in the Archean are reported in other locations (e.g. Bleeker, 1999; Polat et al., 1998; Percival, 2007). Therefore, the precise genesis of Archean crust in the GAB remains unclear and more investigation with regards to the involved petrogenetic and geodynamic processes need to be conducted.

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Diversification of magmatism during early Earth is manifested by the transition from Na-rich TTG suites to K-rich melts and mantle-derived sanukitoid rocks after 2.7 Ga (Corfu and Stone, 1998; Bédard, 2006; Laurent et al., 2014). Similar progression recorded at different times in other cratons points to a global process in the late Archean (Smithies and Champion, 2000; Laurent et al., 2011; Heilimo et al., 2013). These changes are generally succeeded by major episodes of crustal growth (Hawkesworth et al., 2018). The detection of sanukitoid magmatism in the GAB has implications for Archean geodynamics and crustal growth. High compatible and fluid-mobile elements typical of these rocks support subduction-related mantle wedge contribution that provides additional evidence for subduction (Shirey and Hanson, 1984; Smithies and Champion, 2000; Martin et al., 2010; Laurent et al., 2014). Moreover, post-tectonic sanukitoid intrusions underpin a major shift in igneous magmatism (Halla et al., 2009;

Laurent et al., 2014; Martin et al., 2014). The lower SiO2 coupled with higher MgO, FeO, Mg# and increased compatible transition metals e.g. Ni, Co of these rocks are assumed to represent mantle contribution (Shirey and Hanson, 1984; Stern, 1989). Their steep REE, in particular HREE, are attributed to high pressure melting in the presence of garnet (Rapp et al., 2010). Thus, sanukitoid rocks tend to occur along, or adjacent to, terrane boundaries, such as intrusive suite between the Caiçara and Anta terranes presented by Beguelli Jr. (2012). Geochemically, volcanic rocks in greenstone belt are characterized by: (i) plume-related tholeiitic basalt-komatiite in oceanic and continental plateau (Arndt, 1994; Polat et al., 2009), and (ii) subduction- related calc-alkaline basalts, andesites, dacites and rhyolites with other subordinated rock types e.g. boninites, adakites, Nb-rich basalts (Kerrich et al., 1998; Wyman et al., 2000; Percival et al., 2003; Polat and Kerrich, 2006). The derivation of tholeiitic basalt-komatiite associations in greenstone belts has been attributed to high-temperature plume activity (Condie, 1994; Polat et al., 2009). Alternatively, the subduction-related origin of this association in fore-arc settings have been proposed based on Phanerozoic boninites (e.g. Barberton greenstone belt; Parman et al., 2001). These rocks, formed by hydrous melting of the refractory mantle at shallow depths (Crawford et al., 1989), are intrinsically related to early subduction in intraoceanic fore-arc settings (Pearce et al., 1992).

The comparable chemistry of ultramafic rocks to modern boninites (SiO2 > 53%, Mg# >60, TiO2 <0.5 wt. %; Crawford et al., 1989) support a subduction-related origin of komatiites for the southern GAB greenstone belts (Borges et al., 2017). The flat REE patterns and Nb-rich contents of mafic sequences, similar to back-arc and Nb-rich basalts augment the involvement of subduction processes for the genesis of volcanic rocks in these sequences (Borges et al., 2017). Therefore, volcanic sequences characterize fore- to back-arc rifts (Borges et al., 2017) during accretionary extension, whereas calc- alkaline intrusions are proposed to form at transient compression by shallow subduction in the southern GAB (Chapter 3).

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Fig. 9. A. Previous εNd(T) data for supracrustal rocks and felsic magmatism in the GAB (Arndt et al., 1989; Pimentel and Fuck, 1994, 1996b; Queiroz et al., 2000; Fortes et al., 2003; Pimentel et al., 2003b; Tassinari et al., 2006; Beguelli Jr., 2012; Jost et al., 2008, 2013, Borges et al., 2017). B. εHf(T) isotope data for the southern GAB (Chapter 3). C. δ18O isotope composition for granitic intrusions in the southern GAB (Chapter 3). D. Histogram with crystallization ages for igneous rocks in the GAB (all references cited above).

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Early whole-rock Sm-Nd isochron age of 2825 ± 98 Ma and mostly juvenile εNd (0.6) data reported for komatiite flows in the CGB (Arndt et al., 1989) attest to their fairly juvenile source. Other studies propose a depleted mantle source originated from an oceanic crust endorsed by whole-rock Sm- Nd isochron at 2998 ± 70 Ma and εNd(T) of 2.4 for lower metavolcanic rocks of the CGB (Fortes et al., 2003). The only record of reworking could be related to the Nd isotopic signature presented by the latter authors for metasedimentary rocks at the CGB. This data agrees with the above arguments for subduction rather than a plume-related origin of these rocks. In the southern GAB, the 2061 Ma Pink Syenite intrusion is compatible with the onset and prevalence of reworking. The dashed line in Fig. 9B shows the evolution of 2870 Ma crust (Uvá TTG) towards strongly negative εHf(t) values represented by the Serra Negra granite. The trend from a 2870 Ma primitive mantle has a slope that corresponds to a 176Lu/177Hf value of 0.02, which agrees with that of a mafic crust (Chapter 3). Compression is expressed in Phanerozoic orogens by decreasing magmatism, isotopic flare-ups and distinct bulk rock geochemistry (e.g. Kemp et al., 2009; Collins, 2011). The transition to unradiogenic Nd/Hf isotopes in accretionary margins at this time is related to increased contamination in the upper plate. In the northern GAB, the increasingly unradiogenic Nd signature from west to the eastern TTG terranes favours compression. This is compatible with geochemical results on TTG terranes from Vargas (1992). Moreover, diagnostic geochemical and geological parameters provided by Kerr et al. (2000) for volcanic sequences from distinct tectonic settings support the arc-like signature of lower mafic-ultramafic rocks in the northern GAB greenstone belts (e.g., Jost et al., 1999). Nevertheless, the lack of available Nd isotopic signature for the supracrustal rocks, in particular Guarinos and Pilar de Goiás (Jost et al., 2008) hampers more detailed correlations. Thus, calc-alkaline intrusions are proposed to form by episodic compression during shallow subduction and eastward migration of the arc in the northern GAB (e.g. Queiroz et al., 2008).

5.2. Linking crustal architecture to gold mineralization Crust generation and destruction by subduction-collision in convergent margins sculpt continental crust through Earth’s history (Wilson, 1965). The evolution of continental crust develops a diverse lithospheric architecture that effectively controls the development of preferential pathways for magmas and fluids (Begg et al., 2009). The crustal architecture plays a major role in the location of lithospheric-crustal scale structures controlling large Au-related hydrothermal systems (Blewett et al., 2010). The potential for mineralization worldwide in cratonic blocks is illustrated by the occurrence of gold (Groves and Batt, 1984; Robert et al., 2005; Ispolavot et al., 2008; Bateman et al., 2008), base metals (Ashley et al., 1988; Hannington et al., 1999; Cantwell et al., 1999), iron (Khan and Naqvi, 1996; Angerer and Hagemann, 2010; Duuring et al., 2012; Angerer et al., 2012), and nickel (Fiorentini et al., 2012). As recognized by previous studies, the addition of juvenile material into the crust prior to a gold- related event enhances the gold fertility in the terrane (Cassidy et al., 2005; Bateman and Bierlein, 2007;

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Hronsky et al., 2012). Studies in the Yilgarn craton suggests the input of juvenile crust (εNd > 0) before and after gold mineralization acted as a potential source of this metal (Mole et al., 2015). A similar U- Pb-O-Hf isotope approach was applied in this study to provide insights into potential links between crust-mantle evolution and the development of gold systems in the GAB. The geochemistry of TTGs from the GAB indicates their formation in volcanic arcs associated with subduction zones (Vargas, 1992; Queiroz et al., 2008; Borges et al., 2017). The hypothesis that an older Mesoarchean crust might have existed, Based on U-Pb and Sm-Nd TDM model ages (Queiroz et al., 2008; Jost et al., 2013; Borges et al., 2017), is not considered feasible for two reasons. First, the reworking of these rocks indicated by geochronology and isotopic geochemistry is more likely while the magmas are still hot rather than ~ 300 Ma years after their emplacement. In addition, a lack of exposed Mesoarchean crust or inherited U-Pb ages older than 3100 Ma, argues against the presence of continental crust before this time. A more plausible explanation for the extraction of these primitive melts involves crustal reworking occurring shortly after emplacement of the original Archean crust. In the northern GAB, two major magmatic events are represented by ~ 2770 Ma with juvenile- like εNd and ~ 2700 Ma with εNd values, indicative of crustal contribution for the formation of these magmas. The new zircon in-situ U-Pb-O-Hf work indicates crustal formation in southern GAB likely involved the negligible reworking of an older crust of approximately 2870 Ma (Chapter 3), which overlaps with older magmatic event proposed by Queiroz et al. (2008) for the northern GAB. This is consistent with previous observations from Jost et al. (2013). The geodynamics of the southern GAB can be described by the following sequence as: (i) 2960-2920 Ma forearc-arc-back-arc assembly with positive εNd (+2.16 to +2.77) indicative of juvenile arc, followed by (ii) ca. 2800 Ma continental arc with slightly negative εNd (-0.15) that indicate minor crustal contribution during genesis of tonalite magmatism (Borges et al., 2017). Gold deposits in the GAB are spatially located in the greenstone belts (Jost et al., 2001), especially along with major structures adjacent to their boundaries with the TTGs and at intersections of multiple regional structures (i.e., Cox, 1999; Blewett et al., 2010). The metasedimentary and metavolcanic host rocks of gold deposits are mostly characterized by unradiogenic εNd (Fortes et al., 2003) indicative of an old reworked source. In the southern GAB, the emplacement of syn- to late- orogenic Pink Syenite with unradiogenic εHf signature coupled with mantle-like δ18O values indicates a link between mantle, crust, gold-related deformation (Chapter 3). Similar geochemistry, geochronology and isotopic signatures with younger intrusions, e.g. the Serra Negra granite, and the correlation of the latter with newly discovered Cu-Au deposits in the region (e.g. Fazenda Nova mine; Marques, 2017) have implications for regional exploration and it supports the importance of juvenile contribution for the generation of gold deposition in the GAB.

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6. Conclusions The results provided by the characterization of gold mineralization in the Faina greenstone belt and the use of in-situ U-Pb-O-Hf isotope methodology in protracted magmatism in the southern GAB provide important insights that assist to establish a more comprehensive tectono-magmatic framework of this crustal fragment of still unclear origin. Archean geodynamic models for the GAB are approached by evidence for continental crust formation. Based on geochemical and isotopic data, the crustal reservoirs of the GAB seem to remain fairly consistent through time, suggesting that the older 3000 to 2800 Ma source had minor to negligible input of new material. Reworking of a deep-seated source during opening and closing of greenstone belt paleobasins may imply that associated magmatism acted as a potential source for gold systems in the Goiás Archean Block area.

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Thomson, M.L. and Fyfe, W.S., 1990. The Crixás gold deposit, Brazil: Thrustrelated postpeak metamorphic gold mineralization of possible Brasiliano cycle age. Economic Geology 85, 928– 942. Thomson, M.L., 1991. Wall-rock alteration related to Au mineralization in the low amphibolite facies: Crixás Gold Mine, Goiás, Brazil. Can. Mineral. 29, 461–480. Tomazzoli, E.R., 1985. Geologia, Petrologia, deformação e potencial aurífero do greenstone belt de Goiás - GO. Unpublished MSc Dissertation, University of Brasília, p. 206. Tomazzoli, E.R. and Nilson, A.A., 1986. Contribuição à geologia, metamorfismo e deformação do Greenstone Belt de Goiás, GO: Brazilian Congress of Geology, 34, Goiânia, Proceedings, v.2 p.615-629. Tomazzoli, E.R., 1992. A Faixa cidade de Goiás(GO): geocronologia. SBG, Brazilian Congress of Geology, São Paulo, Proceedings, v.2, p. 178-179. Tomazzoli, E.R., 1997. Aspectos geológicos e petrológicos do enxame de diques Morro Agudo de Goiás. Unpublished MSc Dissertation, University of Brasília, p. 265. Tomlinson, K.Y., Stevenson, R.K., Hughes, D.J., Hall, R.P., Thurston, P.C., Henry, P., 1998. The Red Lake green-stone belt, Superior Province: evidence of plume-relatedmagmatism at 3 Ga and evidence of an older enrichedsource. Precambrian Research, 89: 59–76. Valeriano, C.M., Machado, N., Simonetti, A., Valladares, C.S., Seer, H.J., Simões, L.S.A., 2004. U-Pb geochronology of the southern Brasília belt (SE-Brazil): sedimentary provenance, Neoproterozoic orogeny and assembly of West Gondwana. Precambrian Research, 130:27-55. Valeriano, C.M., Pimentel, M.M., Heilbron, M., Almeida, J.C.H., Trouw, R.A., 2008. Tectonic evolution of the Brasília Belt, Central Brazil, and early assembly of Godwana. Geological Society, London, Special Publications, 294:197-210. Vargas, M.C., 1992. Geologia dos granito-gnaisses dos Terrenos Granito-Greenstone da Região de Crixás, Guarinos, Pilar de Goiás e Hidrolina, Goiás. Unpublished MSc Dissertation, University of Brasília, p. 172. Van Kranendonk, M.J., Collins, W.J., Hickman, A., Pawley, M.J., 2004. Critical tests of vertical vs. horizontal tectonic models for the Archean East Pilbara GraniteGreenstone Terrane, Pilbara Craton, Western Australia. Precambrian Res. 131, 173–211. Van Kranendonk, M.J., Smithies, R.H., Hickman, A.H., Champion, D.C., 2007. Review: secular tectonic evolution of Archean continental crust: interplay between horizontal and vertical processes in the formation of the Pilbara Craton, Australia. Terra Nova 19, 1–38. Van Kranendonk, M.J., 2011. Onset of plate tectonics. Science, 333(6041): 413-414. Van Kranendonk, M.J., Kröner, A., Hoffman, E.J., Nagel, T., Anhaeusser, C.R., 2014. Just another drip: re-analysis of a proposed Mesoarchean suture from the Barberton Mountain Land, South Africa. Precambrian Res. 254, 19–35.

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Chapter 5 Conclusions and future work

This final chapter provides the key findings of each chapter followed by scientific and exploration implications for gold systems in the Faina greenstone belt and in the Goiás Archean Block (GAB). This chapter concludes with suggestions for future work which would help to solve some remaining issues.

5.1. Lithostratigraphy and structural settings in the Faina greenstone belt The Faina greenstone belt consists of a lower volcanic sequence of komatiites and tholeiitic basalts unconformably overlain by two sedimentary sequences that progress from lower conglomerate and immature sandstone to pelites until culminating with an upper chemically-precipitated sequence (e.g. limestone, BIF, chert). All lithological units comprising the Faina volcano-sedimentary sequence have been metamorphosed to greenschist facies. Initial shortening during Paleoproterozoic deformation led to the development of moderately N-dipping S1 foliation (S0 transposed into S1), regional-scale

NNW-striking F1 folds, barren S1-parallel parasitic veins and greenschist metamorphism. Progression to D2 is attested by the development of ~ E-W-striking, tight to isoclinal F2 folds and subtly S-dipping axial planar S2 foliation. Structures produced during D2 also include gently W-plunging intersection lineation between S1 and S2 foliations, and EW-trending, subtly S-dipping thrust faults that caused the stacking of the stratigraphy. The preferred orientation of quartz and gold grains, subtly plunging west, define the Lm2 mineral stretching lineation that implies a general sense of transport in the NNW-SSE direction. Deformation of early formed V1 parasitic veins during D2 is accompanied by new fluid influx reflected by the development of gold-bearing V2 veins. Subsequent deformation during the D3 event produced moderately N-dipping S3 foliation and NW-trending shear zones, which locally comprise LS- mylonites. The orientation of kinematic indicators, e.g. σ-feldspar porphyroclasts in quartzite, indicate a dextral sense of displacement. Earlier-formed structures are transposed into parallelism within D3 shear zones. Subtly NW-plunging Lm3 mineral stretching lineation indicate mass-transport in the NW-

SE direction. During the D3 event, minor hydrothermal fluid flow within V2 veins caused the precipitation of the paragenetically late galena-pyrite-stibnite±tourmaline assemblage. The D1 to D3 structures are cross-cut and locally offset (≤ m-wide) by D4 structures. The latter include moderately S- dipping thrust faults and associated sub-horizontal S-plunging slickenlines, as well as steeply W- dipping reverse faults and associated sub vertical S-plunging slickenlines. The D4-related planar features are characterized by fault-fill veins and breccias. Late D4 structures are interpreted to be associated with the Neoproterozoic Brasiliano/Pan-African orogeny. Metal endowment proposed for the D4 event, based on available whole rock and trace element data, reveal an Ag, W, Mo, Pb, Mn, U, and Zn enrichment but no significant Au concentration. Given the paucity of information concerning

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the hydrothermal alteration of D4-related structures, in addition to the apparent lack of associated gold, this system is not further investigated in the present study.

5.2. Hydrothermal alteration mineralogy and zonation of the Cascavel and Sertão gold deposits

Hydrothermal alteration surrounding folded V2 veins in the Cascavel and Sertão deposits is largely controlled by the host lithology. Quartzite and biotite schist from Cascavel developed dominant quartz-white mica-pyrite±chalcopyrite and white mica-biotite-tourmaline-siderite-pyrite assemblages, respectively. Alteration hosted in carbonaceous schist and BIF from the Sertão comprise a dominant white mica-siderite-pyrite-arsenopyrite±chalcopyrite assemblage. Increase in Fe content and pervasive K-metasomatism are evidenced towards proximal alteration zones. The chemically precipitated metasedimentary rocks at Sertão likely acted as a favourable trap for gold deposition due to the high Fe content. Gold mineralization in this deposit is accompanied by the progressive replacement of Fe-rich carbonates, e.g. ankerite, Fe-dolomite and siderite by Fe-rich sulfide assemblages, e.g. pyrite±chalcopyrite±arsenopyrite. This suggests sulfidation via redox reactions likely played an important role in gold precipitation mechanisms. Temperature estimates of gold-bearing hydrothermal fluids are between 330-420°C at Cascavel and 320-430°C at Sertão.

At Cascavel, ‘free’ gold in deformed quartz veins hosted in quartzite is mostly controlled by F2 fold hinges, with minor disseminated mineralization in hydrothermally altered wall-rock. The NW- trending D3 shear zones, locally deforming D2-related orebodies, are also associated with minor fluid influx represented by paragenetically late minerals (e.g., tourmaline, galena, stibnite).

At Sertão, refractory gold is locally present within V2 veins and disseminated sulfide-rich quartz- siderite-white mica proximal alteration zones that surround the D3 shear zones. Gold-bearing D2 structures, such as fold hinges, are locally reoriented parallel to NW-trending D3 shear zones. Gold mineralization in the Cascavel and Sertão deposits is characterized by: (i) structural control of gold mineralization via shear zones and folds, (ii) equivalent white mica±fuchsitic mica- pyrite±chalcopyrite assemblages surrounding gold-bearing structures and veins ubiquitous in proximal alteration zones of both deposits, (iii) overall increase of Fe content towards proximal alteration zones, and (iv) ore fluid temperature estimates between 330-420°C at Cascavel and 320-430°C at Sertão. Accordingly, the structural settings, hydrothermal alteration mineralogy and zonation, combined with the temperature range of hydrothermal alteration and mineralization at the Cascavel and Sertão deposits are comparatively similar to other orogenic gold systems in Brazil and worldwide (Hagemann and Cassidy, 2000; Lobato et al., 2016). Therefore, both Cascavel and Sertão deposits represent epigenetic, mesozonal orogenic gold systems. A Paleoproterozoic age for gold mineralization is consistent with the deformation history and alteration characteristics of similar gold systems in other greenstone belts of the Goiás Archean Block (Tassinari et al., 2006; Jost and Queiroz; 2008, Marques et al., 2013).

5.3. Petrogenesis of intrusive rocks in the southern Goiás Archean Block

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Isotopic and elemental measurements in granitic rocks portray remarkable secular evolution across terranes of the southern GAB. Disparate evolutionary paths favour collage of terranes amalgamated by younger tectonic processes and imply allochthonous growth of the craton (e.g. Pimentel et al., 2000; Jost and Fortes, 2001; Jost et al., 2014). Constraints on crustal growth recorded in the southern GAB assist to clarify the still uncertain origin of this cratonic fragment and reassess existing geodynamic models proposed for the area. Secular whole-rock chemistry and zircon Hf-O isotope composition highlight shifts in magma sources, depth of partial melting and differentiation. Zircon Hf isotope compositions define four major domains at: (i) 2870-2820 Ma are mostly near chondritic, with more radiogenic values with decreasing age (2820 Ma Paus de Choro granite), (ii) 2060 Ma show a steep shift to the unradiogenic Pink Syenite intrusion in the Faina greenstone belt, (iii) 630-610 Ma indicate wide range of unradiogenic compositions in K-rich granites and, (iv) 530 Ma define a steep trend to strongly unradiogenic Serra Negra granite. Conversely, oxygen isotopes in zircons are: (i) largely mantle-like until 2820 Ma, after which they rise to heavier values marking the onset of subduction, (ii) overall mantle-like at 2060 Ma, (iii) abruptly 18O-rich at ~ 600 Ma due to widespread supracrustal reworking that at last sharply switches to (iv) mantle-like values at 530 Ma. Pink Syenite intrusion marks the progression from Na-rich Archean TTGs to K-rich Neoproterozoic arc magmatism. Low REE content and La/Yb ratio of Pink Syenite are consistent with a low degree of partial melting as predicted for metasomatised slab-derived melts. For the first time, felsic intrusive magmatism is registered in the Faina greenstone belt. Following 1000 Ma hiatus, magmatism at ~ 600 Ma bordering the southern GAB is denoted by the emplacement of K-rich granites (i.e., Rio Caiapó and Itapuranga). LILE enrichment, HFSE depletion and negative Ti anomaly featured by K-rich granites are in agreement with arc-related magmatism along continental margins (Hawkesworth et al., 1993). Magmatic activity at this time is roughly simultaneous with the peak of collisional processes involving West Gondwana (~ 660-600 Ma). High Sr/Y ratio, low Y content and Rb/Sr ratio suggest subduction-related derivation of the youngest magmatic pulse (Serra Negra granite) by deep melting of a primitive source at convergent margins. Based on the chemical and isotopic evidence from the Uvá TTG (Archean) and the Serra Negra granite (Cambrian), a co-parental relation is proposed for the two magmatic events.

5.4. Tectonic-magmatic evolution and geodynamics in the GAB Archean crustal growth recorded in the GAB persisted for over 150 Ma (Pimentel et al., 1996; Queiroz et al., 2008; Beguelli Jr., 2012; Jost et al., 2013; Chapter 3). The ca. 200 Ma of near-continuous Archean juvenile magmatism is incompatible with plume-related models, which are defined by pulses on scales of tens of m.y. (Condie, 1998). Instead, relatively continuous subduction processes predict the long-lived magmatism recorded in the Goiás Archean Block. The most consistent model accounting for protracted TTG magmatism between ca. 3000-2800 Ma in the GAB, therefore, comprises subduction

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processes. Accordingly, a model proposed for the Archean encompasses: (i) early intraplate rifting; (ii) development of greenstone basins, and (iii) subsequent subduction and collision during ~ 2800 Ma and ~ 2700 Ma deformation (in the southern and northern GAB, respectively). Chemical and isotopic shifts manifested by superchondritic Hf compositions and high δ18O signature after ~ 2800 Ma (i.e. near the Meso-Neoarchean transition) implies major tectonic changes as the onset of subduction processes. Archean greenstone belts are mostly represented by komatiites and, to a minor extent, arc-like tholeiitic basalts. The paucity of arc basalts can be explained by the fore-arc and arc low preservation potential during collision (Clift and Vannucchi, 2004), which results in arc destruction and back-arc preservation. Likewise, minor Paleoproterozoic felsic magmatism is preferentially distributed along accretionary/continental margins, e.g. Pink Syenite intrusion in the Faina greenstone belt. The former shows LILE enrichment and HFSE depletion assumed to denote the contribution of an enriched source. It also displays high Sr/Y ratio and low Y content that favour garnet-bearing residuum associated with deep melting at convergent margins (Condie, 2008; Martin, 1993). Moreover, low REE contents and La/Yb ratio of this intrusion indicate a low degree of partial melting (Arth and Hanson, 2975; Arth, 1976). If elevated Rb/Sr ratios indicate an increased granitic contribution (e.g., Dhuime et al., 2012), the low Rb/Sr ratio of the Pink Syenite is an additional evidence to support the emplacement of this intrusion along deep-seated suture zones. Therefore, subduction of oceanic crust confined between the Uvá and Caiçara continental blocks is proposed to generate the intrusion of the Pink Syenite during the Paleoproterozoic orogeny. The existence of a rigid crust able to maintain topography and thereby provide the source for erosion at high elevation, result in the formation of basins after sediment deposition at low elevation. Accordingly, supracrustal material extensively available in the Neoproterozoic led to the reworking of sedimentary rocks testified in the isotopic record of K-rich granites. The transient appearance of K-rich granites, therefore, marks the onset of pervasive sedimentary reworking. Serra Negra magmatism designates ultimate tectonic shift and interaction with older, presumably Archean, source.

5.5. Links between crust-mantle evolution and mineral systems The concentration of specific mineralization types is known to be spatially and temporally controlled by steep lithospheric structures and transient geodynamic stresses (McCuaig and Hronsky, 2014; Wyman et al., 2016). Paradigm established based on the concept of mineral systems involves steep/deep structures acting as pathways to focus ore fluids/melts from a fertile source at depth, e.g. metasomatised mantle above subduction zones (McCuaig and Hronsky, 2014). Thus, recognition of these pathways and potential sources can be used as a tool for exploration targeting at a lithospheric scale (Begg et al., 2010; Hronsky et al., 2012; Mole et al., 2014). The formation of greenstone belt-hosted orogenic gold deposits is widely attributed to the final stages of ~ 50 m.y. deformation and magmatic events (e.g. Goldfarb et al., 2001). In the southern GAB, Pink Syenite intrusion brackets syn- to late-phases of regional deformation linked with Au

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mineralization (Chapter 3). By inference, the Pink Syenite may be considered as a potential host for gold mineralization, thereby interesting from an exploration point of view. Similar magmatism is recorded in greenstone belts of the northern GAB (e.g., Jost et al., 1993; Queiroz, 2000). Accordingly, syn-collisional gold mineralization events occur in other areas affected by the Paleoproterozoic orogeny during the assembly of Atlantica, e.g. West Africa, Amazon and São Francisco cratons, and the Guiana Shield (Teixeira et al., 2007; Petersson et al., 2017; Eglinger et al., 2017; Tedeschi, 2018). Gold and minor orthomagmatic mineral deposits (e.g. the Boa Vista Ni deposit at the NW edge of the Crixás greenstone belt; Costa Jr. et al., 1997) tend to cluster along craton margins with TTG terranes and adjacent to major tectonic unconformities in greenstone belt sequences. Orebodies comprising the 7 Moz Crixás deposits in the northern GAB, for example, are mostly located near the contact between graphite-rich metasedimentary and metavolcanic rocks (Fortes, 1996; Jost et al., 2010). However, taken that several margins seem to be devoid of mineralization, the need for datasets combining detailed tectonic, geological and geophysical constraints is emphasized. One of the most prominent structural features in the southern GAB, the Faina fault, has been interpreted as a splay from the Moiporá-Novo Brasil shear zone limiting the GAB to the west (Resende et al., 1998). The northeast striking Faina fault, which offsets the two southernmost greenstone belts (Faina and Goiás), corresponds spatially to changes in age, chemistry and isotopic signature. The irregular distribution of sedimentary rocks suggests the latter was a growth fault likely active before opening and formation of greenstone basins in the Paleoproterozoic. A deformation corridor that is located parallel to the Faina fault hosts the Sertão gold deposit, therefore, ore-forming fluids likely exploited this early structure during subsequent deformation events. This is illustrated by undeformed mafic dikes intrusive in the Sertão deposit area (e.g. Chapter 2). The concentration and distribution of mineral deposits in greenstone belts of the GAB can potentially be conjectured by internal back-arc settings. Back arcs superimposing metasomatised lithospheric crust may provide the metal fertility required for the development of gold systems and it also may explain the occurrence of other mineralization types in the GAB, e.g. VHMS in the Crixás greenstone belt (Jost et al., 2014).

5.6. Comparisons with Archean cratons worldwide Other Archean cratons appear to record similar isotopic, chemical and stratigraphic evolution to that of the Goiás Archean Block. Examples include the Dharwar craton (Lancaster et al., 2015; Moyen et al., 2003), Slave craton (Kusky et al., 2013), Pilbara craton (Van Kranandonk et al., 2010), and Superior craton (Björkman, 2017). Comparisons with additional data provided in this thesis led to a better understanding of the nature of geological processes in parallel with the crust-mantle evolution, e.g. the compositional change at the Archean Proterozoic boundary (Keller and Schoene, 2012; Martin and Sigmarsson, 2010). Evidence presented in this study support connection of the GAB with the highly metal endowed São Francisco craton. Links with the latter, known as the oldest craton in the South

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American Platform, would have several practical applications, e.g., it enhances prospective exploration targeting and, therefore, may boost important discoveries in the GAB. However, more research focusing on deciphering the tectonic history of this area should be undertaken to better constrain this hypothesis.

6. Future work The GAB represents one of the most potential areas for gold exploration in central Brazil. Similarities with the São Francisco craton reinforce the potential for future gold discoveries. Nevertheless, the Goiás Archean Block still lacks a more detailed and systematic geological framework linking gold mineralization to the structural architecture and deformation history, particularly the under- constrained connection between the northern and southern domains of the craton. Deciphering the nature of geological, isotopic and geophysical boundaries is fundamental for a better understanding of the crustal evolution and geodynamic models for the GAB. In order to build this framework, further progress is proposed focusing on three main topics involving: (1) regional crustal and structural architecture, (2) isotope geochemistry and geochronology of intrusive rocks, and (3) nature of ore- forming fluids and timing of gold mineralization.

6.1. Regional crustal and structural architecture: Local current and historical exploration targeting by many companies involved drilling areas previously known for artisanal mining. Despite the apparent success of this approach (e.g., Cascavel, Sertão, Pilar and Mina III mines), accurate and detailed geology-based strategies are required for a more effective exploration targeting. Thus, recognition of structures, their kinematics, timing, isotopic signature and relation to hydrothermal alteration and gold mineralization are essential for future exploration success and a better understanding of the gold potential in the GAB.

6.2. Isotope geochemistry and geochronology of intrusive rocks: Limited exposure (particularly for Archean TTG terranes), extensive weathering soil profiles and dense vegetation cover in some areas hinder a broader constraint of distinct magmatic events and better definition of the contacts of these igneous bodies. Hence, complementary isotopic and geochemical investigations may help define the nature of cryptic localities, e.g. boundary between the northern and southern GAB, as well as their fertility in terms of mineral systems.

6.3. Nature of ore-forming fluids and timing of gold mineralization in the southern GAB: One of the questions that remain unanswered at present relates to the absolute age dating of gold mineralization in the Faina greenstone belt. Attempts provided in this thesis were unsuccessful in covering this knowledge gap. The effort of applying the Re-Os methodology to bracket timing of gold mineralization in the Sertão deposit was hampered by unfeasibility of arsenopyrite analysed (see Appendix of this chapter for details). Therefore, further geochronological investigations are required to constrain the age

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of the hydrothermal alteration and mineralizing systems in the Faina greenstone belt. At the deposit scale, additional studies focusing on the precise P-T-X conditions and evolution of Au-related hydrothermal fluids (e.g. combined fluid inclusion, laser ICP-MS analyses, and oxygen-hydrogen isotope studies) of greenstone belt-hosted gold systems in the southern GAB are also suggested. Detailed work should aim for the characterization of structural features and associated hydrothermal systems developed during Neoproterozoic deformation within the GAB.

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Appendices

Chapter 2 Table A2.1. Representative WDS microprobe analyses of hydrothermal silicates in the Cascavel deposit

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Table A2.1. Representative WDS microprobe analyses of hydrothermal silicates in the Cascavel deposit (continued)

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Table A2.2. Microprobe analyses of hydrothermal chlorite at the Cascavel deposit

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Table A2.3. Representative WDS microprobe analyses of sulfide minerals and gold at Cascavel deposit

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Table A2.4. Microprobe analyses of hydrothermal chlorite at the Sertão deposit

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Table A2.5. Representative WDS microprobe analyses of hydrothermal silicates from the Sertão deposit

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Table A2.6. Representative WDS analyses of gold and sulfides from the Sertão deposit

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Chapter 3

Methodology

Oxygen analyses were performed using a 20 keV Cs+ ion beam with an intensity of 3 nA in aperture illumination mode and a spot diameter of about 15µm. A normal incidence electron gun was used to avoid sample charging. The 16O and 18O isotopes were simultaneously monitored with Faraday detectors with 1010 Ω and 1011 Ω (channels L’2 and H’2, respectively), at a mass resolution of ~2500. The drift and stability of the analysis were monitored using Temora 2 zircon (δ18O = 8.2‰ VSMOW; Black et al., 2003), which on average yield a reproducibility of about 0.5‰ (2SD). The instrumental mass fractionation (IMF) was corrected using the Temora 2 zircon according to the procedure proposed by Kita et al. (2009). The Penglai zircon (δ18O = 5.3 ± 0.1‰; Li et al., 2010) was used as a secondary standard in order to monitor the accuracy of the measured isotopic ratios, which yield an average value of about 0.3‰ (2SD). The uncertainty of δ18O was calculated by propagating the errors on instrumental mass fractionation determination and an internal error on each sample spot analysis. Hafnium isotope analysis was conducted by laser ablation multi-collector ICP-MS, using a Thermo Neptune Plus multi-collector ICP-MS combined with 193 nm Ar-F excimer laser sampling system in the School of Earth Sciences at UWA. Analytical conditions included laser energy of ∼5 J/cm2, ablation rate of 4 Hz over an ablation period of 60 s. Spots of 25, 35, 40 and 50 µm diameter were employed, depending on the size of the grains and the zone of interest. Analytical protocols used in the laboratory follow Kemp et al. (2009). The data reproducibility was evaluated by quality control analysis using the reference zircons Penglai, OGC, FC1 and Mud Tank zircon. The mean 176Hf/177Hf values and associated 2σ uncertainties (standard deviations) determined for the standard reference materials are: Penglai 0.282910 ± 0.000022 (0.282906 ± 0.0000010; Li et al., 2010), OGC 0.280636 ± 0.000019 (0.280633 ± 34; Kemp et al., 2017), FC1 0.282184 ± 0.000019 (0.282184±16; Woodhead and Hergt, 2005), and Mud Tank Carbonatite 0.282488 ± 0.000012 (0.282506 ± 0.00026; Woodhead and Hergt, 2005). All data are reported relative to the solution 176Hf/177Hf value established for Mud Tank zircon. The εHf values for analysed sample zircons were calculated using a 176Lu decay constant of 1.865 x 10-11 yr-1 (Scherer et al., 2001), and the chondritic values of Bouvier et al. (2008).

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Table A3.1. Mean weighted analyses of standard materials used for U-Pb dating

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Table A3.2. Results of standard materials used for the Lu-Hf analyses

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Table A3.2. Results of standard materials used for the Lu-Hf analyses (continued)

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Table A3.3. Results of standard materials used to calibrate in-situ oxygen analyses

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Table A3.3. Results of standard materials used to calibrate in-situ oxygen analyses (continued)

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Fig. A3.1. Distribution in bi-log diagrams of compatible (Ni, Co, V, Sr) versus incompatible (Rb, Ba) elements indicate that most of Neoproterozoic samples evolved mostly by partial melting, rather than by fractional crystallization. A close relationship is observed between the Uvá and Paus de Choro samples. The striking feature of this data is the lack of correlation between the Serra Negra granite and other Neoproterozoic samples, but rather with Archean samples, e.g., Uvá and Paus de Choro, suggesting mostly partial melting (B and D), with minor fractional crystallization (A and C) can be invoked to indicate co-parental features between these magmas.

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Fig. A3.2. Plots of εHft of the magmas versus geochemical signatures. Grey field outlines the composition of TTGs based on Moyen (2011). Near chondritic εHf of Archean rocks contrasts from lower values in Paleo- and Neoproterozoic times. (A) K2O/Na2O ratio and εHf values of the Uvá and Caiçara TTGs are consistent with reference. The K-rich composition of the Paus de Choro granite agrees with origin from TTG melts and two-mica mineralogy. Neoproterozoic granites have high K2O/Na2O ratios, excluding the Serra Negra granite. (B) A general increase in Th content is observable from Archean-Paleoproterozoic to Neoproterozoic samples, with exception of the Serra Negra granite, which plots within the TTG field. (C) High Sr/Y ratios of the Serra Negra granite, Pink Syenite and Caiçara orthogneiss denote deep source for their genesis.

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Fig. A3.3. Plots of δ18O versus U-Pb dataset. The correlation between lower δ18O with high discordance (A) and Th/U ratios (B) indicate that the oxygen values are a product of secondary alteration instead of magmatic features (please see text).

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Table A3.4. Whole-rock geochemistry results of major, trace and REE analyses

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Table A3.4. Whole-rock geochemistry results of major, trace and REE analyses (continued)

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Fig. A3.5. CL images of Itapuranga I zircons used for in-situ U-Pb, Hf-O study. Circles indicate U-Pb (black), Hf (yellow) and O (orange) analyses for all samples

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Fig. A3.6. CL images of Itapuranga II zircons used for in-situ U-Pb, Hf-O study.

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Fig. A3.7. CL images of Rio Caiapó zircons analysed in the in-situ U-Pb, Hf-O study

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Fig. A3.8. CL images of Serra Negra zircons analysed in the in-situ U-Pb, Hf-O study

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Fig. A3.9. CL images of Pink Syenite zircons analysed in the in-situ U-Pb, Hf-O study

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Fig. A3.9. CL images of Pink Syenite zircons analysed in the in-situ U-Pb, Hf-O study (continued)

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Fig. A3.10. CL images of Paus de Choro zircons analysed in the in-situ U-Pb, Hf-O study

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Fig. A3.11. CL images of Uvá zircons analysed in the in-situ U-Pb, Hf-O study

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Fig. A3.12. CL images of Caiçara zircons analysed in the in-situ U-Pb, Hf-O study

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Figure A3.13. Weighted average plots of zircon U-Pb dating

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Figure A3.13. Weighted average plots of zircon U-Pb dating (continued)

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Figure A3.14. Results of SHRIMP U-Pb analyses on zircon

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Table A3.5. Results of SHRIMP U-Pb analyses on zircon

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Table A3.5. Results of SHRIMP U-Pb analyses on zircon (continued)

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Table A3.5. Results of SHRIMP U-Pb analyses on zircon (continued)

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Table A3.5. Results of SHRIMP U-Pb analyses on zircon (continued)

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Table A3.6. Results of zircon SIMS O analyses on zircon Drift corrected or raw Raw data from CIPS SIMS corrected ratios Final delta Analysis n (- to +) ratios if no drift 18O/16O σ int in rel.% 18O/16O σ abs 18O/16O σ abs δ18O 2σ abs O_JB1_IT1-1 38.5 0.002022 0.009203565 0.002022 1.86146E-07 0.002022 3.1674E-07 8.42 0.32 O_JB1_IT1-11 51.5 0.002024 0.01320273 0.002024 2.67271E-07 0.002024 3.70424E-07 9.33 0.37 O_JB1_IT1-12 52.5 0.002023 0.01283957 0.002023 2.59806E-07 0.002023 3.64997E-07 8.89 0.36 O_JB1_IT1-14 53.5 0.002023 0.0111877 0.002023 2.26368E-07 0.002023 3.41999E-07 8.83 0.34 O_JB1_IT1-15 60.5 0.002024 0.01386126 0.002024 2.80571E-07 0.002024 3.80108E-07 9.21 0.38 O_JB1_IT1-16 61.5 0.002023 0.01172686 0.002023 2.3726E-07 0.002023 3.49287E-07 8.75 0.35 O_JB1_IT1-17 62.5 0.002023 0.01257243 0.002023 2.54332E-07 0.002022 3.61072E-07 8.61 0.36 O_JB1_IT1-18 63.5 0.002024 0.01528924 0.002023 3.09393E-07 0.002023 4.01803E-07 8.95 0.40 O_JB1_IT1-19 64.5 0.002023 0.01352192 0.002023 2.73547E-07 0.002023 3.74855E-07 8.64 0.37 O_JB1_IT1-3 39.5 0.002024 0.01793955 0.002024 3.63112E-07 0.002024 4.44515E-07 9.21 0.44 O_JB1_IT1-4 40.5 0.002024 0.01527482 0.002024 3.09148E-07 0.002023 4.01642E-07 9.12 0.40 O_JB1_IT1-5 41.5 0.002024 0.008248435 0.002024 1.66924E-07 0.002023 3.05964E-07 8.99 0.31 O_JB1_IT1-6 42.5 0.002023 0.01284799 0.002022 2.59862E-07 0.002022 3.64958E-07 8.45 0.36 O_JB1_IT1-7 49.5 0.002023 0.01108509 0.002023 2.24232E-07 0.002022 3.40539E-07 8.56 0.34 O_JB1_IT1-9 50.5 0.002023 0.0163529 0.002023 3.30809E-07 0.002022 4.18461E-07 8.63 0.42 O_JB1_SC-1 -60.5 0.002026 0.008471634 0.002026 1.71668E-07 0.002026 3.08875E-07 10.40 0.31 O_JB1_SC-10 -47.5 0.002025 0.01501712 0.002025 3.04161E-07 0.002025 3.97952E-07 9.94 0.40 O_JB1_SC-11 -46.5 0.002026 0.01196784 0.002026 2.42455E-07 0.002026 3.53096E-07 10.16 0.35 O_JB1_SC-12 -45.5 0.002025 0.008705449 0.002025 1.76321E-07 0.002025 3.1138E-07 9.91 0.31 O_JB1_SC-13 -44.5 0.002025 0.01560132 0.002025 3.15944E-07 0.002025 4.07E-07 9.78 0.41 O_JB1_SC-14 -43.5 0.002025 0.01031888 0.002025 2.08965E-07 0.002025 3.30924E-07 9.75 0.33 O_JB1_SC-15r -34.5 0.002027 0.01252575 0.002027 2.53901E-07 0.002027 3.61151E-07 10.72 0.36 O_JB1_SC-16r -33.5 0.002027 0.008395617 0.002027 1.70168E-07 0.002027 3.08091E-07 10.62 0.31 O_JB1_SC-17 -32.5 0.002026 0.01322666 0.002026 2.68025E-07 0.002026 3.71157E-07 10.41 0.37 O_JB1_SC-18 -31.5 0.002025 0.01178179 0.002025 2.38637E-07 0.002025 3.50447E-07 9.94 0.35 O_JB1_SC-19r -30.5 0.002026 0.01405627 0.002026 2.84732E-07 0.002025 3.8333E-07 10.04 0.38 O_JB1_SC-6 -59.5 0.002026 0.01612181 0.002026 3.26621E-07 0.002026 4.15409E-07 10.21 0.41 O_JB1_SC-7 -58.5 0.002026 0.0101706 0.002026 2.0604E-07 0.002026 3.29162E-07 10.14 0.33 O_JB1_SC-8 -57.5 0.002026 0.01040748 0.002026 2.1088E-07 0.002026 3.32252E-07 10.34 0.33 O_JB1_SC-9 -56.5 0.002025 0.01422538 0.002025 2.8807E-07 0.002025 3.85765E-07 9.75 0.38 O_JB1_SN-10 3.5 0.002017 0.01265172 0.002017 2.55196E-07 0.002017 3.61164E-07 5.74 0.36 O_JB1_SN-11 4.5 0.002017 0.01189971 0.002017 2.40064E-07 0.002017 3.50666E-07 5.89 0.35 O_JB1_SN-1c -20.5 0.002017 0.01012852 0.002017 2.04289E-07 0.002017 3.27185E-07 5.69 0.33 O_JB1_SN-1r -21.5 0.002017 0.0143588 0.002017 2.89562E-07 0.002016 3.86169E-07 5.52 0.39 O_JB1_SN-2r -19.5 0.002018 0.00973898 0.002018 1.96512E-07 0.002017 3.22469E-07 6.09 0.32 O_JB1_SN-3 -18.5 0.002017 0.0112716 0.002017 2.274E-07 0.002017 3.42134E-07 5.94 0.34 O_JB1_SN-4 -17.5 0.002017 0.008943218 0.002017 1.80394E-07 0.002017 3.1284E-07 5.75 0.31 O_JB1_SN-5 -9.5 0.002018 0.02101983 0.002018 4.24109E-07 0.002017 4.95161E-07 6.05 0.49 O_JB1_SN-6c -7.5 0.002017 0.01158539 0.002017 2.337E-07 0.002017 3.46325E-07 5.80 0.35 O_JB1_SN-6r -8.5 0.002017 0.01358798 0.002017 2.74013E-07 0.002016 3.7465E-07 5.50 0.37 O_JB1_SN-7 -6.5 0.002017 0.01118849 0.002017 2.25661E-07 0.002017 3.40927E-07 5.65 0.34 O_JB1_SN-8 -5.5 0.002018 0.0140871 0.002018 2.84224E-07 0.002017 3.82267E-07 6.01 0.38 O_JB1_SN-9 2.5 0.002017 0.008832083 0.002017 1.78161E-07 0.002017 3.11564E-07 5.78 0.31

210

Table A3.6. Results of zircon SIMS O analyses on zircon (continued) Drift corrected or raw Raw data from CIPS SIMS corrected ratios Final delta Analysis n (- to +) ratios if no drift 18O/16O σ int in rel.% 18O/16O σ abs 18O/16O σ abs δ18O 2σ abs O_JB1_IT2-1 5.5 0.002023 0.0115469 0.002023 2.33618E-07 0.002023 3.46831E-07 8.79 0.35 O_JB1_IT2-13 28.5 0.002023 0.007276496 0.002023 1.47244E-07 0.002023 2.95674E-07 8.91 0.29 O_JB1_IT2-17 29.5 0.002024 0.01340511 0.002024 2.71332E-07 0.002024 3.73344E-07 9.22 0.37 O_JB1_IT2-18 30.5 0.002023 0.007148219 0.002023 1.44633E-07 0.002023 2.94359E-07 8.80 0.29 O_JB1_IT2-2 6.5 0.002023 0.01212514 0.002023 2.45271E-07 0.002022 3.54748E-07 8.60 0.35 O_JB1_IT2-3 14.5 0.002024 0.009685827 0.002024 1.96038E-07 0.002024 3.22797E-07 9.15 0.32 O_JB1_IT2-4 15.5 0.002023 0.0125228 0.002023 2.53376E-07 0.002023 3.6044E-07 8.83 0.36 O_JB1_IT2-5 16.5 0.002024 0.008234242 0.002024 1.66651E-07 0.002023 3.05836E-07 9.09 0.31 O_JB1_IT2-6 17.5 0.002023 0.0121047 0.002023 2.4493E-07 0.002023 3.54566E-07 8.89 0.35 O_JB1_IT2-7 18.5 0.002023 0.009889422 0.002023 2.00112E-07 0.002023 3.25239E-07 8.91 0.32 O_JB1_IT2-8 26.5 0.002024 0.01153125 0.002024 2.33348E-07 0.002023 3.46685E-07 8.97 0.35 O_JB1_IT2-9 27.5 0.002024 0.0138314 0.002024 2.79953E-07 0.002024 3.79649E-07 9.19 0.38 O_JB2_PC-10 -36.5 0.002018 0.01176285 0.002018 2.3735E-07 0.002018 3.68343E-07 6.23 0.37 O_JB2_PC-11c -35.5 0.002018 0.007499881 0.002018 1.51371E-07 0.002018 3.19829E-07 6.45 0.32 O_JB2_PC-11r -34.5 0.002013 0.007331798 0.002013 1.47594E-07 0.002013 3.1741E-07 3.83 0.32 O_JB2_PC-12c -33.5 0.002018 0.01211477 0.002018 2.44437E-07 0.002018 3.72936E-07 6.17 0.37 O_JB2_PC-12r -32.5 0.002019 0.01152521 0.002019 2.32705E-07 0.002019 3.65507E-07 6.88 0.36 O_JB2_PC-13 -28.5 0.002018 0.01495253 0.002018 3.01768E-07 0.002018 4.12835E-07 6.43 0.41 O_JB2_PC-16c -26.5 0.00202 0.01177535 0.00202 2.37817E-07 0.00202 3.68836E-07 7.13 0.37 O_JB2_PC-16r -27.5 0.002011 0.01207036 0.002011 2.42734E-07 0.002011 3.71114E-07 2.84 0.37 O_JB2_PC-2 -51.5 0.002017 0.009842419 0.002017 1.98555E-07 0.002017 3.44571E-07 6.00 0.34 O_JB2_PC-20c -25.5 0.002013 0.0131429 0.002013 2.64561E-07 0.002013 3.85945E-07 3.82 0.38 O_JB2_PC-20r -24.5 0.002018 0.008393583 0.002018 1.69352E-07 0.002017 3.28642E-07 6.12 0.33 O_JB2_PC-21c -17.5 0.002019 0.01289307 0.002019 2.60313E-07 0.002019 3.83664E-07 6.84 0.38 O_JB2_PC-21r -20.5 0.002019 0.00944828 0.002019 1.90726E-07 0.002018 3.40267E-07 6.63 0.34 O_JB2_PC-22c -16.5 0.002018 0.01546007 0.002018 3.11975E-07 0.002018 4.20332E-07 6.31 0.42 O_JB2_PC-22r -15.5 0.002018 0.01177491 0.002018 2.37566E-07 0.002017 3.68454E-07 6.10 0.37 O_JB2_PC-3c -50.5 0.002017 0.01129768 0.002017 2.27852E-07 0.002017 3.62189E-07 5.74 0.36 O_JB2_PC-4 -49.5 0.002018 0.01613801 0.002018 3.25623E-07 0.002018 4.30542E-07 6.22 0.43 O_JB2_PC-5 -48.5 0.00202 0.01570464 0.00202 3.17204E-07 0.00202 4.24403E-07 7.25 0.42 O_JB2_PC-6 -44.5 0.002018 0.01570729 0.002018 3.16913E-07 0.002018 4.23982E-07 6.16 0.42 O_JB2_PC-7 -43.5 0.00202 0.009477845 0.00202 1.91425E-07 0.00202 3.40786E-07 7.17 0.34 O_JB2_PC-8 -42.5 0.00202 0.01103109 0.00202 2.22801E-07 0.00202 3.59355E-07 7.21 0.36 O_JB2_PC-9c -40.5 0.00202 0.01137911 0.00202 2.29839E-07 0.00202 3.63769E-07 7.25 0.36 O_JB2_PC-9r -41.5 0.002011 0.008485952 0.002011 1.7066E-07 0.002011 3.28539E-07 2.87 0.33 O_JB2_UV-1c 18.5 0.002016 0.01261637 0.002016 2.54412E-07 0.002016 3.79424E-07 5.58 0.38 O_JB2_UV-1r 19.5 0.002033 0.01262899 0.002033 2.56769E-07 0.002033 3.82728E-07 13.88 0.38 O_JB2_UV-2c 20.5 0.002017 0.01182157 0.002017 2.38499E-07 0.002017 3.69046E-07 6.05 0.37 O_JB2_UV-2r 21.5 0.002018 0.007447574 0.002018 1.50294E-07 0.002018 3.19279E-07 6.28 0.32 O_JB2_UV-3 22.5 0.002012 0.0159697 0.002012 3.21321E-07 0.002012 4.26771E-07 3.36 0.43 O_JB2_UV-4 25.5 0.002015 0.01275184 0.002015 2.5695E-07 0.002015 3.80973E-07 4.82 0.38 O_JB2_UV-5c 26.5 0.002018 0.009004437 0.002018 1.81677E-07 0.002017 3.35156E-07 6.11 0.33 O_JB2_UV-5r 27.5 0.002031 0.009469368 0.002031 1.92352E-07 0.002031 3.42639E-07 12.93 0.34 O_JB2_UV-6 28.5 0.002028 0.0245463 0.002028 4.97701E-07 0.002027 5.72543E-07 11.12 0.57 O_JB2_UV-7 29.5 0.002015 0.01063424 0.002015 2.14315E-07 0.002015 3.53657E-07 4.97 0.35 O_JB2_UV-8c 33.5 0.002016 0.01390208 0.002016 2.80335E-07 0.002016 3.97267E-07 5.56 0.40 O_JB2_UV-8r 32.5 0.002015 0.01082493 0.002014 2.18076E-07 0.002014 3.55866E-07 4.59 0.35 O_JB2_UV-9 34.5 0.002015 0.01061139 0.002015 2.13862E-07 0.002015 3.5339E-07 5.00 0.35

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Table A3.6. Results of zircon SIMS O analyses on zircon (continued) Drift corrected or raw Raw data from CIPS SIMS corrected ratios Final delta Analysis n (- to +) ratios if no drift 18O/16O σ int in rel.% 18O/16O σ abs 18O/16O σ abs δ18O 2σ abs O_JB2_CC-11c 50.5 0.00201 0.0217815 0.00201 4.3785E-07 0.00201 5.20043E-07 2.42 0.52 O_JB2_CC-11r 51.5 0.00201 0.01386872 0.00201 2.78758E-07 0.00201 3.95509E-07 2.30 0.39 O_JB2_CC-12c 53.5 0.002009 0.01750639 0.002009 3.51649E-07 0.002009 4.49749E-07 1.67 0.45 O_JB2_CC-12r 52.5 0.002028 0.01373018 0.002028 2.78402E-07 0.002028 3.97015E-07 11.12 0.40 O_JB2_CC-16c 54.5 0.002015 0.01742445 0.002015 3.51147E-07 0.002015 4.49929E-07 4.94 0.45 O_JB2_CC-16r 55.5 0.002012 0.01083482 0.002012 2.1803E-07 0.002012 3.55584E-07 3.45 0.35 O_JB2_CC-3c 35.5 0.002011 0.01095593 0.002011 2.2034E-07 0.002011 3.56881E-07 2.88 0.36 O_JB2_CC-3r 36.5 0.002009 0.01186941 0.002009 2.38496E-07 0.002009 3.68174E-07 1.98 0.37 O_JB2_CC-4c 39.5 0.002012 0.01211809 0.002012 2.43779E-07 0.002012 3.71867E-07 3.16 0.37 O_JB2_CC-4r 40.5 0.002012 0.01776568 0.002012 3.57526E-07 0.002012 4.54685E-07 3.55 0.45 O_JB2_CC-5c 42.5 0.00201 0.0154573 0.00201 3.1063E-07 0.002009 4.18548E-07 2.12 0.42 O_JB2_CC-5r 41.5 0.002014 0.0102614 0.002014 2.06655E-07 0.002014 3.48907E-07 4.25 0.35 O_JB2_CC-6c 43.5 0.002017 0.01005364 0.002017 2.02751E-07 0.002016 3.46923E-07 5.63 0.35 O_JB2_CC-6r 46.5 0.002015 0.01041899 0.002015 2.09907E-07 0.002014 3.50927E-07 4.62 0.35 O_JB2_CC-7 47.5 0.002011 0.01313275 0.002011 2.64138E-07 0.002011 3.85479E-07 2.96 0.38 O_JB2_CC-9c 48.5 0.002016 0.007434363 0.002016 1.49924E-07 0.002016 3.18931E-07 5.57 0.32 O_JB2_CC-9r 49.5 0.002012 0.01355884 0.002012 2.72863E-07 0.002012 3.91623E-07 3.53 0.39 O_JB2_SY-1 -14.5 0.00202 0.01679216 0.00202 3.39185E-07 0.00202 4.41078E-07 7.29 0.44 O_JB2_SY-12 1.5 0.002015 0.01365708 0.002015 2.75256E-07 0.002015 3.93601E-07 5.07 0.39 O_JB2_SY-13_23 4.5 0.002018 0.01746625 0.002018 3.52532E-07 0.002018 4.51286E-07 6.51 0.45 O_JB2_SY-14 5.5 0.002018 0.01470558 0.002018 2.96731E-07 0.002018 4.09131E-07 6.23 0.41 O_JB2_SY-18c 6.5 0.002018 0.009198484 0.002018 1.85675E-07 0.002018 3.37448E-07 6.57 0.34 O_JB2_SY-18r 7.5 0.002033 0.01219892 0.002033 2.47968E-07 0.002033 3.76833E-07 13.65 0.38 O_JB2_SY-19c_21c 8.5 0.002017 0.01011257 0.002017 2.04008E-07 0.002017 3.4774E-07 5.99 0.35 O_JB2_SY-19r_21r 11.5 0.002031 0.0127942 0.002031 2.59816E-07 0.002031 3.8453E-07 12.67 0.38 O_JB2_SY-2 -10.5 0.002017 0.008924097 0.002017 1.79993E-07 0.002017 3.34167E-07 5.77 0.33 O_JB2_SY-20_19 12.5 0.002014 0.01095451 0.002014 2.20596E-07 0.002014 3.57326E-07 4.19 0.36 O_JB2_SY-21_16 13.5 0.002022 0.01360736 0.002022 2.7514E-07 0.002022 3.94169E-07 8.31 0.39 O_JB2_SY-22_14 15 14.5 0.002016 0.01668645 0.002016 3.36356E-07 0.002016 4.38531E-07 5.20 0.44 O_JB2_SY-23c_13 15.5 0.002017 0.01303627 0.002017 2.62991E-07 0.002017 3.85316E-07 6.01 0.38 O_JB2_SY-3 -9.5 0.002018 0.01009675 0.002018 2.03737E-07 0.002018 3.47637E-07 6.24 0.35 O_JB2_SY-4 -8.5 0.002019 0.01034251 0.002019 2.08786E-07 0.002019 3.50717E-07 6.67 0.35 O_JB2_SY-5 -7.5 0.002019 0.01312388 0.002019 2.64915E-07 0.002018 3.86756E-07 6.61 0.39 O_JB2_SY-6 -6.5 0.002018 0.01210859 0.002018 2.44372E-07 0.002018 3.72942E-07 6.41 0.37 O_JB2_SY-7c -2.5 0.002017 0.01775592 0.002017 3.5816E-07 0.002017 4.55589E-07 5.90 0.45 O_JB2_SY-7r -1.5 0.002018 0.01428963 0.002018 2.88387E-07 0.002018 4.03155E-07 6.41 0.40 O_JB2_SY-8 -0.5 0.002013 0.0123937 0.002013 2.49464E-07 0.002013 3.7574E-07 3.74 0.37 O_JB2_SY-9 0.5 0.002018 0.0102068 0.002018 2.05976E-07 0.002018 3.48973E-07 6.32 0.35

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Chapter 4

Table A4.1. Previous deformation history of the GAB

213

Chapter 5

Arsenopyrite analyses Petrography, SEM and EPMA In order to conduct arsenopyrite Re-Os geochronology, three ore-bearing samples from the Sertão deposit were selected for detailed investigation. Microscale textural and mineralogical relationships were defined by transmitted and reflected optical petrography supported by semi-quantitative analyses using the Vega 3 Tescan Scanning Electron Microprobe (SEM-EDS) at the Centre for Microscopy, Characterisation and Analysis (CMCA) at UWA. Images were taken in secondary electron (SE) and backscatter (BSE) modes to identify internal discontinuities, i.e. fractures and/or inclusions, growth patterns. An accelerating voltage of 20kV, a working distance of 15 mm, and a beam intensity of 1.5 nA were used to obtain the images. Electron dispersive spectroscopy (EDS) analyses utilized the Oxford Instruments AZTEC® analytical suite. Qualitative data was collected from electron probe micro- analyzer (EPMA) analyses using the JEOL JXA-8530F Hyperprobe fitter with five wavelength- dispersive spectrometers (WDS) at the Centre for Microscopy, Characterisation and Analysis (CMCA). Operating conditions for WDS analyses included an accelerating voltage of 20 kV, a beam current of 40 nA, and a counting time of 40 s on peak. Mean atomic number and ZAF corrections were used throughout analyses.

214

Fig. A5.1. Hand sample (A, B) of carbonaceous schist used for LA-ICP-MS analyses (STO5:128). Typical hydrothermal assemblage is illustrated by transmitted (C, F) and reflected (D, E) optical photomicrographs.

215

Fig. A5.2. Compositional zoning evidenced by EPMA elemental maps of arsenopyrite hosted in carbonaceous schist (A, B and E) and BIF (C and D).

216

Laser ablation inductively-coupled plasma mass spectrometry (LA-ICP-MS) on arsenopyrite Arsenopyrite samples representative of gold mineralization in the Sertão deposit (STO5:128) were drilled out of thin section block and prepared in four 1-inch diameter epoxy mounts. The mounts were polished and labelled as plug 1, 2, 3 and 4 (see Fig. A5.3. for details). This investigation aimed to underpin arsenopyrite chemical zonation and test the suitability of the grains for the Re-Os method. Evaluation of the Re concentration and possible correlation with Au distribution were conducted via multi-element LA-ICP-MS spot analysis and elemental mapping of arsenopyrite. These were carried out on an Agilent HP 7700 laser coupled to a HP 4500 quadrupole inductively-coupled plasma mass spectrometer (ICP-MS) at CODES Analytical Laboratories, UTAS (University of Tasmania). Analytical procedures followed the methodology proposed by Danyushevsky et al. (2011). A total of four elemental mapping LA-ICP-MS analyses were carried out on arsenopyrite from the Sertão deposit, and the results are presented in Table A5.1 and illustrated in Fig. A5.4 to A5.7. Arsenopyrite grains were ablated on 16-51 µm diameter spots using the laser at 1-10 Hz and energy density of 2.7 J/cm2. All data reduction calculations and error propagations were undertaken within the software LADR designed at CODES Analytical Laboratories, UTAS. Standard and unknowns analyses were used to compare values for each element and to correct for differences in spot size between the primary. Measured isotopes included 34S, 57Fe, 59Co, 60Ni, 65Cu, 86Zn, 75As, 77Se, 95Mo, 121Sb, 185Re, 187Re, 197Au, 208Pb, and 209Bi.

The dataset reveals common S-depleted rims that are enriched in Co, Ni, Sb, As, Sb and, to a minor extent, Cu, Pb, Mo, Au and Re. Negligible Zn and Se, as well as variable Fe and Bi contents, are reported for all samples. Poor correlation prescribes distribution of Mo and Re. However, no clear correlation link Au and Re concentrations. Arsenopyrite grains exhibit low Re contents with abundances ranging from 10-2 to 10-1 ppm. Variable Au contents have abundance ranging from 100 to 103 ppm. Gold distribution at the micron- to the nanoscale is consistent with Au either in the lattice or as nanoparticles (e.g. Ciobanu et al., 2011, 2012).

217

Fig. A5.3. SEM-BSE images of plugs from sample STO 5:128 for arsenopyrite LA-ICP-MS work. Elemental mapping analyses are indicated in red.

218

Fig. A5.4. LA-ICP-MS elemental mapping with ppm concentration measured for Plug 1, circle #5

219

Fig. A5.5. LA-ICP-MS elemental mapping with ppm concentration measured for Plug 3, circle #7.

220

Fig. A5.6. LA-ICP-MS elemental mapping with ppm concentration measured for Plug 3, circle #8.

221

Fig. A5.7. LA-ICP-MS elemental mapping with ppm concentration measured for Plug 3, circle #16.

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Table A5.1. Results of arsenopyrite LA-ICP-MS analyses Filename Sample Minerals Primary Spot Frequency Energy Fluence Spot ICPMS Comment Name (one or more) Secondary Size Line method eg.pyrite, quartz Sample (μm) (Hz) (mJ) (J/cm2) Image eg lazirc D18JU20b001 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b002 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b003 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b004 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b005 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b006 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b007 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b008 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b009 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b010 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b011 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b012 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b013 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b014 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b015 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b016 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b017 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b018 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b019 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b020 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b021 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b022 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b023 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b024 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b025 3-7_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20b026 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b027 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b028 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b029 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b030 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20b031 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c001 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c002 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c003 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c004 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c005 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T

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Table A5.1. Results of arsenopyrite LA-ICP-MS analyses (continued) Filename Sample Minerals Primary Spot Frequency Energy Fluence Spot ICPMS Comment Name (one or more) Secondary Size Line method eg.pyrite, quartz Sample (μm) (Hz) (mJ) (J/cm2) Image eg lazirc D18JU20c006 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c007 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c008 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c009 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c010 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c011 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c012 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c013 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c014 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c015 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c016 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c017 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c018 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c019 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c020 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c021 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c022 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c023 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c024 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c025 1-5_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c026 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c027 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c028 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c029 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c030 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c031 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c032 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c033 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c034 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c035 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c036 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c037 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c038 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c039 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c040 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c041 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c042 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c043 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c044 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c045 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T

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Table A5.1. Results of arsenopyrite LA-ICP-MS analyses (continued) Filename Sample Minerals Primary Spot Frequency Energy Fluence Spot ICPMS Comment Name (one or more) Secondary Size Line method eg.pyrite, quartz Sample (μm) (Hz) (mJ) (J/cm2) Image eg lazirc D18JU20c046 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c047 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c048 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c049 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c050 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c051 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c052 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c053 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c054 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c055 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c056 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c057 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c058 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c059 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c060 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c061 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c062 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c063 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c064 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c065 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c066 3-8_AspyMap arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c067 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c068 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c069 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c070 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c071 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c072 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c073 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c074 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c075 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T

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Table A5.1. Results of arsenopyrite LA-ICP-MS analyses (continued) Filename Sample Minerals Primary Spot Frequency Energy Fluence Spot ICPMS Comment Name (one or more) Secondary Size Line method eg.pyrite, quartz Sample (μm) (Hz) (mJ) (J/cm2) Image eg lazirc D18JU20c076 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c077 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c078 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c079 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c080 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c081 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c082 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c083 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c084 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c085 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c086 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c087 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c088 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c089 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c090 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c091 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c092 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c093 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c094 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c095 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c096 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c097 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c098 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c099 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c100 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c101 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c102 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c103 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c104 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c105 img3-16 arsenopyrite Sample 16µm 10Hz 100mJ 2.7 image gs short 12.5%T D18JU20c106 smp blk gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c107 BLANK gas blank 51µm 1Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c108 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c109 GSD-1G glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c110 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T D18JU20c111 STDGL3 glass Primary 51µm 10Hz 100mJ 2.7 Spot gs short 12.5%T

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