Petrogenesis and Tectono-Magmatic Evolution of S-Type and A-Type Granites in the New England Batholith

Petrogenesis and Tectono-magmatic Evolution of S-type and A-type granites in the New England Batholith

A Thesis submitted to The University of Newcastle for the Degree of Doctor of Philosophy

B. Landenberger B.Sc. (Hons) July 1996

Frontispiece: View looking east from the Devil’s Forehead (near the summit of Chaelundi Mountain in the Guy Fawkes River National Park) over the headwaters of Chandlers Creek in Chaelundi State Forest. The foreground outcrop forms part of the Chaelundi Complex A- type granite suite. I hereby certify that the work embodied in this thesis is the result of original research and has not been submitted for a higher degree to any other University or Institution.

(Signed)______ACKNOWLEDGMENTS

Several people have lent helping hands during the course of this project with technical and logistical problems, while others have helped to prevent the onset of insanity. Bill Collins and Robin Offler supervised the project and have been invaluable help in clarifying ideas and helping with organisational aspects of the thesis work. Bill Collins provided fervent inspiration in times when things seemed to be going in circles. Discussions with Paul Dirks, Sue Keay, Terry Farrell, Martin Hand, Dick Flood, Stirling Shaw, and Ron Vernon have also aided in developing the ideas presented herein.

Many staff members of the Geology Department have provided technical assistance and advice, including Richard Bale, Esad Krupic, Hope Ruming, Geraldene MacKenzie and Jan Crawford. I also thank Doug Todd for his help with XRF analysis and Dave Phelan for his help with microprobe analysis. Isotopic analyses were carried out at the Centre for Isotope Studies (CIS - North Ryde) under the guidance of Dave Whitford, Bo Zhou and Steve Craven. Stereographic projections were produced with the computer program GeOrient which was kindly supplied by Rod Holcombe (Department of Earth Sciences, University of Queensland).

During the first three years of the project, I was supported by an Australian Postgraduate Research Award. Funding for external analytical work was supported by several ARC and University (SRC) research grants to Bill Collins and Robin Offler.

Finally, I would like to thank my wife Debbie for her help, support, and most of all patience and understanding over the last six years.

For Debbie, Justine and Daniel -i-

TABLE OF CONTENTS

LIST OF FIGURES...... v

LIST OF PLATES...... ix

LIST OF TABLES...... x

ABSTRACT...... xi

CHAPTER 1. INTRODUCTION ...... 1 1A1 Preamble - Physiography of the New England Tableland...... 1 1A2 Regional geological setting ...... 2 1A3 A Brief Tectonic History ...... 4 1A3A1 Pre- late Carboniferous ...... 4 1A3A2 Post- late Carboniferous ...... 5 1A4 Previous studies of the New England Batholith ...... 6 1A5 Aims of the project...... 8

CHAPTER 2. STRUCTURAL RELATIONS OF THE HILLGROVE PLUTONIC SUITE 10 2A1 Introduction ...... 10 2A2 Structural Setting ...... 12 2A2A1 Distribution of the Hillgrove Suite ...... 12 2A2A2 Structural history of the Tia Complex...... 13 2A3 Structural Framework...... 14 2A3A1 A definition of domains...... 15 2A3A2 Structural continuity outside the Tia Complex - Macroscopic structural features...... 16 2A3A3 Meso- and Microscopic Deformation Features...... 19 2A3A4 A correlation of post-accretion deformation structures - the regional perspective...... 32 2A4 Conclusions and Regional Implications...... 43 -ii-

CHAPTER 3. AGE RELATIONS OF THE HILLGROVE PLUTONIC SUITE ...... 47 3A1 Introduction ...... 47 3A2 Previous geochronology of the ...... 49 3A3 Structural sequence associated with the Hillgrove Suite - a brief review ..... 53 3A4 Sample Selection, Analytical Methods ...... 54 3A4A1 Zircon U-Pb...... 54 3A4A2 Biotite Rb-Sr ...... 57 3A6 Discussion...... 61

CHAPTER 4. PETROGENESIS OF THE HILLGROVE PLUTONIC SUITE...... 66 4A1 Introduction ...... 66 4A1A1 S-type granites in the SNEFB ...... 67 4A1A2 Previous petrogenetic studies of the Hillgrove Suite ...... 67 4A2 Defining the Hillgrove Suite ...... 68 4A2A1 Previous definitions of the Hillgrove Suite ...... 68 4A2A2 Definition of the Hillgrove Supersuite ...... 71 4A3 Petrography & Mineral Chemistry ...... 78 4A3A1 General Petrography of Hillgrove Supersuite granitoids ...... 78 4A3A2 Biotite ...... 82 4A3A3 Other ferromagnesian phases: Amphiboles & pyroxenes...... 85 4A3A4 Peraluminous phases other than biotite: garnet, cordierite, muscovite...... 87 4A3A5 Oxides and other opaque phases ...... 90 4A3A6 Plagioclase...... 91 4 A3A7 Alkali Feldspar...... 92 4A3A8 Quartz...... 93 4A4 Geochemistry ...... 96 4A4A1 Defining the geochemical characteristics - variations within the Hillgrove Supersuite, and comparisons with other granitoids...... 96 4A4A2 The Bakers Creek Suite...... 110 4A4A3 Rocks of accretion complex ...... 114 4A5 Petrogenetic Constraints...... 115 4A5A1 Parental magmas of the Hillgrove Suite ...... 115 -iii-

4A5A2 Identifying the source rocks...... 118 4A5A3 Quantification the partial melting process...... 133 4A5A4 Contamination within the Hillgrove Supersuite - the possibilities...... 140 4A6 Conclusions...... 143

CHAPTER 5. PETROGENESIS OF THE CHAELUNDI COMPLEX A-TYPE GRANITOID SUITE: DERIVATION BY PARTIAL MELTING OF A DEHYDRATED CHARNOCKITIC LOWER CRUST...... 144 5A1 Introduction ...... 144 5A2 Geological Setting ...... 145 5A3 Petrography of the Suites ...... 149 5A3A1 The I-type suite ...... 149 5A3A2 The A-type suite...... 155 5A4 Geochemistry ...... 156 5A5 Petrogenetic Constraints...... 161 5A5A1 Fractionation ...... 161 5A5A2 Parental magmas ...... 172 5A5A3 Geochemical contrasts between the Chaelundi Complex I- and A-type parental magmas...... 172 5A5A4 Nature of the source rocks...... 174 5A6 Petrogenetic Models...... 175 5A6A1 Previous Models ...... 175 5A6A2 Charnockitization of the lower crust and generation of A-type magmas176 5A7 Conclusions...... 184

CHAPTER 6. PETROGENESIS OF BASALTIC ENCLAVES IN THE A-TYPE WOODLANDS QUARTZ MONZONITE ...... 185 6A1 Introduction ...... 185 6A2 Geological Setting ...... 186 6A3 Microgranitoid enclaves in A-type granites of the New England Batholith ...... 188 6A4 Field relationships of the Woodlands Quartz Monzonite and its enclaves .... 188 -iv-

6A4A1 Age and intrusive relationships...... 188 6A4A2 Outcrop detail ...... 190 6@5 Petrography and Mineral Chemistry ...... 194 6A5A1 The host quartz monzonite...... 194 6A5A2 Enclave matrix...... 194 6A5A3 Enclave phenocrysts ...... 197 6A5A4 Enclave xenocrysts ...... 202 6A5A5 Pyroxene thermometry...... 203 6A6 Geochemistry ...... 204 6A6A1 The host quartz monzonite...... 204 6A6A2 Enclaves...... 209 6@7 Enclave Petrogenesis ...... 210 6A7A1 Restite, Cumulate or Mingled Magma? ...... 210 6A7A2 Composition of the enclave parent magma...... 215 6A7A3 Post-emplacement chemical evolution of enclave magmas ...... 218 6A8 Conclusions...... 231 6A8A1 A model of enclave formation for the Woodlands example ...... 231 6A8A2 General implications for the origin of microgranitoid enclaves .... 236

CHAPTER 7. A TECTONIC SYNTHESIS...... 238

REFERENCES ...... 242

APPENDICES...... 266 Appendix A: Analytical methods...... 267 Appendix B: Structural data...... 270 Appendix C: Modal data for granitoids...... 273 Appendix D: Mineral analyses (Electron-microprobe)...... 277 Appendix E: Whole-rock geochemical analyses...... 290 Appendix F: Isotopic data ...... 304 Appendix G: Catalogued sample list with grid references...... 306 -v-

LIST OF FIGURES

Figure 1A1. Simplified map of the southern New England Fold Belt...... 3 Figure 2A1. Areal distribution of the Hillgrove Suite and defined structural domains...... 11

Figure 2A2. D3 form-surface map for accretion complex metasediments...... 34

Figure 2A3. D5 form-surface map for accretion complex metasediments...... 36 5 Figure 2A4. Contoured stereoplot of poles to S5 overlain with L5 stretching lineations...... 37

7 Figure 2A5. Stereoplot of poles to D7 mylonite shear (C) planes overlain with L7 stretching lineations...... 37

Figure 2A6. Map of D7 shear zones in Hillgrove Suite granitoids and metasediments. . . 40 Figure 2A7. Unwinding the New England orocline...... 45 Figure 3A1. Geological map of southern New England Fold Belt...... 48 Figure 3A2. (a) Standard isochron plot of whole-rock Rb-Sr data for Hillgrove Suite granitoids. (b) Best isochron plot of whole-rock Rb-Sr data for Hillgrove Suite granitoids...... 51 Figure 3A3. Concordia plot of zircon U-Pb data from the Tia Granodiorite and Abroi Granodiorite...... 55 Figure 3A4. Map of the Hillgrove Plutonic Suite and associated Bakers Creek Suite with sample locations and Rb-Sr biotite ages ...... 59 Figure 3A5. Contour map of the biotite Rb/Sr ages from Hillgrove Suite granitoids not

mylonitized during D7...... 63 Figure 3A6. Schematic cross-sections of the Wollomombi Zone...... 65 Figure 4A1. Map of plutons of the Hillgrove Supersuite...... 73 Figure 4A2. (a) Streckheisen plot of Hillgrove Supersuite and Bundarra Suite granitoids. (b) Modal variation within the Hillgrove suite...... 79 Figure 4A3. Biotite compositions of Hillgrove Supersuite and Bundarra Suite granitoids...... 84 Figure 4A4. (a) Classification of Rockisle Suite amphiboles. (b) Rockisle Suite amphibole compositions and Bakers Creek pyroxene compositions plotted in the pyroxene quadrilateral...... 86 -vi-

Figure 4A5. Garnet, cordierite, and ilmenite compositions from the Hillgrove Suite and Bundarra Suite plotted in terms of ternary Fe-Mg-Mn contents...... 89 Figure 4A6. Ternary plot of feldspar compositions for Hillgrove Supersuite and Bundarra Suite granitoids...... 89 Figure 4A7. Major element Harker plots...... 98 Figure 4A8. Trace element Harker plots...... 101 Figure 4A9. Molecular Ca-Na-K and ACF plots...... 104 Figure 4A10 Inter-element ratio Harker plots...... 105 Figure 4A11. REE spidergram...... 109 Figure 4A12. AFM and tectonic discrimination plots for gabbros of the Bakers Creek Suite.112 Figure 4A13. Q-Ab-Or and An-Ab-Or plots...... 124 Figure 4A14. Batch partial melting trends...... 127

87 86 Figure 4A15. eNd vs Sr/ Sri plot for granitoids and gabbros of the Hillgrove Supersuite and accretion complex metasediments...... 131 Figure 4A16. Harker plots of residuum compositions calculated by major element modelling.138 Figure 5A1. (a) Geological map of the southern New England Fold Belt showing the outcrop area of the granitoids younger than 240 ma...... 146 Figure 5A1. (b) Geological map of the Chaelundi Complex ...... 148 Figure 5A2. Streckheisen plot of representative modes from the two suites of the Chaelundi Complex...... 150 Figure 5A3. Plot of biotite compositions for the A-type suite...... 152 Figure 5A4. Plot of hornblende compositions for the A-type suite...... 153 Figure 5A5. Ternary feldspar plot for the A-type suite...... 154 Figure 5A6. Selected major elements plots...... 157 Figure 5A7. Selected trace elements plots...... 158 Figure 5A8. REE - chondrite normalized plots ...... 160 Figure 5A9. Plot illustrating the combined effect of plagioclase and orthoclase fractionation on Ba and Eu contents within the A-type suite...... 168 Figure 5A10. Selected plots of inter-element ratios vs silica...... 171 Figure 5A11. Y vs Nb tectonic discrimination diagram for both suites ...... 173 -vii-

Figure 5A12. Ce/Nb vs Y/Nb discrimination plot for the two major subgroups of A-type granites ...... 173 Figure 5A13. Schematic representation of partial melting processes during I-type magma ...... 180 Figure 5A14. Spidergram plots comparing the A-types and average I-type of the Chaelundi Complex with C-type magmas...... 183 Figure 6A1. Geological map showing the distribution of leucoadamellites and related rocks in the southern New England Fold Belt...... 187 Figure 6A2. Geological map of the Kookabookra - Wards Mistake area ...... 189 Figure 6A3. Mineral chemistry plots for enclaves of the Woodlands Quartz Monzonite and host monzonite...... 198 Figure 6A4. Pyroxene quadrilateral plot of enclave clinopyroxene and orthopyroxene phenocrysts, and their secondary alteration products...... 199 Figure 6A5. Classification of both primary matrix amphibole and secondary amphiboles after clinopyroxene phenocrysts ...... 199 Figure 6A6. Ternary plot of feldspar compositions from the Woodlands Quartz Monzonite and enclaves...... 200 Figure 6A7. Pyroxene thermometry of enclave phenocrysts from the Woodlands Quartz Monzonite...... 200 Figure 6A8. Major element plots for Woodlands Quartz Monzonite and enclaves...... 205 Figure 6A9. Trace element plots for Woodlands Quartz Monzonite and enclaves...... 206 Figure 6A10. Inter-element ratio plots for Woodlands Quartz Monzonite and enclaves...... 207 Figure 6A11. REE chondrite normalized plot for Woodlands Quartz Monzonite and enclaves.208 Figure 6A12. gNd vs initial 87Sr/86Sr plot of isotopic analyses of the Woodlands Quartz Monzonite, enclaves, and a single enclave clinopyroxene separate...... 214 Figure 6A13. Tectonic discrimination diagrams for Woodlands enclaves...... 217 Figure 6A14. Mg# vs XCa and ASI vs Mg# plots showing the relationship between bulk enclave compositions, enclave matrix and phenocrysts...... 228 Figure 6A15. Schematic representation of the three stage model for formation of the Woodlands enclaves...... 233 -viii-

Figure 7A1(a-c). Schematic profiles of the southern New England Fold Belt from early Carboniferous to late Carboniferous...... 239 Figure 7A1(d-e). Schematic profiles of the southern New England Fold Belt from early Permian to middle Triassic...... 241 -ix-

LIST OF PLATES

Plate 2A1. D5 macro- and mesoscopic features of metasediments from the Wollomombi Zone.20

Plate 2A2. D5 microstructures in metasediments...... 22

Plate 2A3. Mesoscopic features of D5 D7 and D8 structures in Hillgrove Suite granitoids. 25

Plate 2A4(a-d). Microscopic features of D7 in Hillgrove Suite granitoids...... 29

Plate 2A4(e-h). Microscopic features of D7 in Hillgrove Suite granitoids...... 30 Plate 4A1. Photomicrographs of Hillgrove Supersuite granitoids and associated rocks...... 81 Plate 5A1. Photomicrographs comparing the microstructure of the mafic end-members of the two suites ...... 151 Plate 6A1. Outcrop features of enclaves in the Woodlands Quartz Monzonite...... 191 Plate 6A2. Outcrop detail of mafic microgranitoid enclaves...... 193 Plate 6A3(a-d). Woodlands Quartz Monzonite - Enclave microstructure...... 195 Plate 6A3(e-h). Woodlands Quartz Monzonite - Enclave microstructure...... 196 -x-

LIST OF TABLES

Table 1A1 Simplified summary of the New England Batholith...... 7 Table 2A1. Distinguishing characteristics of accretion complex and Permian basin metasediments...... 27 Table 2A2. Comparative timing of structural and magmatic events throughout the Tablelands Complex...... 33 Table 3A1 Summary of previous geochronological work on the Hillgrove Suite and associated rocks...... 50 Table 3A2. Biotite Rb-Sr data from Hillgrove Suite granitoids and high grade metasediments. 60 Table 4A1. Characteristics of S- and I-type granites...... 66 Table 4A2. Characteristics of Hillgrove Supersuite and Bakers Creek Suite members...... 72 Table 4A3 Distinctive mineralogy of the Hillgrove, Rockisle and Bundarra plutonic suites ...... 80 Table 4A4. Comparative major and trace element chemistry for LFB I-type, LFB S-type and New England Batholith I-type granites, with the Hillgrove, Rockisle and Bundarra Suite granites...... 107 Table 4A5. Comparison of the average analysis of Bakers Creek Suite gabbros, with mid-ocean -ridge, island arc calc-alkaline and island arc tholeiitic basalts...... 111 Table 4A6. Major element modelling...... 137 Table 4A7. Trace Element Modelling ...... 140 Table 5A1. Comparative mineralogy and geochemistry of the I- and A-type suites...... 156 Table 5A2 Major element modellingA ...... 163 Table 5A3. Trace element Modelling ...... 165 Table 5A4. Rare-earth-element Modelling ...... 167 Table 6A1. Summary of isotopic results from the Woodlands Quartz monzonite and enclaves 213 Table 6A2. Modelling results - Woodlands enclaves ...... 222 Table 6A3. Comparison of average bulk (whole-rock, XRF) enclave compositions and matching average matrix compositions...... 226 -xi-

ABSTRACT

The late Carboniferous heralded a fundamental change in tectonic and magmatic styles within the southern New England Fold Belt (SNEFB). Westward migration of arc magmatism during the early and middle Carboniferous, was ensued by rapid easterly migration during the late Carboniferous, establishing a new arc within the original accretion complex (Tablelands Complex) of the SNEFB. This event was accompanied by the onset of high-T/low-P metamorphism and uplift in parts of the accretion complex.

Subduction-accretion fabrics developed earlier in the Carboniferous (D1-D2) were initially overprinted by D3 (~311 Ma) during major uplift of the southern Tia Complex. Biotite grade D3 fabrics were in turn folded during D4. This thermal perturbation culminated with intrusion of granitoids of the Hillgrove Supersuite and gabbros and diorites of the Bakers Creek Suite (of island arc tholeiite affinity) in the latest Carboniferous. Intrusion of these arc magmas was accompanied by further compressional deformation (D5) causing uplift of the entire Wollomombi Zone along its eastern margin. Preliminary zircon U-Pb ages presented herein, provide the first tight constraints on the emplacement and crystallization of the Hillgrove Supersuite granitoids, and also the D5 deformation, establishing a latest Carboniferous age.

Geochemical characteristics, combined with Sr and Nd isotopic compositions of granitoids and various potential source rocks, provide tight constraints on possible sources for the Hillgrove Supersuite granitoids. Only volcanogenic greywackes of intermediate composition (~65% SiO2) overlap with the isotopic composition of the granitoids, inferring that these metasediments are the most likely source. Calculated melt fertilities of various potential source rocks (based on the proportional component of the ternary Q-Ab-Or minimum melt composition at 5 kb) also indicate that these intermediate greywackes are the most likely sources to produce large volumes of partial melt. Significantly, the isotopic characteristics and calculated melt fertilities preclude the involvement of pelites and felsic greywackes (~70%SiO2), which have previously been inferred as granitoid sources. The isotopic and chemical immaturity of these sediments (87Sr/86Sr 0A7048 to 0A7070, gNd +2 to

-1, high Na2O and low ASI), explains the unusual character of Hillgrove Supersuite granitoids, which are isotopically primitive (87Sr/86Sr 0A7040 to 0A7065, gNd -1 to +4), only mildly peraluminous (ASI 1A00 - 1A15), and relatively high in Na2O (3 - 4%) compared to -xii- most S-types. Major and trace element modelling indicate that the more mafic magmas (68-

70% SiO2) of the suite were produced by ~48% partial melting of the intermediate greywacke source, under water-undersaturated conditions involving biotite breakdown at granulite facies conditions and mid crustal depths (~5 kb).

Hillgrove Suite, with higher CaO, Al2O3, TiO2 and lower K2O, ΣFeO contents. These granites plot on an isotopic mixing curve between intermediate greywacke (87Sr/86Sr 0A7048 to 0A7070, gNd +2 to -1) and coeval gabbros of the Bakers Creek Suite (87Sr/86Sr 0A7027, gNd +9.5). Accordingly, mantle-derived magmas are considered to have been a contributor to the most primitive granites. Another possible minor magma source are seawater-altered metabasalts, which are common in the deeper parts of the accretion complex. These metabasalts are likely to have undergone small degrees of partial melting (via amphibole breakdown), contributing a minor melt component to the primary S-types, thus causing an isotopic shift towards higher gNd and higher initial 87Sr/86Sr.

After the D5 event and intrusion of the Hillgrove Supersuite, compressional tectonics gave way to rifting in the early Permian, with the development of early Permian basins such as the Nambucca and Manning basins within the Tablelands Complex, and the Sydney Basin further to the west. During the early and middle Permian, there was little manifestation of arc magmatism within the SNEFB. However, intrusion of the Bundarra Plutonic Suite, and extrusion of bimodal volcanics in the rift basin sequences, probably represent inter-arc or back-arc rifting during establishment of an intra-oceanic island arc further to the east.

Compressional tectonics dominated the late Permian, with climactic deformation during the Hunter-Bowen Orogeny. This event initially generated large-scale folding of earlier fabrics within the accretion complex (F6). On a regional scale, this folding was related to development of the Texas - Coffs Harbour Orocline and dispersal of discrete structural blocks within the Tablelands Complex. F6 folds in the Wollomombi Zone and S1 cleavage in the adjacent early Permian Nambucca Block, are both truncated by D7 ductile shear zones -xiii- which represent the culmination of this deformation. D7 involved westward tilting of the entire Wollomombi Zone with up to 8 km of uplift, along mylonite zones which truncated many plutons of the Hillgrove Supersuite.

Rb-Sr dating of biotite from high-grade D7 mylonite zones constrain the age of D7 uplift to the late Permian (258-266 Ma). Crustal tilting during D7 has also produced the large range of Rb-Sr biotite ages for Hillgrove Supersuite granitoids not directly affected by D7 mylonitization. Slow thermal relaxation after the high-T/low-P event which accompanied intrusion of the Hillgrove Supersuite, combined with crustal tilting at ~260 Ma, produced a pattern of biotite ages which decrease from west to east, as the major late Permian shear zones are approached. The oldest biotite ages in this range (296 Ma) are within error of the age of granitoid intrusion, while the youngest ages (257 Ma) record the age of uplift.

This climactic deformation was followed by re-establishment of arc-related volcanism in the early Triassic, which involved minor crustal extension and major plutonism, with intrusion of I- and A-type granites of the New England Batholith. I- and A-type granites of the Chaelundi Complex were generated at this time, in a subduction-related tectonic setting. Although isotopic ages of the suites are indistinguishable (233-235 Ma), field relations indicate that the A-type is younger. The most mafic granitoids from each suite have similar silica contents (66-68% SiO2), slightly LREE enriched patterns without Eu anomalies, low Rb/Sr and K/Ba ratios, and high K/Rb ratios, suggesting that both represent parental magmas. The A-type is distinguished mineralogically by abundant orthoclase and sodic plagioclase (total >60%), ferro-hornblende, annite and allanite. In contrast, the I-type has more hornblende and biotite, which are more magnesian in composition, and less feldspar. The parental magmas of both suites have many similar geochemical characteristics, although the A-type has slightly higher alkalis, Zr, Hf, Zn and LREE, and lower CaO, MgO, Sr, V, Cr, Ni and Fe3+/ΣFe. The geochemical features characteristic of leucocratic A-type granites, such as high Ga/Al, Nb, Y, HREE and F contents, are only manifest in the more felsic members of the A-type suite. These features were produced by ~70% fractional crystallization of feldspar, hornblende, quartz and biotite.

Both granite suites were generated by water-undersaturated partial melting of a similar source, but the A-type parent magma resulted from lower a conditions during partial H2O -xiv- melting. Generation and rapid ascent of the earlier I-type magma during disequilibrium partial melting produced a relatively anhydrous, but not refractory, charnockitic lower crust. Continued thermal input from mantle-derived magmas, during ongoing subduction, partially melted the ‘charnockitized’ lower crust at temperatures in excess of 900EC, to produce A-type magmas. As the I- and A-type granites intruded penecontemporaneously, a tonalitic source model for genesis of the Chaelundi A-type, is untenable.

Basaltic enclaves preserved in the nearby Woodlands Quartz Monzonite, also of A-type affinity, provide evidence that basaltic magmas of island arc affinity were still providing the heat source necessary for partial melting in the lower crust during the Triassic. Combined petrographic, geochemical and isotopic data provide unequivocal evidence that the enclaves present in the Woodlands Quartz Monzonite, originated as a coeval basaltic magma, that mingled with the host quartz monzonite. The preserved basaltic phenocryst assemblage (augite + hypersthene + calcic plagioclase) and the geochemical character of the enclaves, suggest that the parent magma was of high-alumina basalt affinity. Variations in chemistry and mineralogy of the enclave suite are the result of several magmatic differentiation processes which have affected the original basaltic magma. Modelling suggests that internal fractional crystallization was the primary process responsible for differentiation of the enclave suite, together with concomitant processes of physical exchange with the host granitoid. These processes include diffusional exchange at the molecular scale, as well as exchange of phenocrysts, and minor late metasomatism.

The unique preservation of the basaltic phenocryst assemblage and the basaltic geochemical character of these enclaves are the result of arrested hybridism. Although the enclaves preserved in the Woodlands Quartz Monzonite are an unusual example, their similarities to enclaves present in other granitoid types (particularly I-types) in the New England Batholith and elsewhere, suggest that this model may be applied to many examples of microgranitoid enclaves. 1

CHAPTER 1. INTRODUCTION

1A1 Preamble - Physiography of the New England Tableland.

The New England region of northeastern New South Wales is dominated by a tableland which, in the Armidale - Walcha area averages an elevation of approximately 1000 metres above sea level. This plateau slopes gradually westwards, merging with the western slopes and plains of central NSW. In contrast, the eastern descent to the coastal plains is marked by the abrupt ‘great escarpment’, where local relief is generally of the order of 600 metres or greater.

John Oxley, the Colonial Surveyor-General, was the first white man to set eyes on the Macleay gorges which form the eastern escarpment of the New England tableland in the area to the southeast of Armidale. On reaching the brink of the gorge country near Moona Plains in the spring of 1818 he recorded that:

“The whole indeed is a great natural spectacle, and is an indubitable mark of the vast convulsions which this country must at one period have undergone.” (Oxley 1826).

Although the eastern escarpment of the New England plateau is largely a product of Tertiary uplift associated with the Tasman Sea opening, together with subsequent erosion and incision by eastward flowing drainage (Ollier 1982a), Oxley’s assertion that this part of the country had undergone ‘vast convulsions’, was indeed apt.

The southern New England Fold Belt (SNEFB) is a major tectonic element in eastern Australia and dominates the New England tableland of northeastern New South Wales. Prior to Tertiary uplift associated with breakup of the Gondwana supercontinent, the SNEFB had protracted tectonic and magmatic histories stretching from the middle Devonian to the late Triassic. However, the dominant phase of deformation and magmatic activity affecting the central New England tablelands began in the late Carboniferous. Climactic deformation occurred in the late Permian (Landenberger et al. 1995) and compressional tectonics had essentially ceased by the early Triassic, while ‘post-orogenic’ arc-related igneous activity continued to at least the late Triassic (Landenberger & Collins 1996). 2

Breakup of the Gondwana supercontinent began during the Cretaceous (Veevers et al. 1993), and involved the rifting of eastern Australia to form the Tasman Basin, along with uplift of the eastern Australian highlands (Ollier 1982). This uplift and subsequent erosion were the dominant processes that produced the present physiographic features of the New England tableland.

1A2 Regional geological setting

The New England Orogen, as defined by Day et al. (1978) refers to the eastern part of the Tasman Orogenic Zone. Three provinces have been recognized (see inset, Fig. 1A1); the Yarrol in Queensland (also referred to as the northern New England Orogen or Fold Belt), the Gympie in southeastern Queensland (the ‘Gympie province’), and the New England in northeastern NSW (also referred to as the southern New England Orogen or Fold Belt). The term southern New England Fold Belt (SNEFB) will be used herein in reference to the New England province of the New England Orogen.

Prior to the application of modern plate tectonic theory, interpretation of the fold belt was largely based on the geosynclinal concept, and the terms miogeosyncline and eugeosyncline were applied (e.g. Voisey 1959) to subdivisions within the SNEFB with different tectonic styles. Leitch (1974) was the first to associate these subdivisions with an arc - forearc - subduction zone tectonic regime. Two basic subdivisions (Leitch 1974) are separated by the Peel - Manning Fault System (PMFS) which is a major tectonic feature of the SNEFB (Fig. 1@1). This suture was originally termed the Great Serpentinite Belt. The subdivisions applied by Leitch (1974) were: the Tamworth Belt (or ‘Zone A’ - west of the PMFS) which was interpreted as an arc and forearc system, while the Tablelands Complex (or ‘Zone B’ - east of the PMFS) was interpreted as a subduction/accretion complex (Fig. 1@1). Although this original subdivision was somewhat simplistic and has since been refined, there is still general consensus of this interpretation.

The SNEFB is bounded to the south and south-west by the Permo-Triassic Sydney and Gunnedah Basins, and to the north and northwest by the Clarence-Moreton Basin and Great Australian Basin which are Triassic to Cretaceous in age (Fig. 1@1). 3

Gympie Province

Yarrol Province

Great Australian Tablelands Complex New England Province Basin New England Fold Belt

Clarence-

Moreton N Basin B C

Tamworth Belt A

T M Lorne F Basin Pacific Sydney Basin S Ocean

Peripheral basins 0 50 100 Late Permian - Triassic intrusives and extrusives km Early-Middle Permian intrusives Early Permian rift basins Figure 1. 1. Simplified map of the southern New Late Carboniferous England Fold Belt showing the areas involved in intrusives this study. Inset "A" is the area relevant to chapters 2, 3 and 4, inset "B" is relevant to Accretionary prism chapter 5 and inset "C" is relevant to chapter 6. metasediments HMT = Hunter-Mooki Thrust. Arc & fore-arc basin volcanics PMFS = Peel-Manning Fault System. sequences

} Pre-Permian & metasediments 4

1A3 A Brief Tectonic History

1A3A1 Pre- late Carboniferous

The Devonian - late Carboniferous development of the southern New England Fold Belt is widely accepted as incorporating a convergent margin plate boundary, with a magmatic arc in the west flanked by a fore-arc basin, and a subduction complex in the east (e.g. Murray et al. 1987). A major change in the tectonic pattern occurred in the late Carboniferous (Collins et al. 1993), and so this summary is divided into pre - and post - late Carboniferous sections.

Serpentinites within the PMFS contain the oldest rocks in the SNEFB. The serpentinites contain various exotic blocks of blueschist, eclogite, metabasites and plagiogranites, with the latter having Cambrian U-Pb zircon ages (Aitchison 1992). Despite these relatively old ages, most rocks in the SNEFB are Devonian and younger.

The relatively intact stratigraphic sequences of the Tamworth Belt range from middle Devonian to Permian in age on the basis of in situ faunas (Flood & Aitchison 1988). The style of volcanism within the belt changed from intra-oceanic island arc (basaltic to andesitic) in the Devonian to continental margin arc (dacitic to rhyolitic) in the middle Carboniferous (McPhie 1987, Roberts & Engel 1987), and finally to an early Permian bimodal volcanism associated with rifting (Leitch 1988).

In contrast to the Tamworth Belt, sequences in the Tablelands Complex are severely disrupted (e.g. Collins et al. 1993), partly as a result of subduction/accretion involving dismemberment of stratigraphic sequences. The tectono-stratigraphic sequence youngs dominantly eastward, although the younging direction within each tectonic package is generally westward (e.g. Korsch 1977). This disruption is further complicated by several phases post-accretion deformation involving rearrangement of structural blocks and oroclinal bending. The age of these tectono-stratigraphic sequences has only recently been clarified by radiolarian biostratigraphy (e.g. Aitchison 1990) and ranges from late Devonian to late Carboniferous.

Korsch (1977) subdivided the accretion complex sequences into three broad lithological and 5 chronological categories, which were termed ‘associations’. In chronological order, these are:

(I) the Woolomin Association (?late Devonian) which is dominated by chert, jasper, pelites and metabasic rocks along with rarer sandstones, argillites and minor limestone;

(ii) the Sandon Association (?early Carboniferous) which is a lithic sandstone- rich unit with abundant chert, metabasites and rare limestones;

(iii) the Coffs Harbour Association (late Carboniferous) which is dominated by volcanolithic greywackes, siliceous siltstones and pelites, while cherts and metabasic rocks are extremely rare, and limestones are absent.

This progression from an ocean floor metabasite and pelagic sediment dominated sequence, to sequences dominated by arc derived detritus, is generally thought to reflect an increasing supply rate of arc detritus from the forearc which was emergent by the late Carboniferous (e.g. Fergusson 1985). In conjunction with this progression, is a change in the nature of the arc detritus from andesitic in the Woolomin Association, to dacitic in the Sandon Association, to rhyodacitic in the Coffs Harbour Association, reflecting geochemical evolution of the arc source area (Korsch 1984).

The metamorphic grade of exposed rocks within the subduction/accretion complex is relatively low - reaching only prehnite-pumpellyite or lower greenschist facies (Offler & Hand 1988), although restricted occurrences of blueschist facies rocks do occur (e.g. in the southern part of the Tia Complex - Hand 1988).

1A3A2 Post- late Carboniferous

In the late Carboniferous a dramatic change in tectonic style occurred. Subduction was terminated in an intense orogenic event associated with high T/low P metamorphism and granitic magmatism within the Tablelands Complex. The metamorphism and progressive deformation associated with this event are best recorded in the Wongwibinda and Tia 6 metamorphic complexes which reached a peak metamorphic grade of upper amphibolite facies at this time. Granitoids of the Hillgrove Plutonic Suite also intruded at this time (~300 Ma - Collins et al. 1993) and occur in the core of these metamorphic complexes, as well as at higher crustal levels. The underlying cause for this abrupt change is still a matter of speculation, but there is a general acceptance that the original Carboniferous magmatic arc had migrated eastwards, placing the subduction/accretion complex in an arc or back-arc setting (e.g. Collins et al. 1993).

Intrusion of the Bundarra Plutonic Suite (~285 Ma - Collins et al. 1993) in the early Permian was immediately followed by widespread rifting and bimodal volcanism associated with formation of the Barnard Basin (Leitch 1988) which includes the Nambucca Block and several other smaller basins which lie unconformably on, or are faulted against, older rocks of the fold belt. Climactic deformation ensued in the late Permian and involved uplift, oroclinal bending and a rearrangement of fault-bounded blocks (Collins et al. 1993). This deformation was followed by intrusion of the voluminous, post-tectonic I- and A-type granite suites of the New England Batholith, and extrusion of their volcanic equivalents.

1A4 Previous studies of the New England Batholith

The New England Batholith (NEB) is defined here as including all granitoid suites that are late Carboniferous or younger, which are intrusive into rocks of the SNEFB. This definition excludes the Cambrian plagiogranites which occur as exotic blocks within serpentinites of the PMFS (Aitchison 1992). While some authors have previously excluded the Hillgrove Plutonic Suite from this definition (e.g. Korsch 1977) on the basis of its aerial distribution, the currently recognized extent of the suite and its spatial association to younger granitoid suites does not warrant its exclusion from the definition.

Although granitic rocks of the NEB were first described early this century (Andrews & Mingaye 1907), more intensive petrological investigations were not initiated until the 1960's, when a succession of doctoral studies concentrated on different portions of the batholith (e.g. Chappell 1966, Flood 1971, Neilson 1970 and Shaw 1964). These studies were summarized by Wilkinson (1969), and were the basis of the original subdivision into suites by Binns et al. (1967). These studies were published during the 1970's in a succession of papers dealing 7 with the petrological aspects of the batholith (e.g. Chappell 1978, Flood & Shaw 1975, 1977, 1979, and O’Neil et al. 1977). The definitive subdivision into five major suites, along with an unnamed group of leucoadamellites by Shaw & Flood (1981) presented broad petrogenetic models for the various granitoid types, and still stands as the most extensive summary of data collected on the batholith. The doctoral studies of Hensel (1982) were published in 1985 (Hensel et al. 1985) and presented the first Nd isotopic data for the batholith. While this work essentially supported earlier studies, a redefinition of suites was presented and alternative petrogenetic models were proposed for some suites. Since the early 1980's few studies have been published, excepting studies of the Barrington Tops Granodiorite (Eggins & Hensen 1987) and the Clarence River Plutonic Suite (Bryant & Arculus 1993). A summary of the named suites of the batholith along with broad petrogenetic models is presented in table 1A1.

Table 1A1 Simplified summary of the New England Batholith.

Suite Inferred Genetic Inferred source rock type Age type Shaw & Flood Hensel et al. Shaw & Flood Hensel et al. 1981 1985 1981 1985 Hillgrove Hillgrove Permo- S-type meta- meta- Carboniferous sedimentary sedimentary Bundarra Bundarra Permo- S-type meta- meta- Carboniferous sedimentary sedimentary Clarence River Nundle Permo-Triassic I-type meta-igneous mixed

Moonbi Permo-Triassic I-type meta-igneous meta- New England sedimentary Supersuite Uralla Permo-Triassic transitional mixed metasedimentary Leucoadamellites Post-orogenic Triassic A-type meta-igneous meta- (‘Stanthorpe suite sedimentary Suite’ Korsch 1977)

The foliated rocks of the Hillgrove Plutonic Suite (Binns et al. 1967) were first recognized as distinctive by Andrews (1907) due to their ‘gneissic’ nature. Early workers assigned a Silurian age to the suite (Raggat 1938, Voisey 1942) on the basis that intrusion of the granites (then called the ‘eastern belt’) and the apparently synchronous deformation had not affected 8 rocks of ‘post-Silurian’ age (in reference to the Nambucca Block sequences which are now known to be Permian in age).

Early isotopic dating of the suite (K-Ar biotite, Cooper et al. 1963, Binns & Richards 1965) revealed middle to late Permian ages, while later whole-rock Rb-Sr dating showed somewhat older ages (e.g. Flood & Shaw 1977). In short, the age of the Hillgrove Suite has proved to be somewhat enigmatic. Despite the foliated nature of granitoids belonging to the Hillgrove Plutonic Suite, few investigations have been conducted on the structural relationships of these granites to the enclosing accretion complex rocks and their bearing on regional deformation events. Dirks et al. (1992, 1993) detailed the structural relationships of the Tia Granodiorite to deformation events within the Tia Complex. Likewise, Farrell (1992) detailed the structural relationships of the Abroi Granodiorite and the Wongwibinda Complex. These studies identified several phases of pre-, syn- and post-emplacement deformation. Larger scale investigations linking the events within these two complexes to each other and to other parts of the accretion complex are lacking.

1A5 Aims of the project

The numerous geological problems associated with the Hillgrove Plutonic Suite, the oldest suite of the New England Batholith, inspired most of the aims of this project. The lack of a regional tectonic correlation of deformation events affecting the plutons of the Hillgrove Suite and their enclosing metasediments; poor constraints on the ages of intrusion and deformation of the suite; and a rudimentary knowledge of the source rocks and processes involved in generation of the magmas, were all problems that needed addressing. Additionally, an initial investigation into the A-type granitoids of the New England Batholith (the youngest group of intrusives in the batholith) which identified penecontemporaneous I- and A-type suites in the Chaelundi Complex (Landenberger 1988), required a more detailed petrogenetic investigation. Also, the unusual occurrence of mafic enclaves in some members of this A-type group, also promised to provide unique insights into both the petrogenesis of the A-type suites and the formation of mafic enclaves in granitoids generally.

In detail, the objectives were to: 9

(i) Conduct broad scale mapping of structures present in the Hillgrove Suite granitoids and the host accretion complex metasediments, thereby providing a correlation of deformation events between the Tia and Wongwibinda metamorphic complexes and areas outside these complexes. This also entailed the delineation between accretion complex metasediments and the Permian sedimentary sequences of the Nambucca Block to the east. (See chapter two).

(ii) Investigate the mylonite zones that truncate many plutons of the Hillgrove suite, specifically to determine the sense of movement on these shear zones, and to determine their relationship to regional structures. (See chapter two).

(iii) Constrain the age of intrusion of the Hillgrove Suite granitoids, and the deformation of this suite and the enclosing accretion complex rocks, using U-Pb zircon and Rb-Sr biotite dating techniques in conjunction with the structural relationships. (See chapter three).

(iv) To carry out a petrogenetic study of the Hillgrove Suite granitoids and potential source rocks in order to refine the models of magma generation for this suite. (See chapter four).

(v) To carry out a detailed petrogenetic study on the A-type granite suite of the Chaelundi Complex and its differences to the enclosing I-type granite suite, as initially identified by Landenberger (1988). (See chapter five).

(vi) To conduct a petrogenetic study of unusual occurrences of mafic enclaves in A-type granites of the New England Batholith - specifically their bearing on the formation of mafic enclaves and their contribution to the petrogenesis of the host granitoids. (See chapter six). 10

CHAPTER 2. STRUCTURAL RELATIONS OF THE HILLGROVE PLUTONIC SUITE

2A1 Introduction

The Hillgrove Plutonic Suite (Binns et al. 1967) is composed of a series of small to medium- sized, peraluminous, biotite-granodiorite and adamellite plutons which are hosted by the accretion complex sequences of the Tablelands Complex (Fig. 2A1). Plutons of the suite are generally small (not exceeding 200 km2, generally <100 km2), and comprise only ~11% of the New England Batholith (Shaw & Flood 1981).

Host metasediments intruded by the suite vary in grade from upper amphibolite facies in the Tia and Wongwibinda metamorphic complexes (Dirks et al. 1992, Farrell 1992), to prehnite- pumpellyite facies or lower elsewhere in the Tablelands Complex. Deformation in these sequences is also highly variable. Weakly deformed areas commonly display only one apparent deformation (although bedding and fold plunges are ubiquitously steep). More highly deformed zones such as the Tia and Wongwibinda complexes record several phases of penetrative deformation, with features such as primary bedding and early cleavage commonly destroyed (Dirks et al. 1992).

Granitoids of the Hillgrove Suite are variably deformed, and range from unfoliated varieties (which record mild deformation in their microfabrics - e.g. undulose quartz) to those which are strongly mylonitized. While most higher level plutons of the suite record only one apparent deformation, plutons within the deeper level metamorphic complexes record two deformation episodes (Dirks et al. 1992, Farrell 1992).

This chapter details the structures developed in areas outside the deeper level metamorphic complexes, and provides a broad scale correlation of deformation events that have affected both the Hillgrove Suite granitoids and the host accretion complex sequences. The structural study concentrates on deformation events that have directly affected the Hillgrove Suite, in particular, the faulting which caused large scale uplift of the high grade metamorphic complexes, and produced mylonitic fabrics that bound the plutons of the suite. Before a broad structural framework can be established, the structural framework within the Tia 11

1090 00 10 20 30 40 50 10

3 Plutons of the 0 20 40 00 00 Hillgrove Suite km & mafic complexes 90 90

N 60 70 80 90 80 4 80

Glen 70 70 27 26 5 Bluff CHB Major Wongw 60 60 roads 28 25

Guyra ibinda - Fault 2 50 50 WR WC SZ Dorrigo 1 40 AB 40 7 29 24 6 Ebor 30 30 Bellingen

Armidale 8 60 70 80 90 20 20 9 30 Key to structural blocks TC Tia Complex 10 er Fault NB 10 10 WC Wongwibinda Complex Uralla Chandl WSZ Winterbourne Sub-Zone Borah WRSZ Walcha-Rockvale Sub- 00 11 00 31 Fault Zone AB Armidale Block

32 12 33 Yarrowitch CHB Coffs Harbour Block 90 90 NB Nambucca Block 13 14 YB Yarrowitch Block 80 WR WSZ 80 Major late Permian SZ 15 fault zones Walcha 70 34 70 YB Western limit of Tia Fault 16 Wollomombi Zone 19 60 35 60 (Mihi Line) 20 17 Younger rock units have been 18 excluded for simplicity 50 Kilburnie Fa 50 21

Nowendoc Fault Fault Permian sediments 40 40 Permian TC Early Permian Volcanics ult

Dingo Fault 30 30 22 Gabbros & diorites Late

System Carboniferous 20 20 Hillgrove Suite

10 23 Nowendoc Accretionary metasediments 10 Pre- late MB Carboniferous Serpentinites 00 00 40 50 60 70 80 90 00 Figure 2. 1. Areal distribution of the Hillgrove Suite and associated mafic complexes. Named plutons and complexes are listed below. Hillgrove Suite Mafic complexes 1 Glenifer Adamellite 13 Eastlake Adamellite 24 Dorrigo Mountain Complex 2 Dundurrabin Granodiorite 14 Winterbourne Granodiorite 25 Charon Creek Diorite 3 Henry River Adamellite 15 Argyll Granodiorite 26 Sheep Station Creek Complex 4 Kookabookra Adamellite 16 Kimberley Park Adamellite 27 Mornington Diorite 5 Tobermory Adamellite 17 Garibaldi Adamellite 28 Day's Creek Gabbro 6 Abroi Granodiorite 18 Tia Granodiorite 29 Camperdown Complex 7 Rockvale Adamellite 19 Kilburnie Adamellite 30 Baker's Creek Diorite 8 Hillgrove Adamellite 20 Campfire Adamellite 31 Barney House Gabbro 9 Gara Adamellite 21 Ingleba Adamellite 32 Woodburn Diorite 10 Gostwyck Adamellite 22 Rockisle Granodiorite 33 Cheyenne Complex 11 Blue Knobby Adamellite 23 Murder Dog Adamellite 34 Moona Plains Complex 12 Enmore Adamellite 35 Apsley River Complex 12

Complex will be discussed, since the greatest number of deformation episodes are preserved within this complex.

2A2 Structural Setting

2A2A1 Distribution of the Hillgrove Suite

The Hillgrove Plutonic Suite is defined here as a group of 23 plutons (see Fig. 2A1), which expands the definition of Hensel et al. (1985) to include the Gara, Gostwyck, and Murder Dog Adamellites (which were previously included in a ‘transitional’ group) and the Rockisle Granodiorite (previously included in the Nundle Suite). The reasons for this expansion of the definition are discussed in chapter four.

Granitoids of the Hillgrove Suite crop out over an area spanning from the Dorrigo - Bellingen region in the northeast, to the Niangala region situated to the west of the Tia Complex (Fig. 2A1), a distance of some 200 km. Although this range has been described as a ‘linear’ belt by many workers (e.g. Shaw & Flood 1981), the megascopic outcrop pattern is distinctly curvilinear, largely following the trends of folded fabrics in the host metasedimentary sequences. This large scale ‘folding’ of the Hillgrove Suite is one of two important factors which have determined the megascopic outcrop pattern of the suite, which will be detailed below (section 2A3).

The second factor affecting the megascopic outcrop pattern of the suite, is the truncation of many plutons by late, major faults (e.g. Wongwibinda-Yarrowitch Fault System, Tia and Chandler Faults - see Fig. 2A1). These mylonite zones were responsible for uplift of the high-grade crustal blocks of the Tia and Wongwibinda metamorphic complexes (Dirks et al. 1992, Farrell 1992) and will be discussed in detail below.

Exposed plutons of the Hillgrove Suite were emplaced over a range of crustal depths. Plutons emplaced at shallow depths have narrow contact aureoles (biotite-cordierite, less than 1 km) in otherwise low-grade (prehnite-pumpellyite facies) metasediments (e.g. Gostwyck Adamellite). These plutons are generally composed of adamellite, display no foliation or other mesoscopic deformation features in outcrop and are equivalent to the epizonal granites 13 of Buddington (1959). In contrast, deeper level intrusives (e.g. Tia Granodiorite), commonly of granodioritic composition, are enveloped by regional-scale metamorphic aureoles, and commonly display a weak to strong foliation defined principally by sub-parallel alignment of relict igneous biotite, and recrystallized aggregates of biotite and quartz. This foliation transgresses pluton boundaries and is equivalent to the regional S5 foliation in the surrounding schists and migmatites. Field relationships suggest that granitoid intrusion occurred late syn-D5 (Dirks et al. 1992, Dirks et al. 1993, Farrell 1992). Development of migmatites occurs adjacent to granite intrusions in the Wongwibinda and Tia metamorphic complexes. Peak temperatures of ~700EC (P=2A5-3A0 kilobars) in the Wongwibinda Complex (Farrell 1992) and ~650EC (P=2A5-3A0 kilobars) in the Tia Complex (Dirks et al. 1992) were attained, implying local geothermal gradients of ~70EC/km.

2A2A2 Structural history of the Tia Complex

The Tia Complex is the most complete and best studied area illustrating the protracted structural history of the accretion complex of the SNEFB. Accordingly, all recognized structural events elsewhere will be correlated (section 2A3) with the structural sequence recorded in this complex. Structural studies of the Tia Complex (Hand 1988, Dirks et al. 1992) reveal eight phases of deformation, three of which are penetrative. The following discussion of structural evolution within the Tia Complex is largely based on Dirks et al. (1992), combined with personal field observations.

Early subduction-related deformation events D1 and D2 (D2 is the first penetrative deformation) are preserved only in the southern portion of the complex and record blueschist facies metamorphism. These were followed by SW-over-NE thrusting and uplift during D3

(dated at ~311 Ma by K-Ar white mica, Offler et al. unpublished data). The effects of D3 are widespread, with S3 being the dominant form-surface throughout much of the central and northern parts of the complex. D3 was accompanied by the onset of a thermal overprint of the early blueschist facies fabrics by high-T/low-P metamorphism, which continued during progressive uplift D4 - D5. The effects of D4 are restricted to the central and southeastern half of the complex, and are associated with open to tight folds with wavelengths up to 6 km.

Cleavage formation during D4 is restricted to a crenulation of the S3 form-surface. 14

D5 is the last recorded penetrative deformation in the Tia Complex, and its effects are most intense along the eastern boundary of the complex, forming a 5 km wide shear zone. In this shear zone, S5 has formed an intense mylonitic fabric which has obliterated all previous structures, and records an W-over-E shear sense. Dirks (1992) demonstrated with field evidence that the Tia Granodiorite (a member of the Hillgrove Suite) intruded late in this phase of deformation. This conclusion is supported by similar evidence from mapping in the Wongwibinda Complex, where Farrell (1992) concluded that the Abroi Granodiorite intruded late- syn-D2 (the equivalent deformation phase in at Wongwibinda - see below). S5 within the Tia Granodiorite is variable but generally weak, defined by an alignment of biotite and

5 quartz stringers, and is occasionally accompanied by a weak L5 stretching lineation.

Deformation continued after granite intrusion with D6 causing large scale, open folding of all earlier fabrics. The effects of D6 are unrecognized on outcrop scale, and its effects are only apparent megascopically, as F6 folds have wavelengths of >20 km (see below). The effects of D7 are most easily observed within the Tia Granodiorite, where a steep, E-dipping foliation has truncated the S5 fabric. S7 forms a spaced cleavage (~1 cm) often defined by growth of new fine-grained biotite and secondary muscovite, which together with quartz

7 rods, define a down-dip stretching lineation L7.

D8 in the last recognized phase of deformation in the Tia Complex, and consists of numerous narrow, subvertical ultramylonite and pseudotachylite veins, which are most common in the eastern half of the complex, and were interpreted by Dirks et al. (1992) as a brittle reactivation of the D5 shear zone.

2A3 Structural Framework

The Hillgrove Suite crops out within a number of structural domains or ‘zones’ with distinctive structural styles (as outlined in Fig. 2A1). This discussion of the structural framework will first define each of these domains (section 2A3A1). The macroscopic structural features within each domain, and structural continuity between domains are then examined (section 2A3A2). As noted earlier, the most complete structural sequence in the SNEFB is preserved in the Tia Complex, and hence the deformation history recorded there, provides the basis around which all structural events recognized elsewhere should be 15 correlated. Hence, an initial correlation of structural events is presented at this stage on the basis of these macroscopic features. This pre-emptive correlation is presented here to avoid confusion of the reader, and to facilitate discussion of the meso- and microscopic deformation features, which are described for each structural event, rather than for each domain (section 2A3A3). Following this, a more rigorous and detailed correlation of structural events between each domain, and other areas of the SNEFB is presented (section 2A3A4).

2A3A1 A definition of domains

Although some plutons of the Hillgrove Suite have intruded the Coffs Harbour and Armidale Blocks, most occur within an area called the Wollomombi Zone, which was initially defined by Korsch & Harrington (1981), and includes the Tia and Wongwibinda complexes. The Wollomombi Zone is defined here, as the area bounded to the east by the Wongwibinda - Yarrowitch Fault System (WYFS) and to the west by the thick dashed line shown in figure 2A1. This line marks the approximate position of the ‘Mihi Fault’ (Pogson & Hitchins 1972), which is a delineation of rocks to the west which display no obvious post-accretion deformation, from rocks to the east which are structurally dominated by overprinting deformations. Although there is no direct evidence that a major fault exists here, the boundary shown as a dashed line (Fig. 2A1) serves to mark the spatial onset of deformation fabrics which are not accretion/subduction related, and is referred to herein as the Mihi Line. This definition of Wollomombi Zone specifically excludes rocks of the Yarrowitch Block (Fig. 2A1), which occurs to the east of the WYFS (previously termed the ‘East Yarrowitch Block’, Leitch 1988).

For the purposes of the following structural outline, the Wollomombi Zone has been subdivided into four separate zones of contrasting structural styles:

1. The Tia Complex (TC) is bounded to the east and northeast by the Tia Fault, and to the south and west by the Nowendoc Fault (Dirks et al. 1992).

2. The Wongwibinda Complex (WC) is bounded to the east by the WYFS, but the western boundary is more obscure, being marked by a westward drop in metamorphic grade (Farrell 1992), just west of the marked extension of the Glen Bluff Fault (Fig. 16

2A1).

3. The Winterbourne Subzone (WSZ - new definition) is bounded to the east by the WYFS and to the west by the Tia and Chandler Faults (which have been joined by a dashed line (Fig. 2A1), representing the western margin of the zone which is probably unfaulted in this area).

4. The Walcha - Rockvale Subzone (WRSZ - new definition) is bounded to the east by the Chandler and Tia faults and to the west by the ‘Mihi Line’. This zone includes the Rockvale Block as defined by Korsch (1981), which occurs to the west of the Wongwibinda Complex, and is bounded to the north by the Tobermory Adamellite. The southern margin of the zone is defined by the Nowendoc Fault and units of the small Permian outlier which occurs to the east of the Kilburnie Fault, which are informally termed the Spitzbergen Creek Beds herein.

2A3A2 Structural continuity outside the Tia Complex - Macroscopic structural features

The Wongwibinda Complex (WC) The early deformation sequence of the Tia Complex is not recorded in the Wongwibinda

Complex, and blueschists are absent. However, the ubiquitous steep plunge of both F1 and

F2 folds in the Wongwibinda Complex suggests a vertical orientation of bedding prior to D1

(termed D0 by Farrell 1992), which is consistent with this sequence forming part of a subduction-accretion complex. D1 and D2 in the Wongwibinda Complex both form upright folds, with the initial E-W trending F1 axial planes reoriented by D2, which produced N-S folds and a strong axial plane cleavage (S2). Farrell (1992) demonstrated that the Abroi

Granodiorite (a member of the Hillgrove Suite) intruded the complex late during the D2 event. For this reason, together with the similarity of fabric orientation (indicating E-W compression), D2 is equated with D5 in the Tia Complex. D1 at Wongwibinda has been equated with D3 in the Tia Complex (Farrell 1992), also due to the similarity of structural style and orientation of fold axes, although the intense shearing during D3 in the Tia Complex is not evident at Wongwibinda. 17

The Winterbourne Subzone (WSZ) Eastward from the Tia Complex, the ~25 km wide WSZ records variable, but relatively intense deformation. The Tia Fault (D5) marks the western boundary of this zone, and also bounds the eastern margin of the Tia Complex (Dirks et al. 1992). D5 within the Tia complex has folded an S3 form-surface, and on the eastern margin, S3 is completely transposed into S5, with intense W-over-E shearing (Dirks et al. 1992). The boundary with the WSZ is marked by the sudden appearance of bedding, along with a sudden drop in metamorphic grade on the eastern side. Dirks et al. (1992) suggested that a D7 shear exists in this location. Although S0 is the form surface to the east in the WSZ, the same intense fabric continues across the zone, parallel to S5 in the Tia Complex, exhibiting a northwesterly trend (Fig. 2A3). This same cleavage can be traced into the higher grade rocks of the Moona Plains Complex, where it is continuous with the foliation in the Argyll Granodiorite (Fig. 2A3). Further north, in the region north of the Enmore Adamellite, the same fabric has a northeasterly orientation, and the western boundary of the WSZ is more diffuse. Cleavage and folding gradually lose intensity westward in this area, and cleavage is no longer observed west of the Mihi Line. North of the Enmore region, the Chandler Fault

(a D7 shear) forms the western boundary of the WSZ. However, a weak cleavage, which may be of the same generation, occurs in the Metz area just north of the north of the Gara Adamellite near Metz (within the Walcha-Rockvale subzone - Fig. 2A3). This deformation is here equated with D5 of the Tia Complex, and more evidence for this correlation will be discussed below (section 2A3A4).

The Walcha-Rockvale Subzone (WRSZ) Compared to other subdivisions of the Wollomombi Zone, rocks of the Walcha-Rockvale Subzone are not intensely deformed. Cleavage is often weak or absent, and folding is generally large-scale and of an open character. Included in this zone are the mildly deformed sequences of the Rockvale Block, described by Korsch (1981). Within the Rockvale area, two phases of deformation have been recorded, in addition to a vertical attitude of bedding prior to deformation (Korsch 1981). This compares well with the structural sequence in the Wongwibinda Complex. The dominant cleavage within this northern portion of the WRSZ (termed S1 by Korsch 1981) is oriented east-west, and can be traced eastwards into S1 of the Wongwibinda Complex, and thus is equated here with D3 of the Tia Complex. D2 in the Rockvale area has caused folding about a north-south axis, but 18 has produced little or no cleavage (Korsch 1981). Since this generation of folds can traced eastwards into F2 folds of the Wongwibinda Complex, then the two are equated, and therefore are the equivalent of D5 in the Tia Complex.

South of Rockvale, in the area surrounding the Hillgrove Adamellite and Gara Adamellite, cleavage is sporadic and often absent. Where developed, it parallels southwest-northeast trending folds, and it occasionally forms a weak mylonitic fabric. To the east of the Gostwyck Adamellite, folds of the same generation are generally open, large-scale and steeply plunging. However, in an easterly traverse of Salisbury Waters (from GR711080), there is a general tightening of folds, which is accompanied by the development of an axial plane cleavage in the vicinity of Fairfield (GR739062). Further east, this cleavage intensifies and can be traced into the more intense fabrics of the Winterbourne Subzone. Hence, the sporadically developed cleavage in this central section of the Walcha-Rockvale

Subzone are equated with S5 in the Tia Complex.

Interbedded cherts and metapelites define the form surface in the southern portion of the WRSZ, in the vicinity of Walcha. The sequence here is folded into open to tight, steeply plunging mesoscopic folds, with axial plane cleavage well developed in pelitic units. This cleavage parallels S3 within the Tia Complex (Fig. 2A2). The dominant fabric in the nearby northwestern margin of the Tia Complex is S3, and has a northeasterly trend. Approximately 1 km to the south of Walcha, this fabric is abruptly terminated by an extension of the Nowendoc Fault. Although the fault does not outcrop, the sudden appearance of bedding and dramatic drop in metamorphic grade implies its presence. Although detailed mapping has not been conducted in this area, this fabric is tentatively equated with D3 in the Tia Complex.

The Coffs Harbour Block (CHB) Deformation within the Coffs Harbour Block has been described by both Korsch (1981) and Fergusson (1982). Fergusson (1982) detailed the structure in the central and northern parts of this block, and attributed the main phase of deformation to subduction/accretion processes

(/D1,D2 of the Tia Complex). Typical structures indicative of subduction within the CHB are melange zones (e.g. Gundahl Melange) and in less disturbed zones, large scale, shallowly plunging upright folds with asymmetries indicating a southwesterly vergence. 19

These structures were termed D1 by both Korsch (1981) and Fergusson (1982). They have a general northwesterly trend, and are gently warped across the block, trending more east- west near the coast.

Post-accretionary deformation structures within the Coffs Harbour Block are generally less penetrative than D1 structures. Fergusson (1982) described F2 as open to close kinks in S0 and S1, with steeply inclined axes and near vertical plunges. F2 structures are generally better developed in the southern portion of the Coffs Harbour Block, where gentle flexures, kink bands and chevron folds are relatively common (Korsch 1981). The Dundurrabin Granodiorite, which crops out in the southern part of the Coffs Harbour Block has a foliation defined by aligned biotite flakes and quartz aggregates. This foliation is best developed near the southern margin of the intrusion, and is parallel to the long axis of the pluton and to fabrics of the host accretion complex sequences (Fig. 2A3). In addition, metamorphic biotite within the contact aureole of the Dundurrabin Granodiorite is variably aligned in a weakly developed cleavage which generally parallels the pluton margin. Although the dominant fabrics within the southern Coffs Harbour Block are associated with D1 (Korsch 1981), the occurrence of this secondary cleavage, together with the foliation developed within the

Dundurrabin pluton, indicate that the effects of S5 (/D2 of Korsch 1981) are weakly developed in this block (as depicted in Fig. 2A3). However, more detailed structural mapping is required to elucidate the extent of S5 fabrics in the southern part of the Coffs Harbour Block.

2A3A3 Meso- and Microscopic Deformation Features

The deformation fabrics which are most widespread and correlatable throughout the

Wollomombi Zone (including the Hillgrove Suite granitoids) are attributed to the D5 and D7 events of the Tia Complex. The meso- and microscopic characteristics of the earlier penetrative phases of deformation (D2 and D3), which are essentially restricted to the Tia

Complex and Wongwibinda Complex (D3), have been described in detail by Dirks et al. (1992) and Farrell (1992), and so will not be discussed in detail here. The following discussion of meso-and microscopic features of D5 and D7 within the Wollomombi Zone, firstly describes the characteristics of these fabrics as developed in accretion complex sequences, followed by those developed in the Hillgrove Suite granitoids. 20

Plate 2A1. D5 macro- and mesoscopic features of metasediments from the Wollomombi Zone.

(a) View looking south in Salisbury Waters gorge. The structure in the cliff face is a

macroscopic D5 fold in a greywacke - siliceous lutite sequence. The fold has a upright axial plane oriented ~130E and plunges steeply westwards. The younging direction of the beds is also westward. The trace of the axial plane is marked with a dashed line. The Casuarinas in the river bed are ~5 metres tall.

(b) Well-developed pencil cleavage in pelites of the Wollomombi Zone. S5 is vertical in the picture. The secondary spaced cleavage indicated by the pencil may relate to

a nearby D7 mylonite zone in the Enmore Adamellite (Borah Fault). The pencil is ~15 cm in length.

(c) Tight, steeply plunging mesoscopic F5 folds in a siliceous lutite - pelite repeating sequence in the bed of the Chandler River, Wollomombi Zone.

(d) Isoclinal, steeply plunging mesoscopic F5 folds in a siliceous lutite - greywacke sequence, Wollomombi Zone. Note the truncation of fold limbs by minor shears. (b) (a) Up West West

p lunge

West

10 cm axial

plane she ar plane 21

D5 fabrics

The effects of D5 within the Tia Complex are variable, with S5 only becoming the dominant fabric near the eastern margin of the complex (Dirks et al. 1992, - see discussion above).

East of the Tia Complex, within the Winterbourne Subzone, S5 is the dominant fabric throughout, with the observed variations in intensity merely reflecting the competency of host rock-types.

Psammitic sequences within the Winterbourne Subzone generally display the least deformation. Where bedding is not observed, greywackes are massive, and only rarely develop cleavage. However, these ‘massive’ greywackes usually display microstructural evidence of deformation, mainly defined by strong undulose extinction and deformation bands in quartz phenoclasts (see plate 2A2b). The consistency of orientation of these deformation bands in separate quartz clasts within a given sample, indicates that this is a post-depositional, rather than an inherited feature. Where cleavage does develop within greywackes, it is discontinuous, widely spaced, and is usually defined by muscovite/chlorite or muscovite/biotite growth in the finer-grained matrix of greywackes (see plate 2A2c).

0 F5 folds are best observed in interbedded greywacke - siliceous lutite (argillite) sequences. These interbedded sequences are prevalent throughout much of the Winterbourne Subzone, and are usually accompanied by minor pelitic units and intraformational conglomerates. Bedding ranges from centimetre to metre scale, and both meso- and macroscopic folds are readily observed in the field. Cleavage is usually weak in greywacke units, stronger, but still discontinuous in argillaceous units, and is penetrative in pelitic layers, where it is axial planar

0 to F5 folds. Cleavage refraction is often observed between pelitic and argillaceous horizons. 0 F5 folds are usually tight to isoclinal and similar in style on outcrop scale (see plates 2A1c-d), and are often crosscut by small-scale fine shear bands (plate 2A1d D6?). Larger scale F5 fold hinges are often outlined by more massive greywacke units (plate 2A1a), and display a more

0 cylindrical style. F5 fold plunges are consistently steep (always >70E, see map 1 - rear 0 0 5 pocket), with both measured F5 fold plunges and L5 intersection lineations near parallel to L5 0 5 stretching lineations within adjacent pelitic units. F5 fold plunges and L5 stretching lineations 0 from across the WSZ are plotted stereographically in figure 2A5. The steepness of F5 fold plunges on all scales indicates that bedding was subvertical prior to D5, which is consistent with observations elsewhere in the Wollomombi Zone, and supports the conclusion that this 22

Plate 2A2. D5 microstructures in metasediments.

(a) Mylonitized psammo-pelite (sample A79) with tightly folded pelite bands and lithic clasts with porphyroclast trails (σ) deformed in an intense mylonitic fabric. Crossed polars - base of photograph = 17 mm.

(b) Mildly deformed greywacke (sample A4) showing quartz with deformation bands. Crossed polars - base of photograph = 5 mm.

(c) Deformed greywacke (sample A18) with a large volcanic clast (bottom left) and strongly deformed matrix with syn-deformation biotite. Crossed polars - base of photograph = 5 mm.

(d) Mylonitized pelite (sample Y45) showing syn-tectonic biotite growth. Plane polarized light - base of photograph = 625 μm. 03

(a) (b) 14

13 23 folding is postdates accretion within the subduction/accretion complex.

The non-coaxial nature of the strain developed during D5 is recognized on outcrop scale by 5 the presence of a stretching lineation L5, which is most obvious in pelitic units. Although both the minerals defining this lineation and the indicated sense of shear in pelites are only revealed in their microstructure, mesoscopic evidence of the non-coaxial nature of this deformation is often spectacularly revealed in intraformational conglomerate horizons. The muddy intraclasts within these units define a stretching lineation on spaced S5 cleavage planes, with millimetre to centimetre wide clasts stretched to aspect ratios which vary up to 10:1. These intraclasts also show σ porphyroclast asymmetry, which consistently indicates a W-over-E sense of shear for D5, an observation consistent with that for pelitic units (see below). Microscopically, these sheared intraformational conglomerates reveal textures consistent with those observed in hand specimen. The example shown in plate 2A2a (sample A79), shows large stretched intraclasts (the darker zones of the photomicrograph). Although these intraclasts extend beyond the field of view, the asymmetry displayed by smaller lithic clasts in the sample, is consistent with observation on hand specimen scale (in this case indicating dextral shear). Microporphyroclasts within the intraclasts (mostly single quartz grains) also indicate the same sense of shear as their hosts do on a mesoscopic scale. Truncated folds defined by quartz - muscovite (±biotite) rich layers are also visible in this sample, and highlight the intense shearing within the intraclasts and matrix of this sample.

Pelitic sequences typically display the most penetrative effects of D5, and commonly form slates and phyllites. Spectacular examples of these slates can be observed at Apsley Falls, in a ~200 m vertical section, with the sheer gorge walls dominantly formed by the subvertical S5 form-surface. Cleavage in these rocks is a finely spaced solution enhanced cleavage (plate 2A1d), generally defined by muscovite and chlorite (or biotite and muscovite in areas where biotite is developed). Folding, where visible, is often defined by isoclinally folded thin quartzo-feldspathic layers, which are commonly truncated, leaving isolated fold hooks (intrafolial folds). S5 in pelites is always accompanied by an intense down-dip 5 stretching lineation (L5), which is generally defined by fine-grained, stretched quartz- feldspar domains and a general a colour variation on cleavage surfaces (alternating muscovite-rich and chlorite-rich layers - S0?). 24

Although not evident in hand specimen, the microstructure of these strongly cleaved pelites reveals that S5 is a mylonitic fabric. Although S-C relationships are difficult to identify microscopically (S is usually not observed due to the close spacing of C-planes), asymmetry displayed by micro-porphyroclasts may be used to infer the sense of shear. This case is evident in plate 2A2d (sample Y45 - biotite grade slate) where quartz micro-porphyroclasts (σ) show asymmetrical relationships (in this case indicating sinistral shear), consistent with the poorly defined S-C asymmetry (S-planes are occasionally preserved within the finer quartz-rich layers). This example (plate 2A2d) also demonstrates the syn-tectonic growth of biotite within the Winterbourne Subzone (where biotite grade is reached), and shows relatively coarse grained biotite aligned in C-planes, as well as finer-grained biotite which appears to grow across this fabric (syntectonic biotite growth is also evident in plate 2A2c). Placed in their original orientation, the sections indicate W-over-E sense of shear, which is consistent throughout the Winterbourne Subzone.

Within granitoids of the Hillgrove Suite, the effects of D5 are most apparent in the deeper level plutons, such as the Abroi Granodiorite in the Wongwibinda Complex, which displays a weak gneissic foliation that is parallel to the schistosity in the surrounding metasediments. In the Tia Granodiorite, this foliation is more penetrative and shows a weak down-dip

5 stretching lineation L5 (Dirks et al. 1992). In general, foliations associated with D5 are defined by lattice preferred orientation of biotite and the dimensional preferred oriented of deformed quartz grains and recrystallized aggregates of quartz and biotite ± feldspar.

S5 is well developed within the western portions of the Argyll Granodiorite (Fig. 2A3), which occurs at the eastern margin of the Winterbourne Subzone. Within the pluton, it is similar to

S5 in the Abroi Granodiorite, displaying no obvious stretching lineation. This fabric is continuous with cleavage in the surrounding metasediments of the Winterbourne Subzone

(plate 2A3a), supporting the conclusion that the dominant fabric throughout this subzone is S5. Although cleavage in the enclosing metasediments is always strong, the equivalent foliation within the Argyll Granodiorite is variable, and is not apparent in the eastern portion of the pluton. This supports the arguments posed by Dirks et al. (1992) and Farrell (1992), that

Hillgrove Suite intrusion occurred late syn-D5. Other plutons which display a weak (and sporadic) gneissic foliation which is parallel to S5 in the enclosing metasediments, include the eastern portions of the Enmore Adamellite and the Blue Knobby Adamellite, and the southern 25

Plate 2A3. Mesoscopic features of D5 D7 and D8 structures in Hillgrove Suite granitoids.

(a) Intrusive margin of the Argyle Granodiorite at Moona Plains. Partially dismembered

D5 folds show an axial plane parallel to a weak S5 foliation (parallel to the pencil) in the granodiorite.

(b) D8 pseudotachylite veins (arrowed) in the Kookabookra Adamellite adjacent to the Triassic Woodlands Quartz Monzonite. Hammer handle is 3 cm in width.

appearance of C and S D7 fabrics within Hillgrove Suite granitoids. The sense of shear is east over west.

(d) D8 pseudotachylite veins in the Henry River Adamellite adjacent to the Triassic granitoids. (a) (b)

10 cm

S (c) (d)

10 cm

C Up

West

10 cm 26 margin of the Dundurrabin Granodiorite (see Fig. 2A3).

The higher crustal level plutons of the Hillgrove Suite, although lacking evidence of deformation on a mesoscopic scale, do show evidence of subsolidus deformation on a microscopic scale, including slightly kinked and bent magmatic biotite, and quartz exhibiting undulose extinction and less commonly deformation bands (plate 2A4h). The timing of this deformation is uncertain, and may be due to either the D5 or D7 event. Development of similar deformation features within granites of the slightly younger (post-D5) Bundarra

Plutonic Suite to the west, suggests that the effects of D7 were more widespread than the restricted development of shear zones would suggest. The extensive effects of this deformation is also supported by the widespread but variable triclinicity of alkali feldspar throughout the Hillgrove Suite, and by the increase in triclinicity in alkali feldspar phenocrysts as the Kilburnie Fault (a D7 shear zone) is approached (Flood 1971).

D7 fabrics

The major manifestation of D7 is regional-scale faults such as the WYFS, which truncate plutons of the Hillgrove Suite. The effects of D7 within the host metasedimentary sequences are more difficult to recognize, since the fabrics formed are generally concordant with previously developed intense fabrics. In these cases, the D7 shears are often only recognized by sudden changes in metamorphic grade. An example of this is the Tia Fault (D7), which has formed within the D5 shear zone that bounds the eastern margin of the Tia Complex, although recognition of the fault is assisted here by the sudden appearance of S0 as described previously. However, in cases where there is no detectable change in metamorphic grade

(e.g. Borah Fault), these D7 shear zones are difficult to trace laterally beyond the boundaries of plutons.

Where accretionary prism metasediments have been juxtaposed against rocks of the Permian rift basin sequences (i.e. the Nambucca Block), the fault systems which truncate plutons are more easily traced into the host metasedimentary sequences. The distinction between these contrasting sedimentary sequences, has been used to combine the Wongwibinda Fault (the eastern margin of the Wongwibinda Complex) with the Yarrowitch Fault (the eastern boundary of the Winterbourne Subzone) into the Wongwibinda - Yarrowitch Fault System (WYFS - new name, c.f. Collins et al. 1993). This fault system defines the western 27 boundary of the Permian metasediments of the Nambucca Block (see Fig. 2A6). Although the precise locality of this fault often proves difficult to pinpoint in the field, the factors which distinguish accretion complex metasediments from the Permian basin metasedimentary sequences may be used to infer its position. These juxtaposed metasedimentary sequences share many lithotypes (e.g. felsic greywackes, lithic arenites and siliceous siltstones), but are contrasted by many other lithological and structural criteria which were used to distinguish these them in the field. These criteria are summarized in table 2A1.

Accretion complex Permian (Nambucca Block) Sequence metasediments metasediments Distinctive Intraformational conglomerates, Diamictites, orthoconglomerates, lithological mafic and intermediate tuffaceous beds, and greenish units greywackes (& lithic arenites) (chloritic) slates.

Distinctive A S0 form surfaces always have near A S0 form surfaces are usually shallowly structural vertical attitudes. (<45E) dipping. styles A All fold plunges are steep. A Fold plunges generally shallow.

A Folds (F5) tight to isoclinal. A Folds (F1) open to close.

A Penetrative cleavage surfaces (S5) A Penetrative cleavage surfaces

always accompanied by stretching (S1) have no stretching lineation (and

5 lineation (L5). are often crenulated by D2).

Table 2A1. Distinguishing characteristics of accretion complex and Permian basin metasediments.

Where the WYFS has been precisely located within metasedimentary sequences, it is always narrow (<50 metres) relative to the equivalent D7 shear within granitoids. The specific characteristics of the shear zone which are developed, are dependent on the parental lithology involved. At Jeogla Warm Corner (junction of the Oaky and Chandler rivers GR 055067) a tightly folded greywacke - siliceous lutite sequence (F5 - see plate 2A1c) of the

Winterbourne Subzone is transposed from the general trend of S5 in this region (060E - 080E) into an intense near-vertical fabric with an orientation of ~010E, which is accompanied by a near-vertical stretching lineation (pitching 75E south). This zone persists for some 100 metres, in which the precursor rock types appear to have been effectively homogenized, with structures such as bedding no longer apparent. East of this zone, Permian metasediments are readily recognized by the presence of a shallowly dipping sequence of interbedded 28 diamictites and pelites. Further south, the higher grade rocks of the Moona Plains Complex, are juxtaposed against Permian diamictites in the vicinity of Budd’s Mare Creek (GR 000720). There, a high strain mylonite is developed over a narrow zone (<50 m). The shear (C) plane orientation of this zone (150E) intersects cleavage orientations of both the Permian diamictites (~050-060E), and the higher grade rocks of the Moona Plains Complex (~020E) at a high angle. A intense down-dip stretching lineation is developed within this zone (656260), and west-over-east thrusting is indicated by tails on quartz porphyroclasts that originated as chert and jasper pebbles within the Permian diamictites.

Granitoids of the Hillgrove Suite directly affected by D7 shearing display penetrative Type 1 S-C mylonitic (Lister & Snoke 1984) fabrics developed across distinct shear zones that vary in width from 300 metres (e.g. Kilburnie Fault) to 2 km (e.g. Wongwibinda Fault). Faulting has juxtaposed the plutons against low grade metasediments. An example from the Dingo Fault (plate 2A3c, Rockisle Granodiorite - west-over-east, reverse thrusting) shows the typical fabrics developed within these zones. Although the angle between S- and C-planes in this example is ~20E, this angle varies between and within individual shear zones, depending on the local accumulated strain. All D7 shear zones show a strong down-dip mineral elongation lineation defined by elongate biotite and quartz aggregates. S-C relationships, porphyroclast asymmetry (mainly σ) and biotite-fish in mylonites indicate a west-over-east, reverse sense of movement on all shear zones, with the exception of the Kilburnie Fault which dips steeply east and shows east-over-west thrusting.

As a consequence of the range in depth of emplacement of the Hillgrove Suite granitoids, the

D7 microstructures developed within granitoids vary considerably. The fabrics reflect both the amount of accumulated strain and the metamorphic conditions under which shearing occurred. S-C mylonites which developed under lower amphibolite facies conditions at deeper crustal levels (e.g. WYFS at Wongwibinda), contain annealed quartz ribbons and recrystallized biotite which define the C-planes, while both microcline and plagioclase exhibit plastic deformation and grain boundary recrystallization (plate 2A4a-b). Under these metamorphic conditions, igneous plagioclase developed porphyroclast trails of recrystallized oligoclase, with internal fractures also healed by oligoclase, and microcline developed trails of recrystallized microcline aggregates. Partial replacement of microcline by myrmekite has also occurred. Relict igneous biotite was replaced by both the mylonitic biotite (of the same 29

Plate 2A4(a-d). Microscopic features of D7 in Hillgrove Suite granitoids.

(a) Sample W445 - Amphibolite facies mylonite from the WYFS (Abroi Granodiorite) showing C-planes defined by recrystallized quartz and biotite, and plagioclase with porphyroclast tails (σ) composed of oligoclase. Base of plate = 7 mm.

(b) Sample W444 - Amphibolite facies mylonite from the WYFS (Abroi Granodiorite) showing microcline porphyroclasts with healed cracks and recrystallized margins composed of oligoclase. Base of plate = 7 mm.

(c) Sample W445 - Magmatic biotite with marginal recrystallization to muscovite + quartz (top) and mylonitic biotite. Base of plate = 3A5 mm.

(d) Sample A65 - Greenschist facies mylonite from the Chandler Fault (Hillgrove Adamellite) showing brittle behaviour of plagioclase grains with cracks healed by quartz, and C-planes principally defined by partly recrystallized quartz ribbons and green biotite. Base of plate = 7 mm. (a) (b)

Plate 2A4(e-h). Microscopic features of D7 in Hillgrove Suite granitoids.

(e) As (d) showing strongly deformed magmatic biotite with marginal recrystallization into aggregates of green biotite + ilmenite, and brittle fracturing of plagioclase. Base of photo 7mm.

(f) Sample N1 - greenschist facies mylonite from the Kilburnie Fault (Kilburnie Adamellite) showing brittle fracturing of a microcline porphyroclast, with fractures healed by quartz. Base of plate = 3A5 mm.

(g) Sample N3 - greenschist facies mylonite from the Dingo Fault (Rockisle Granodiorite) showing brittle fracturing of feldspars and C-planes defined by poorly recrystallized quartz ribbons. Base of plate = 7 mm.

(h) Sample A180 (Gara Adamellite) showing a relatively undeformed igneous texture with minor strain shadows in quartz. Base of plate = 7 mm. (e) (f) 14

(g) 18(h) 31 composition as its precursor) defining C-planes and by less common muscovite-ilmenite aggregates (plate 2A4c). The presence of retrogressive oligoclase in high grade shear zones at Wongwibinda and Tia suggest that temperatures were >500EC at the beginning of shearing (Dirks et al. 1993, Farrell 1992). The presence of recrystallized hornblende in C-planes from mylonites within the Mornington Diorite (Fig. 2A1) also implies a minimum of lower amphibolite facies conditions (Farrell 1992).

Mylonites developed at higher crustal levels (e.g. Kilburnie Fault, Dingo Fault), have developed microfabrics consistent with shearing at lower metamorphic grade. Highly elongate quartz ribbons and minor mylonitic biotite define the C-planes (plate 2A4g). The quartz ribbons commonly show internal anastomosing slip planes, and are characteristically not annealed (plate 2A4g) or show total to partial recrystallization along these slip planes, into polycrystalline aggregates of low-strain quartz (plate 2A4d,e). Biotite shows limited recrystallization into aggregates of green biotite (low-Ti variety) and ilmenite. Plagioclase and microcline exhibit brittle behaviour with no recrystallization. Both feldspars show fracturing and displacement along cleavage planes and porphyroclast rotation (plate 2A4f), with fractures invariably healed by quartz. These deformation textures, together with the retrograde mineral assemblage developed, indicate that greenschist facies conditions prevailed during shearing (c.f. Gapais 1989).

D8 fabrics As noted above (section 2A3A3) subvertical ultramylonite and pseudotachylite veins are widespread throughout the Hillgrove Suite, but are most numerous in proximity to D7 shear zones, and often parallel D7 fabrics. As noted by Dirks et al. (1992), they are probably closely related to D7, representing the last phases of the D7 uplift, with the change in style of deformation probably representing a transgression into the brittle/ductile transition zone as the rocks were progressively lifted to higher crustal levels (c.f. Dirks et al. 1992). Field examples of these pseudotachylite zones are shown in plate 2A3(b and d). 32

2A3A4 A correlation of post-accretion deformation structures - the regional perspective

The following is a detailed discussion and correlation of post- subduction/accretion structural events (i.e. post D1, D2) across the Wollomombi Zone and the Coffs Harbour Block. In addition to the structural maps referred to in this section (Figs. 2A2 to 2A4), a more detailed fold-out map is attached in the rear pocket, and details most field structural data, including additional data (e.g. cleavage in Permian sequences) not presented in other figures. A ‘time- space’ summary of the deformation (and magmatic) events discussed in this section is presented in table 2A2. The absolute time scale depicted in this table (far left) is largely based on data from Landenberger et al. (1995), which are presented in chapter 3.

D3 structures form the dominant fabrics throughout much of the Tia Complex. The orientation of D3 fabrics within the complex is highly variable due to folding during D4 and

D5 (Fig. 2A2). As depicted in figure 2A2, the only areas outside the Tia Complex where the effects of D3 are clearly evident, are in the Wongwibinda Complex, and parts of the Walcha-

Rockvale Subzone. Although the orientation of S3 is variable throughout the belt, removal

(i.e. unfolding) of the effects of later deformation events (D4 and D5 within the Tia Complex,

D5 in the Wongwibinda Complex, and D6 across the Wollomombi Zone), reveals an original orientation of S3 fabrics (WNW-ESE) consistent with near N-S compression (c.f. Dirks et al. 1992, Farrell 1992).

Fabrics associated with D3 are not evident within the Winterbourne Subzone, where the main fabric is a product of D5. Within this subzone, is likely that any earlier fabrics have been overprinted by S5 (see below). Evidence for fabrics associated with the effects of the D3 are also scarce in areas outside the Wollomombi Zone, where the main preserved deformation fabrics have generally been interpreted as subduction/accretion (D1, D2) related (e.g. the Coffs Harbour Block, Fergusson 1982). However, the effects of this event may be more widespread in areas of the fold belt which are poorly mapped. For example, blueschists have been reported in the southern part of the Yarrowitch Block (to the east of the WYFS, Nano 1987), however this area remains largely unmapped and may reveal a similar deformation history to that recorded in the Tia Complex. Age Magmatic events Recognized structural events

(Ma) Wollomombi Zone 3 Tectonic Cycle

Period Coffs Harbour Block *Tia Complex1 Wongwibinda Winterbourne Walcha-Rockvale I- and Complex2 Sub-zone Sub-zone 5 A-type granites Triassic 255 D8 Ultramylonites D4 Ultramylonites D4 Ultramylonites 4 D W-over-E thrusting D W-over-E thrusting D W-over-E thrusting 265 7 3 3

? ? D3 large - scale ?

D6 large - scale D2 large - scale open folding 270 open folding open folding and oroclinal ? bending

Permian ? ? ? Early Permian 280 bimodal volcanism 3 Bundarra Early Permian rifting and basin formation Suite 290

Hillgrove 300 D5 W-over-E thrusting D2 E-W compression D1 W-over-E thrusting D2 E-W compression D2 SW-NE compression Suite (sporadically 2 developed)

D4 N-S compression D SW-over-NE 310 3 thrusting D1 N-S compression D1 N-S compression

Carboniferous Termination of D -D 1 2 D - pre-D accretion D - pre-D accretion D - pre-D accretion D accretion 1 Carboniferous arc subduction/accretion 0 1 0 1 0 1 1

Table 2. 2. Comparative timing of structural and magmatic events throughout the Tablelands Complex. References are: 1 Dirks et al. 1992, 2 Farrell 1992, 3 Fergusson 1982. * Nomenclature for structural events in the Tia Complex are adopted for the discussion of all areas. Dashes indicate that deformation equivalent to that recorded in the Tia Complex is not apparent. The calibrated time scale (far left) is largely based on age data presented in chapter 3. Tectonic Cycles 1-5 are as defined by Collins et al. (1993). 33 34

1090 00 10 20 30 40 50 10

0 20 40 00 00 D3 Bo yd R iv er Form-surface Map km 90 90

N 60 70 80 90 80 80

70 70

Major roads 60 60 Guyra 50 50

Dorrigo 40 40 Ebor

30 30 Bellingen Armidale 60 70 80 90 20 20 Hillgrove

10 10 Uralla

Generalized trend of S3 00 00

90 90 Major D7 shear zones

80 80

Walcha 70 70 Younger rock units have been excluded for simplicity 60 60

Permian sediments 50 50 Permian Early Permian Volcanics

40 40 Gabbros & diorites Late Carboniferous 30 30 Hillgrove Suite

Accretionary metasediments 20 20 Pre- late Carboniferous Serpentinites oc 10 Nowend 10

00 00 40 50 60 70 80 90 00

. Figure 2 2. D3 form - surface map for accretion complex metasediments. 35

The recognized effects of D4 are restricted to the Tia Complex and hence cannot be correlated across the area. However, for the same reasons that are stated above, D4 fabrics may be developed outside the Tia Complex in areas such as the Yarrowitch Block. It is interesting to note that the E-W fabric produced during D4 is unique within the Tablelands Complex.

Of all the deformation events recorded within the Tia Complex, the penetrative S5 fabrics developed during D5, are the most widespread and continuous across the Wollomombi Zone and into the Coffs Harbour Block (Fig. 2A3). Although D5 fabrics cannot be traced westwards from the Wongwibinda Complex or the Tia Complex into the Walcha-Rockvale Subzone, S5 forms the dominant fabric throughout the Winterbourne Subzone. Unfolding of the F6 folds

(see below) reveals an overall increase in the intensity of D5 fabrics in an easterly direction.

This observation is consistent with the variation in intensity of D5 fabrics within both the Tia Complex (Dirks et al. 1992) and in the Wongwibinda Complex (Farrell 1992). In each case,

S5 intensifies, and gradually becomes the dominant fabric in west to east traverses of these complexes.

D5 is also the first fabric that is developed within granitoids of the Hillgrove Suite. The observation that these fabrics may be traced from the granitoids into the dominant cleavage within the Winterbourne Subzone is further confirmation that this fabric is D5 rather than D3.

However, the intensity of S5 within the granitoids is variable, and generally not strong. S5 is best developed within deeper level intrusives of the suite such as the Tia, Abroi and Argyll

Granodiorites. Within most higher level plutons, S5 is weak or absent, and is often difficult to distinguish from D7 fabrics in the vicinity of D7 shear zones, as the fabrics are often parallel (e.g. in the Abroi Granodiorite). Weak S5 foliations may be observed in the more easterly sections of the Enmore and Blue Knobby Adamellites, and in the southern parts of the Dundurrabin Granodiorite (see Fig. 2A3), where they are parallel to stronger S5 fabrics in the enclosing metasediments. This variation in intensity of S5 throughout the suite, is consistent with the conclusions that the suite intruded late during the D5 event (Dirks et al. 1992, Farrell 1992). 36

1090 00 10 20 30 40 50 10

a pproximate position 00 00 0 20 40 of orocline axis D5 km Form-surface Map 90 90

N 60 70 80 90

r 80 ara 80

70 70

Major roads 60 60

Guyra 50 50

Dorrigo 40 40

Ebor

30 Bellingen 30

Armidale 60 70 80 90 20 20

10 10 Uralla Generalized trend of S5 00 00

90 90 Major D7 shear zones

80 80 F6 fold axis Walcha 70 70 Younger rock units have been excluded for simplicity 60 60 Permian sediments 50 50 Permian Early Permian Volcanics

40 40 Gabbros & diorites Late Carboniferous 30 30 Hillgrove Suite

Accretionary metasediments 20 20 Pre- late Carboniferous Serpentinites 10 Nowendoc 10

00 00 40 50 60 70 80 90 00 . Figure 2 3. D5 form-surface map for Hillgrove Suite granitoids and accretionary metasediments. Contoured densities Poles to C7 of poles to S5 : L7 stretching >1% lineations >2% >4% >8% Max=9.41%

+ +

Equal area projection - 197 points Calculated beta axis 84®275 ( ) 5 L5 stretching lineations 0 F5 fold plunges . Figure 2 5. Plot of poles to D7 mylonite shear (C) planes 7 ® Axial plane of the megascopic F6 fold and L7 stetching lineations ( = spherical mean 66 288) in the Winterbourne area (see figure 2. 3) for Hillgrove Suite granitoids and metasediments.

. 37 Figure 2 4. Contoured plot of poles to S5 - Winterbourne Subzone 38

The effects of D6 are not discernible in the field on a mesoscopic scale, with folding caused by the D6 event, only evident on larger scales. Dirks et al. (1992) described the effects of D6 as large scale folding of all earlier fabrics. F6 folds within the Tia Complex are open vertical folds, with steep plunges, wavelengths of ~15-20 km, and an approximately E-W axial trace.

F6 folding is also evident within the Tia Granodiorite, where the S5 foliation is folded about a similar axis (Dirks et al. 1993). To the north of the Tia Complex, within the Winterbourne

Subzone, the dominant S5 foliation is folded about a major F6 fold axis (see Fig. 2A3) with a similar orientation to F6 axes in the Tia Complex (~110E). There, S5 is folded from a northwesterly orientation immediately to the northeast of the Tia Complex, to a northeasterly orientation further north. This large scale fold structure (with a wavelength of ~40 km) has been previously noted by Haydon (1974). A contoured, equal-area stereo projection of poles to S5 within the Winterbourne Subzone (Fig. 2A4) demonstrates that S5 (which is always near- vertical) is folded about an axial plane of ~110E with a calculated β-axis (plunge) of 846275E.

Further north, S5 within the Abroi Granodiorite is folded about a broad F6 hinge zone, with

S5 within the Wongwibinda Complex changing in orientation from north-south to a northwesterly trend at the northern margin of the complex (Fig. 2A3). North from the

Wongwibinda Complex, within the Tobermory and Kookabookra Adamellites, S5 is generally north-south, but is again folded into a northwesterly orientation further north within the Henry River Adamellite (the most northerly of the Hillgrove Suite plutons - see Fig. 2A3).

The orientation and spatial relationship of S5 within the Coffs Harbour Block cannot be readily explained by the F6 folding of the fabrics of the Wollomombi Zone, since they appear to be folded about an axis which is near N-S rather than E-W. Both the spatial relationship of the Hillgrove granitoids within the Coffs Harbour Block, and the orientation of the associated S5 fabrics, may only be explained by either large scale lateral transposition (dextral) by faulting, or by folding about an axis which is NNW (~160E) in orientation. The anomalous orientation of accretionary fabrics relative to other parts of the accretion complex, and the southwesterly younging direction of successive accretionary packets within the Coffs Harbour Block, has been explained by folding of the entire accretion complex about a large double orocline across the Texas - Coffs Harbour region (Flood & Fergusson 1982). Timing 39 and development of the Texas-Coffs Harbour Orocline (or ‘megafold’) has been discussed by many authors since (e.g. Collins et al. 1993, Harrington & Korsch 1987, Korsch & Harrington 1987, Murray et al. 1987). This ‘double orocline’ defines a Z-shaped megafold which is ~300 km across, with NNW (~160E) trending axial traces (Fergusson & Flood 1982). The axial trace of the western limb of this orocline in the Texas region may be extrapolated further south, and is shown as a dashed line (large dashes) in figure 2A3. This axis bisects the area between the Coffs Harbour Block and the Wollomombi Zone, and is proposed here as the best explanation for both the anomalous spatial distribution of Hillgrove Suite granitoids in the Coffs Harbour Block (c.f. the remainder of the suite), and the orientation of S5 in this block.

Although the orientation of F6 axial traces within the Wollomombi Zone are markedly different to that of the trace of the Texas orocline (110E and 160E respectively) their formation may well have been concurrent, as both events postdate D5 and predate D7. The broader implications, and relative timing of these events will be discussed in more detail below (section 2A4).

The effects of D7 are characterized mesoscopically by the restricted development of intense

S-C mylonitic fabrics within granitoids of the Hillgrove Suite (section 2A3A4). These D7 zones usually occur on the eastern margins of plutons, and so many plutons of the suite are truncated (Fig. 2A6). Major faults include the Tia, Chandler, Kilburnie and WYFS. These fault zones vary in width from 300 metres (e.g. Kilburnie Fault) to 2 km (e.g. Wongwibinda Fault). Faulting has juxtaposed the plutons against low grade metasediments of the accretion complex, or low grade Permian metasediments in which no contact thermal effects are present. In a few cases plutons are juxtaposed against their higher level equivalents (e.g. the

WYFS within the Kookabookra Adamellite). All D7 shear zones show a strong down-dip mineral elongation lineation. Kinematic indicators (details below - section 2A3A4) within these mylonites, indicate a west-over-east, reverse sense of movement on all shear zones, with the exception of the Kilburnie Fault which dips steeply east and shows east-over-west thrusting (see Fig. 2A6). Estimation of total strain on the Wongwibinda Fault, using S-C angle relationships over the width of the shear, suggests a minimum displacement of 8 km (Farrell 1992). The present variation in exposure levels of the Hillgrove Suite and surrounding 40

90 00 10 20 30 40 50 10 10

86 Major D7 shear zones 0 20 40 00 00 with stretching lineations km 32 and inferred tectonic 90 50 90 transport directions

N 60 70 80 90

r 80 ar 80 54 a

Glen 70 60 70

Bluff

Roads 60 Wongwibinda - 60

Guyra Fault 50 50 70 68 Dorrigo 46 40 40 64 63 Ebor

30 30 58 36 Bellingen 72 Armidale 62 60 70 80 90 62 20 64 20 ult 84 a 10 dler F 10 Uralla an Ch 75 Inferred tectonic Borah 00 00 transport direction 45 Fault 40 7 90 Ya 90 L7 stretching lineations

rrowitch

80 80 Major D7 shear zones

Walcha 60 70 70

Tia Fault Younger rock units have been 68 40 excluded for simplicity 60 60 42 Permian sediments 70 50 Kilburnie Fault 70 50 Permian Early Permian Volcanics

Nowendoc Fault Fault 40 40 Gabbros & diorites Late 64 Dingo Carboniferous 30 30 Hillgrove Suite

Fault Accretionary metasediments 20 System 20 Pre- late Carboniferous Serpentinites endoc 10 Now 10

00 00 40 50 60 70 80 90 00

. Figure 2 6. Map of D7 shear zones in Hillgrove Suite granitoids and metasediments, with 7 generalized L7 stretching lineations indicated. Inferred tectonic transport directions are shown where appropriate - the dominant component is reverse in all cases. Named fault zones are 41 metasediments is the result of crustal tilting and uplift during this west-over-east directed thrusting on (D7) fault systems such as the Wongwibinda-Yarrowitch, Chandler, and Tia faults (Fig. 2A6). The best example is in the Wongwibinda Complex, where this westward tilting has exposed a sequence of accretion complex rocks which increase in metamorphic grade eastwards, are eventually intruded by the Abroi Granodiorite, which is in turn truncated by the WYFS.

The effects of D7 within the metasedimentary sequences are more difficult to recognize, mainly due to the intense fabrics developed in these rocks prior to the D7 event. D7 shears within accretion complex rocks which have intense pre-existing fabrics (primarily S5), have often exploited this weakness, and hence have simply reactivated earlier shear zones at similar metamorphic grade, generating a new fabric which is parallel to, and therefore indistinguishable from its precursor. In such cases, D7 shear zones prove difficult to trace laterally beyond faulted pluton boundaries (see discussion above, section 2A3A3). This, in turn, renders the broad scale correlation of fault zones problematic. In most cases, since S5 is a weak fabric in the granitoids, D7 shears are at an angle to the earlier fabric. For example,

S7 in the Tia Granodiorite, where S7 crosscuts the earlier S5 fabric which is folded around an

F6 fold hinge (Dirks et al. 1993). However in the case of the Wongwibinda Complex (Farrell

1992) the Wongwibinda Fault (D7) is parallel to the earlier S5 fabric.

The reactivation of broadly folded (S5) fabrics by D7 shear zones, particularly in metasedimentary sequences, gives the impression that these D7 shears were folded during

D6 (see Fig. 2A6 - in particular the Tia and Chandler faults). This variation in fault strike is 7 also evident in equal area projections of L7 stretching lineations and poles to C7 shear planes throughout the Wollomombi Zone (Fig. 2A5). However, the many cases where D7 shears have crosscut S5 fabrics (and hence the broad structure of F6 folds), both within granitoids (see example above) and metasedimentary sequences (Wongwibinda -Yarrowitch Fault System - see discussion below), unequivocally demonstrate that these shear zones post-date D6.

D8 The subvertical ultramylonite and pseudotachylite veins reported from both the Wongwibinda Complex (Farrell 1992) and the Tia Complex (Dirks et al. 1992) occur in many plutons of the Hillgrove Suite. Although the orientation of these late shears is rather random, they often 42 parallel D7 fabrics, and are more numerous in proximity to D7 shear zones, and hence probably represent the last phases of uplift associated with D7 (c.f. Dirks et al. 1992).

The western boundary of the Tia Complex is defined by the Nowendoc Fault (Gunthorpe

1970), which was tentatively correlated with D8 by Dirks et al. (1992). In the southern part of the Tia Complex, this fault is occupied by serpentinite masses which have recorded several phases of movement (Hand 1988), but further north in the complex, the position of this fault is only inferred by the hiatus in metamorphic grade. Fracturing within schists of the Tia Complex, near the inferred position of this fault (Dirks et al. 1992), indicate an E-over-W sense of shear (consistent with the westward drop in metamorphic grade). The inferred E-over-W reverse movement and implied orientation of this fault, are similar to the Kilburnie Fault which crops out only 10 km to the west, which also shows an E-over-W sense of shear.

This close comparison implies that the Nowendoc Fault has acted is a D7 shear zone (at least in the earlier phases of movement) in this northern part of the Tia Complex.

Deformation in the early Permian basins - the Nambucca Block The degree of deformation manifest within early Permian sequences is highly variable. Permian sequences of the Manning Block (Fig. 2A1) show little apparent mesoscopic deformation, with sporadic development of cleavage (R. Jenkins pers. comm. 1994), while others such as the Nambucca Block are multiply deformed. Leitch (1978) detailed the structural succession in the Nambucca Block, describing five deformation events. Although K-Ar dating of fine-grained metasediments from the Nambucca Block revealed a range of ages, Leitch & McDougall (1979) concluded that the major deformation recorded in these rocks occurred in the interval 250-255 Ma. Realistically, these data provide a minimum age constraint, and either record a cooling age (c.f. discussions in chapter 3) or the age of the youngest deformation event.

The most widespread deformation fabric within the Nambucca Block is S1 (Leitch 1977) S1 within the Nambucca Block has a general E-W orientation (Leitch 1977). Within the western margin of the belt, adjacent to rocks of the Wollomombi Zone (see map - rear pocket) S1 varies in orientation from NE-SW to NW-SE, presumably reflecting the effects of later folding. Both S1 and the folding of this fabric, developed in the Nambucca Block prior to D7 shearing in the Wollomombi Zone, as both are truncated by the WYFS. In addition, since 43 deposition within the Nambucca Block occurred in the early Permian, D1 must also postdate

D5 of the Wollomombi Zone. These constraints place D1 (Nambucca Block) in a broad temporal correlation with D6 of the Wollomombi Zone. Although the timing of D6 is not well constrained, a correlation with early deformation in the Nambucca Block is also supported by the similar orientation of the principle stress fields operating in both cases - that is σ1 is approximately in a N-S orientation. The broader tectonic aspects of deformation within the Nambucca Block will be discussed below (section 2A4). The temporal relationship of the later multiple deformation and fabric development evident in the eastern part of the Nambucca

Block (D2 - D5, Leitch 1978) is unclear, as no direct correlations can be made.

2A4 Conclusions and Regional Implications

Collins et al. (1993) divided the tectonic history of the SNEFB into a series of five tectonic cycles (Table 2A2). The deformation events that have generated tectonic fabrics both within the Hillgrove Suite granitoids and the host accretion complex metasediments, or that have affected the distribution of the suite, occurred during tectonic cycles 2-5. The first tectonic cycle is the subduction/accretion and arc volcanism associated with the Devonian - Carboniferous arc in the Tamworth Belt. This event involved both the outward growth by subduction-accretion, and initial deformation within the complex. This deformation is recorded by subduction-accretion fabrics developed in the Coffs Harbour Block (Fergusson 1982) and other areas of the Tablelands Complex not affected by post-accretion deformation, the D1-D2 blueschist fabrics in the southern Tia Complex (Dirks et al. 1992), and the subvertical aspect of bedding (D0) in areas where accretionary fabrics are not well preserved (e.g. the Wongwibinda Complex).

After the termination of arc volcanism in the Tamworth Belt, parts of the Tablelands Complex underwent a rapid change from high-P/low-T to high-T/low-P metamorphism, which was accompanied by major uplift in the Tia Complex (D3). In the latest Carboniferous, the entire Wollomombi Zone was intruded by granitoids of the Hillgrove Suite and associated mantle derived intrusives with an island arc tholeiite geochemical affinity (see chapter 4).

This intrusive event was accompanied by a compressional event (D5) that uplifted the entire

Wollomombi Zone along a D5 shear zone which dominates its eastern margins. This ductile 44

D5 shear zone occupies the eastern part of the Tia Complex and dominates the Winterbourne

Subzone. Although shearing during the D5 event appears limited within the Wongwibinda

Complex (stretching lineations are only weak or absent, Farrell 1992), the truncation of D5 fabrics at the northern end of the Winterbourne Subzone by the WYFS (see Fig. 2A3) suggests that this wide ductile shear zone may have existed east of the Wongwibinda Complex, and may still exist in basement rocks underlying the Permian sequences of the northwestern parts of the Nambucca Block (the Dyamberin Beds). Little evidence of uplift during D5 can be found in the southern Coffs Harbour Block, where S5 occurs as a weak fabric in the southern part of the Dundurrabin Granodiorite and in the enclosing metasediments, with the distinct absence of any ductile shearing fabrics (i.e. stretching lineations).

After the D5 event and Hillgrove Suite intrusion, the third tectonic cycle was dominated by rifting, accompanied by bimodal volcanism and formation of the early Permian basins throughout the region. Little direct evidence of this event can be found in the accretion complex sequences, although Dirks et al. (1992) noted brittle extensional features consisting of steeply dipping quartz veins which parallel S5 in the Tia Complex, and have a west-down component - the opposite to that of D5.

The fourth tectonic cycle has been correlated with the Hunter-Bowen Orogeny by Collins et al. (1993). This event involved compressive deformation within the accretion complex, and generated folding and faulting within the adjacent Tamworth Belt and in the Sydney Basin (c.f. Collins 1991). The ‘Permian dispersal’ event (Cawood & Leitch 1985), involving large strike slip movements of distinct blocks within the Tablelands Complex also occurred at this time. This dispersal event accompanied formation of the Texas - Coffs Harbour Orocline, and culminated in the climactic uplift recorded by D7 shear zones. The sequence of events described here is depicted in figure 2A7. The Permian dispersal event as depicted in this sequence, began with the northeast transposition and anticlockwise rotation of the Hastings and Nambucca blocks (Fig. 2A7a6d). Formation of the Texas - Coffs Harbour Orocline follows this dispersal and is depicted in the sequence figure 2A7e6f. Orocline formation (together with sinistral strike-slip movement along an unnamed fault to the west of the Kilburnie Fault) can be seen here, to be largely responsible for the current aerial distribution of the Hillgrove Suite. The axis of the western limb of the orocline is depicted (dashed line - Fig. 2A7f) as projecting through the divide between the Coffs Harbour Block and the 45

Hillgrove Suite CB (b) (a) (c) Kaloe Granodiorite

100 km BB

Bundarra Suite 100 km

TB AB Hillgrove Suite

Newcastle N YB

W NB HB Wauchope (f)

(e) (d) BB

CB AB

NB TB YB HB Wauchope

Newcastle

Late Permian dispersal and uplift faults

100 km Early Permian basins Early Permian granitoids Late Carboniferous granitoids Subduction-accretion complex Volcanic arc/forearc Figure 2. 7. Unwinding the New England Orocline, adapted from Collins et al. (1993). The western part of the Yarrowitch Block is part of the Wollomombi Zone (WZ) and separated from it by the WYFS; the WZ is separated by an unnamed fault located farther west than the Kilburnie Fault, and has been reorientated so that major structural trends and Hillgrove Suite plutons are aligned parallel to regional structures; Permian basins are included; the Nambucca Basin was rotated and translated along with the Hastings Block; the WYFS evolved from a strike-slip to dip-slip fault during oroclinal bending. AB Armidale Block; BB Bonshaw Block; CB Coffs Harbour Block; TB Tamworth Belt; HB Hastings Block; NB Nambucca Block; WZ Wollomombi Zone 46

Wollomombi Zone (c.f. Fig. 2A3). Deformation during D6 occurred in the ‘lock-up’ stages of orocline formation, with the broad scale F6 folds depicted as the folding of structures surrounding the Hillgrove Suite plutons in figure 2A7f. It is postulated here, that, in addition to the F6 folding, the development of the widespread S1 cleavage within the Permian sequences of the Nambucca Block occurred at this time (see earlier discussion, section 2A3A3). Alternatively, this cleavage may have developed earlier as a N-S fabric associated with E-W compression during the initial stages of the Hunter Bowen Orogeny, and has since been rotated into an E-W orientation during oroclinal bending (Collins et al. 1993). D6 structures in the Wollomombi Zone and D1 structures in the Nambucca Block, are truncated by the formation of D7 ductile shear zones (Fig. 2A7f), which represents the culmination of this NW-

SE directed deformation. As discussed above, D7 is exclusively an uplift event (with up to 8 km of uplift occurring), involving a westward tilting of the entire Wollomombi Zone, as opposed to the dominantly horizontal crustal movement which preceded it. D7 then, represents the final stages of this progressive sequence of deformation of the fourth tectonic cycle, which culminated in the late Permian. This climactic deformation was followed by the fifth tectonic cycle, which involved minor crustal extension associated with major plutonism involving intrusion of the I- and A-type granites of the NEB (Collins et al. 1993). 47

CHAPTER 3. AGE RELATIONS OF THE HILLGROVE PLUTONIC SUITE

3A1 Introduction

Ductile shear zones (mylonites) have long been recognized as significant features of deformed and metamorphosed terrains and are crucial to regional tectonic interpretations. Clearly, the ability to date the formation of shear zones is highly desirable if chronological constraints are to be placed on regional tectonic events. Which dating method should be applied to the dating of such shear zones, depends on an understanding of the isotope systematics in relation to the processes active during individual shear zone development. Many studies over the last twenty years have applied whole-rock Rb-Sr methods in dating the movement on ductile shears (e.g. Hickman & Glassey 1984). However, the successful use of whole-rock methods in dating ductile shears requires that the rocks be largely recrystallized or exposed to high fluid flow so that the isotopic system is completely reset during shearing (Hickman & Glassey 1984, Shaw & Black 1991). Difficulties are encountered if whole-rock isotopic methods are applied to mylonites that have retained unaltered vestiges of original minerals (e.g. porphyroclasts), and which are likely to have retained their original isotopic ratios. In such cases, the separation and isotopic analysis of neocrystallized minerals in the mylonitic fabric provides a more accurate method for successful isotopic dating of the shearing event. Because these minerals grew during ductile shearing, they should preserve the age of deformation if they grew under metamorphic conditions close to or below the appropriate closure temperature (Cliff 1985). On this assumption, K-Ar ages have been successfully used to date neocrystallized muscovite in ductile shear zones from the Arunta Inlier, central Australia (Dunlap et al. 1991).

This study involves the dating of greenschist and lower amphibolite facies ductile shear zones within Hillgrove Suite granitoids within the tectonically dismembered subduction-accretion complex (Fig. 3A1). Rb-Sr isotopic analysis of both preserved magmatic biotite and neocrystallized biotite from C-planes within mylonites is used to constrain the age of tectonic dispersal and to assess the extent isotopic equilibration during mylonitization.

In addition to age constraints from mylonite zones, regional deformation events are also often bracketed by plutonic events, and can be constrained by using radiometrically dated plutons, 48

Map area

Great Australian Tabl Basin elands Complex New England Fold Belt

Clarence-

Moreton

Basin N

Figure 3. 4

Tamworth

Belt H

M M Lorne Basin Pacific F S T Ocean Sydney Basin 0 50 100 Peripheral basins km Early Permian Rift basins Late Permian - Triassic intrusives and extrusives Arc & fore-arc basin volcanics Early-Middle & metasediments Pre - Permian Permian intrusives sedimentary Accretionary prism sequences Late Carboniferous metasediments } intrusives Figure 3. 1. Locality map of the Hillgrove Plutonic Suite. Inset shows the map area for figures 3.. 4 and 3 5. 49 together with the timing of pluton emplacement relative to each event (Paterson & Tobisch 1988). As such, the Hillgrove Plutonic Suite provides a unique opportunity to tightly constrain the ages of two major deformation events within in the SNEFB, in addition to bracketing others. Granitoid emplacement during D5 (Landenberger et al. 1995, also see chapter 2) allows precise dating of that event by dating pluton emplacement, and in turn the development of mylonite zones within these granitoids during D7, enables precise chronological constraints for this later deformation.

3A2 Previous geochronology of the Hillgrove Suite

Early dating studies on the Hillgrove Suite revealed K-Ar biotite ages of 252-259 Ma (e.g. Binns 1966), which were initially interpreted to be emplacement ages for the suite. Biotite Rb-Sr dating revealed a wider age range from 242 to 296 Ma (Hensel 1982). Whole-rock Rb-Sr dating gave combined isochron ages of 295±25 Ma (Flood & Shaw 1977) to 310±12 Ma (Hensel et al. 1985) and demonstrated that the younger dates represent either a cooling or metamorphic age.

Various explanations for the discrepancy between whole-rock isochrons and biotite ages for the Hillgrove Suite have been postulated. The abundant mesoscopic deformation features and ubiquitous deformation microstructures evident throughout the suite suggest that this post-emplacement deformation may be responsible for resetting the mineral ages (e.g. Flood & Shaw 1977). Resetting of mineral ages by later thermal events has also been postulated (Hensel et al. 1985), however, the absence of adjacent intrusions of appropriate age (252-259 Ma), and the lack of microstructural evidence for thermal resetting, preclude this explanation. Where younger intrusives of the New England Batholith have intruded Hillgrove Suite plutons, causing localized static recrystallization of biotite, both the age of these younger granitoids and the age of recrystallized biotites from the affected plutons are younger than 250 Ma (Shaw & Flood 1981, Hensel 1982).

Although the Rb-Sr whole rock isochrons constructed by Hensel et al. (1985) and Flood & Shaw (1977) demonstrated that mineral ages for the suite represented either cooling or deformational ages, the precise age of crystallization remained poorly constrained due to the large errors involved. These errors are largely the result of variations in initial isotopic ratios 50

Reference Unit Method Age (Ma) Inference Cooper & Richards Hillgrove Adamellite K-Ar biotite 275 Intrusion age (1963) Binns & Richards Abroi Granodiorite K-Ar biotite 257, 212 Intrusion age (1965) Wongwibinda Complex: 259 Migmatite 255 Rampsbeck Schists 264 Tobermory Adamellite Rowley (1975) Bakers Creek Diorite K-Ar biotite 291 Intrusion age Kleeman (1975) Tobermory Adamellite Fission track 273±9 Intrusion age (apatite) Hensel (1982) Hillgrove Suite Rb-Sr biotite 2966241 Thermal (total resetting range) Flood & Shaw Hillgrove Suite ‡ Rb-Sr whole rock 289±25 Intrusion age (1977) Hensel et al. (1985) Hillgrove Suite ‡ Rb-Sr whole rock 312±10 Intrusion age Iizumi & Honma Hillgrove Suite Rb-Sr whole rock 290±4 Intrusion age (1987) recalculation using combined data ‡ above Dirks et al. (1992) Tia Granodiorite Rb-Sr biotite 262±3, Deformation 264±3 age Kent (1994) Rockvale Adamellite U-Pb zircon (SHRIMP) 303±3 Intrusion age

K-Ar biotite deformation 284±6 age

Table 3A1 Summary of previous geochronological work on the Hillgrove Suite and associated rocks. for individual members of the suite (Flood & Shaw 1977), a point which will be discussed later (see chapter 4). Separation of these data into two groups by Iizumi & Honma (1987), lowered the error range to give an age of 290±4 from two separate isochrons with different initial 87Sr/86Sr ratios. However, separation of these data into two groups on the basis of 87Sr/86Sr ratios appears somewhat subjective, as the data form a continuous range of initial isotopic ratios rather than a bimodal distribution. The degree of variation in isotopic ratios is well demonstrated by plotting the data on the ‘best isochron diagram’ of Provost (1990) (Fig. 3A2b), a feature of the data that is not easily discernible on a conventional isochron diagram (Fig. 3A2a). In particular, Fig. 3A2b shows that samples with high 87Rb/86Sr values, have 87Sr/86Sr values which plot below the trend defined by samples with lower 87Rb/86Sr, 51

0.74 O

MaO O

0.73 296±7 O

O O

OO O Sr/ Sr OO O O 870.72 86 O O O

O O OO O OO OO O O OOO O

O 0.71 O

0.70 0 2 4 6 8 10 87Rb/ 86 Sr 0.708 0.7100.712 0.716 0.72 0.73 0.76 0.8 1 ¥

320

0.7064

310

0.7060

300 Age (Ma)

0.7056 Sr/ Sr 87 86 290

0.7052

280

0.7048

270 0 0.5 1.0 2 3 4 6 10 40 87Rb/ 86 Sr Figure 3. 2. Conventional isochron and ‘Best Isochron’ plots (Provost 1990) of Rb-Sr whole rock data for the Hillgrove Suite. Purple lines on the lower diagram are lines of equal 87 Sr/ 86 Sr for the radial X-axis. 52 perhaps due to the low total Sr contents of these samples. Exclusion of samples with 87Rb/86Sr ratios of >10, and exclusion of the Rockisle Granodiorite which belongs to a separate suite (see chapter 4), results in a well constrained isochron of 296±7 (model 3

87 86 isochron, Sr/ Srinitial = 0A7059±A0010, MSWD=259000), an age which is within error of the zircon U-Pb data discussed below. Even so, the best isochron plot (Fig. 3A2b) demonstrates, that even with the exclusion of these samples, initial ratios for the suite vary from ~0A7044-0A7064, which results in the very large MSWD.

Zircon U-Pb dating has had limited application in geochronological and petrogenetic studies in the southern New England Fold Belt. The only U-Pb isotopic study prior to 1990 involved dating of exotic blocks of plagiogranite in serpentinites of the Peel-Manning Fault System (PMFS), along with granitoids of the Nundle Suite, the most southerly suite of the New England Batholith. Results obtained for the plagiogranites (isotope dilution) were generally discordant and regarded as a minimum age (436 Ma, Kimbrough et al. 1993). Zircons from Nundle Suite granitoids also produced poor results (due to low U contents) with an inferred crystallization age of 281 Ma for the Barrington Tops Granodiorite, which contrasts with the biotite K-Ar age for this suite of 264 Ma (Cooper et al. 1963), and biotite Rb-Sr ages of 265±2 Ma and 269±2 Ma (Mason & Kavalieris 1984). Aitchison & Ireland (1995) also reported ages for exotic blocks of plagiogranite in serpentinites of the PMFS. Crystallization ages for these rocks range from early Cambrian to mid-Devonian. These plagiogranites probably represent a dismembered ophiolite sequence of intraoceanic island arc origin (Aitchison & Ireland 1995), and record some of the early history of the fold belt.

The only other U-Pb zircon dating reported for the New England Batholith is that by Kent (1994). SHRIMP (sensitive high resolution ion microprobe) U-Pb data for magmatic zircons from the Rockvale Adamellite (a Hillgrove Suite member) revealed a crystallization age of 303±3 Ma (Kent 1994).

The unreliable results obtained by both mineral (Rb-Sr and K-Ar) and whole-rock isotopic (Rb-Sr) methods for determination of the crystallization age of the Hillgrove Suite, necessitates age determination by isotopic systems which are unaffected by younger metamorphic/deformation events, have high closure temperature, and are insensitive to initial isotopic composition. Therefore, the best method for age determination for the Hillgrove 53

Suite is that of U-Pb zircon analysis. Closure temperatures for U-Pb zircon are much higher (700±50EC, Mattinson 1978) than for Rb-Sr or K-Ar in biotite (320±40EC and 280±40EC respectively, Harrison & McDougall 1980), and hence will be unaffected by the suggested low cooling rates for the granitoids (Landenberger et al. 1995). Additionally, new zircon growth during D7 mylonitization is unlikely at the temperatures (greenschist and lower amphibolite facies) implied for this deformation. The possibility of contamination by inherited zircons is discussed below (section 3A5).

3A3 Structural sequence associated with the Hillgrove Suite - a brief review

Structural studies in the Tia metamorphic complex (Hand 1988, Dirks et al. 1992) reveal early subduction-related deformational events D1 and D2 (blueschist facies), followed by thrusting and uplift during D3 (dated at ~311 Ma by K-Ar white mica, Offler et al. unpublished data). Thermal overprinting of the early blueschist facies fabrics by a high-T/low-P event, began in D3 and continued during progressive uplift D4 - D5. Intrusion of the Hillgrove Suite granitoids and associated mafic intrusives occurred in the late

Carboniferous and was synchronous with the D5 event. The deformation sequence in the Wongwibinda Complex differs in that early blueschist facies fabrics are not developed, with

D1 and D2 equivalent to D3 and D5 (see chapter 2) in the Tia Complex (Farrell 1992).

Although subduction-accretion fabrics are rare, the ubiquitous steep plunge of both F1 and

F2 folds in the Wongwibinda Complex suggests an originally vertical orientation of bedding prior to D1, which is consistent with this sequence forming part of a subduction-accretion complex. Deformation throughout the Wollomombi Zone continued after granite intrusion with broad (10 km scale) open folding (D6) of earlier fabrics, and development of mylonite zones (D7) which truncate many plutons of the Hillgrove Suite, juxtaposing them against rocks of higher crustal level (Landenberger et al. 1995).

The major aims of this section of the study are to constrain both the age of the emplacement of Hillgrove Suite (and hence the age of D5), and the deformation event (D7) that produced mylonitization of the granitoids. Tight constraint of these two deformation events, in turn can be used to place broad constraints on other deformation events within the SNEFB. In addition, an explanation was required for the large range in mineral isotopic ages previously obtained from granitoids of the Hillgrove Plutonic Suite and associated metamorphic rocks. 54

3A4 Sample Selection, Analytical Methods and Results

3A4A1 Zircon U-Pb

Zircons were separated from two members of the Hillgrove Suite through crushing to <212 μm, followed by washing and magnetic (Frantz) and heavy liquid (sodium polytungstate) separation techniques. Zircons were then forwarded to the Research School of Earth Sciences (RSES) at ANU where magmatic zircons were separated, dissolved and analysed by isotope dilution mass spectrometry (IDTIMS) technique. Although refinement of the analyses of these zircons still continues, preliminary results were presented by Collins et al. (1993), and are also presented herein.

The zircon U-Pb ages presented below are preliminary data only, the results forwarded from RSES at ANU (M. Fanning pers. comm. 1994). As the data are still being collected and refined, raw data is as yet not available and is not presented herein. Descriptions of the results to date are as follows.

Tia Granodiorite (Y62): Colourless elongate euhedral zircons dominate the zircon population retrieved from the Tia Granodiorite sample. In general the grains have little or no internal structure. Some grains have cavities that are interpreted as resulting from trapped gas vapour in a relatively rapid cooling process. The zircons are considered to represent a simple igneous population with no evidence for inherited components.

Three fractions were analysed by IDTIMS (isotope dilution thermal ionization mass spectrometry), each consisting of four abraded water clear grains. The fractions weighed 21, 17 and 16 μg respectively. The analyses are slightly discordant with 206Pb/238U ages of 290A4, 291A6 and 295A7 Ma respectively and 207Pb/206Pb ages of 302A9, 301A4 and 301A1 Ma respectively. On a standard concordia plot (Fig. 3A3), the high precision of these analyses is highlighted and the discordant nature gives rise to a slight dispersion in 206Pb/238U ages. The 207Pb/206Pb ages are similar and a discordia regression line has no excess scatter (MSWD of <0A1) but is poorly defined as the analyses effectively plot as a single point. An alternate way to treat the IDTIMS analyses is to assume that the discordance arises from loss of 0.050 Abroi 206 238 Granodiorite Pb/ U (Sample W306) 310 0.049 Preliminary intercepts at 1093±491 Ma and 300±4 Ma (MSWD = 0.8) 305

0.048 300

0.047 295 Tia Granodiorite 290 (sample Y62) 0.046 Preliminary weighted mean 207Pb/ 206 Pb = 302±4 Ma 285 207Pb/ 235 U 0.045 0.32 0.33 0.34 0.35 0.36 0.37 Figure 3. 3. Concordia plot and calculated ages (IDTMS) for magmatic zircons from the Tia Granodiorite and Abroi Granodiorite. These data are preliminary only. 55 56 radiogenic Pb at the present day. This assumption is valid as the 207Pb/206Pb ages are similar and the regression line calculated above has a lower intercept within uncertainty of the origin (-84 + 425 Ma). A weighted mean of the 207Pb/206Pb ages gives 302 ± 4 Ma. It is clear from the IDTIMS analyses and the concordia diagram that the fractions have lost small amounts of radiogenic Pb and are imperceptibly discordant. Thus the weighted mean of the 207Pb/206Pb ages, equivalent to forcing a discordia regression line through the origin gives and age of 301A6 ± 3A6 Ma, and this is interpreted as the crystallization age of the simple population of igneous zircons from the Tia Granodiorite.

Despite no optical evidence for the presence of inherited centres to the zircons from Y62, there are many cases of older inherited components that are optically undetected. A selection of zircons were mounted in epoxy and analysed using SHRIMP (sensitive high mass resolution ion microprobe). SHRIMP ages are calculated from 206Pb/238U because determination of the 204Pb and 207Pb ratios by the ion microprobe is severely hampered by poor counting statistics in Palaeozoic zircons. In general, there is insufficient radiogenic 207Pb to determine 207Pb/206Pb ratios to a suitable degree of precision or accuracy. The rate of decay of 238U is more favourable and determination of radiogenic 206Pb/238U ratios can be made using the measured 207Pb/206Pb ratios as outlined in Compston et al. (1992).

The ion microprobe analyses cluster within uncertainty of concordia and there is little evidence for a significant common Pb content. Correction for common Pb can be made by extrapolation along an inferred mixing line with common Pb. A weighted mean of the radiogenic 206Pb/238U ages has no excess scatter at 288A5 ± 5A2 Ma. This is significantly younger than the 207Pb/206Pb age of 301A6 ± 3A6 Ma as calculated for the IDTIMS analyses. It must be noted that the SHRIMP 206Pb/238U ages overlap with the three IDTIMS 206Pb/238U ages (290.4, 291.6 and 295.7 Ma), and the discrepancy between the 206Pb/238U age of 288.5 ± 5.2 Ma and the 207Pb/206Pb age of 301A6 ± 3A6 Ma arises from the assumptions of concordance. That is the more precise IDTIMS analyses are determined to be discordant, whereas the SHRIMP analyses are apparently within uncertainty of concordia.

Abroi Granodiorite (W306 - sample submitted by T.R. Farrell)

The zircons from the Abroi Granodiorite, sample W306, are generally colourless grains, 57 elongate to equant in outline with bipyramidal terminations, or fragments of such grains. In section, some grains can be seen to be zoned from centre to rim, others are remarkably structureless, while yet others contain inclusions of apatite and possible trapped gas vapour trails. The zircon population is simple, homogeneous and considered to show little evidence of inheritance.

Five fractions have been analysed by IDTIMS, consisting of 2 to 4 zircons each, and weighing between 6 and 12 micrograms per fraction. The smaller sample sizes have less total Pb per analysis and so the results are not as precise as the Tia Granodiorite IDTIMS data. The results are plotted on a standard concordia plot (Fig. 3A3). There is some dispersion in the analyses with the 206Pb/238U ages ranging from about 293 to 306 Ma. Two of the analyses appear reversely discordant with 207Pb/206Pb ages of 268 and 281 Ma respectively (206Pb/238U ages of 293 and 297 Ma), though the analyses are within uncertainty of concordia at the 95% confidence level. The array of the data points on the concordia plot suggests the presence of a subtle, small inherited component in the fractions analysed. A regression line fitted to all five analyses has no excess scatter giving concordia intercepts at 1039 ± 491 Ma and 300 ± 4 Ma (MSWD of 0A8). The lower intercept is identical to the weighted means of the 206Pb/238U ages at 301 ± 2 Ma and 207Pb/235U ages at 301 ± 2 Ma. Despite the range in 207Pb/206Pb ages (270 to 325 Ma), the analyses are not as precisely defined and are effectively concordant and within uncertainty at 300 ± 4 Ma. The inferred inheritance will be checked using SHRIMP.

3A4A2 Biotite Rb-Sr

Both relict igneous biotite and fine-grained recrystallized (mylonitic) biotite derived from

C-planes were extracted from granitoids deformed in D7 shear zones to determine the age of shearing and uplift. Biotite and muscovite were also extracted from granitoids and high-grade metasediments not directly affected by the D7 shearing to clarify the thermal history of the granitoids and surrounding metamorphic complexes, and to resolve the inconsistencies in previously published mineral ages for the suite. After dissolution, spiked separates were column separated by ion-exchange methods, dried and analysed on a VG354 7-collector mass spectrometer operating in fully automated mode, at the Centre for Isotope Studies, North Ryde, NSW. 58

Model biotite ages were calculated for most samples assuming an initial Sr87/Sr86 ratio of 0A7060, which is the average isotopic ratio for granites and psammitic metasediments at the time of Hillgrove Suite intrusion (Flood & Shaw 1977, Hensel et al. 1985). Although whole-rock isotopic ratios would elevate with time, the high 87Rb/86Sr ratios of the analysed samples would suppress any variation in the calculated age if different initial ratios are used in the calculation. Because of likely higher initial ratios for the single pelite sample analysed, the biotite age for this sample was calculated using whole-rock/biotite paired analyses.

Deformed magmatic biotite extracted from high grade D7 mylonites from shear zones adjacent to the deeper level metamorphic complexes (Wongwibinda Fault, Fig. 3A4) range in age from 258-266 Ma (Table 3A2). Recrystallized (mylonitic) biotite defining C-planes, yields ages either within error of the relict igneous biotite from the same bulk sample (W445F 258A5±2A1 Ma, W445 260A8±2A1 Ma), or slightly younger (W444F 260A5±2A4 Ma, W444 265A8±2A1), hence defining a slightly narrower age range. This slight variation in calculated ages for mylonitic and relict igneous biotites may relate to differences in diffusion rates of rubidium and strontium, combined with the larger distance over which diffusional equilibration of isotopes must occur in the larger relict igneous grains. This explanation is supported by the larger total strontium content of the fine-grained micas (Table 3A2), which suggest an influx of strontium-bearing fluids during shearing, and emphasises the importance of fluid transport to isotopic re-equilibration in shear zones (Hickman & Glassey 1984).

Analyses of a two-mica granite (W442) from migmatites within the Wongwibinda Complex, in rocks unaffected by D7 shearing, give ages within error of those obtained from mylonitized samples. The sample preserves the syn-emplacement (D2) foliation that is dominant throughout much of the Wongwibinda Complex and shows no evidence of post-emplacement deformation or metamorphism. In addition, both biotite and muscovite from this sample give ages within error of each other (260A4±2A1 Ma and 264A6±2A2 Ma respectively). Since the intrusive age is constrained to 296-306 Ma in age by U-Pb zircon data, the mica ages for these undeformed rocks must represent cooling ages. Given the differences in the Rb-Sr closure temperature for these two minerals (320±40EC for biotite, Harrison and McDougall 1980, 500±50EC for muscovite, Jäger 1979) the close agreement of biotite and muscovite ages also implies that final cooling was rapid, hence implying rapid exhumation. 59

Coffs Harbour Block

251 257 30° 00' S 243

268 - 277 tem

258 - 266* 270 Sys

247 260 Dorrigo WC 251 266 251 263 274 241 Armidale 277 259 ult 279 Fa 284 259 296 lt au 296 F le Nambucca Block nd 266 286 ha r 290 C 292 Borah 281 Fault

tc i

o r

r Wollomombi a Kilburnie Y

Fault Zone - 275 T ia 31° 00' S F a u

lt e

289 262 n o

N Z

o 264

w a e

r e 288 271 269 d u n

n t d i u

o b Yarrowitch S i c Hastings Block

w Block s 292 TC g a 277 Dingo F r n

a r u o a

lt Y

W Fault Port Macquarie 272

Tasman Sea 152° 00' E 0 50 Late Permian fault systems km Early Permian rift basins N Bakers Creek Suite Hillgrove Suite granitoids Figure 3. 4. Map of the Hillgrove Plutonic Suite and associated mafic complexes with sample locations and Rb-Sr biotite ages for samples analysed (bold type). Younger rock units have been excluded for simplicity. Data in italics are from Dirks et al. 1992b and Hensel 1982. * Represents the age range obtained from six samples from a traverse of the Wongwinbinda Fault. Sample sites for zircon U-Pb dating are indicated as . 60

Sample Description Grid Rb Sr Rb87/Sr86 Sr87/Sr86 Age(Ma) Reference ppm ppm W308 Mylonite, Wongwibinda Fault 41356554 755A34A362 612A92 2A991099 262A1±2A1 W306 Mylonite, Wongwibinda Fault 41316553 567A03A261 614A03 2A960864 258A1±2A1 W444 Mylonite, Wongwibinda Fault 41396559 701A13A533 728A76 3A461961 265A8±2A1 W444F Mylonite, Wongwibinda Fault 41396559 500A066A39 21A957 0A787370 260A5±2A4 W445 Mylonite, Wongwibinda Fault 41396559 534A01A679 1381A15A829398 260A8±2A1 W445F Mylonite, Wongwibinda Fault 41396559 531A626A47 59A344 0A924261 258A5±2A1 A65 Mylonite, Chandler Fault 40006172 639A64A027 612A37 3A019447 265A5±2A1 C22 Mylonite, Chandler Fault 40486205 557A55A007 400A59 2A182796 259A1±2A1 N1 Mylonite, Kilburnie Fault 35055547 706A94A903 494A98 2A614675 271A0±2A2 N2 Mylonite, Dingo Fault 35115338 403A229A96 43A048 0A885072 292A3±2A4 N3 Mylonite, Dingo Fault 35175327 506A319A91 75A749 1A010608 282A6±2A3 A180 Gara Adamellite, unfoliated, 38686164 538A18A071 229A50 1A674903 296A2±1A7 mildly deformed A60 Hillgrove Adamellite, 39156213 743A03A788 808A92 3A894969 277A1±2A2 unfoliated, deformed W442 Migmatite, WC, deformed 41356497 806A86A014 451A61 2A378722 260A4±2A1 W442M Migmatite, WC, deformed 41356497 273A819A73 40A734 0A859349 264A6±2A2 W443 Hornfels, WC, undeformed 41286468 684A74A959 464A85 2A379380 250A8±2A0 W443M Hornfels, WC, undeformed 41286468 211A2 140A04A3758 A735480 276±10 W443* Hornfels, WC, undeformed 41286468 191A790A86A1292 0A742357

Table 3A2. Biotite Rb-Sr data from Hillgrove Suite granitoids and high grade metasediments. Sample number suffixes refer to: F = fine-grained recrystallized biotite, M = muscovite, * = whole rock values for W443. WC = Wongwibinda Complex. Errors on ages are calculated from external errors of 0A8% for Rb87/Sr86 ratios and 0A008% for Sr87/Sr86 ratios. All ages calculated using an initial Sr87/Sr86 ratio of 0A7060, except for W443 for which a whole rock analysis was paired to give an initial ratio of 0A72049. Grid references are Australian Map Grid (AMG). Note that samples with the prefix ‘W’ were analysed in conjunction with Farrell, and were presented by Farrell (1992). All tabulated data above were also presented in Landenberger et al. (1995).

In contrast to these high-grade mylonites, samples from the higher-level D7 mylonite zones (e.g. Kilburnie Fault & Dingo Faults) which have deformational fabrics and retrograde mineralogy consistent with shearing under greenschist facies conditions, have relict igneous biotites that range in age from 271 Ma to 292 Ma, suggesting variable degrees of isotopic resetting. The restricted growth of recrystallized (mylonitic) biotite in these samples precluded its separation for analysis. 61

Those Hillgrove Suite granitoids not affected by D7 mylonitization preserve a range of biotite ages from 260 Ma to 296 Ma (Table 3A2, Fig. 3A4). This range encompasses most previously reported mineral ages (Binns 1966, Dirks et al. 1992, Hensel 1982). The highest level plutons (contact aureole plutons) that are most distant from the D7 shear zones, preserve biotite ages of 296±2 Ma (Gara Adamellite). These ages are within error of the emplacement age of ~300 Ma (zircon U-Pb, see above), suggesting rapid cooling (due to higher emplacement levels) to temperatures below the biotite Rb-Sr blocking temperature. In contrast, biotites from granitoids and metasediments in the deeper level metamorphic complexes, record ages identical to those of the D7 shear zones. Distinct from their counterparts at higher crustal levels, these granitoids were emplaced in rocks which were above the Rb-Sr biotite blocking temperature. The preserved biotite ages show that these rocks did not cool to below this closure temperature until after uplift, when they were juxtaposed against cooler rocks at higher crustal levels. Previously published and unpublished mineral ages that are <250 Ma are from rocks affected by static recrystallization related to nearby intrusion of younger granites (225-250 Ma; Shaw et al. 1991).

The pelitic hornfels sample from within the Wongwibinda Complex (W443) has textural modification (obliteration of earlier cleavage and development of randomly oriented biotite) consistent with static recrystallization which may be associated with localized intrusion of lamprophyre dykes which post-date D7 (Ashley et al. 1993), and hence yields a younger age (~251 Ma).

3A6 Discussion

The preliminary zircon U-Pb ages presented herein, provide the first tight constraints on the age of emplacement and crystallization of the Hillgrove Suite granitoids. Earlier dating attempts using mineral Rb-Sr and K-Ar techniques only represented cooling or deformational ages, and whole-rock Rb-Sr techniques were hampered by variations in initial isotopic ratios for the suite. The intrusive age of ~300 Ma for the suite is supported by the single U-Pb SHRIMP age (303±3 Ma) for the Rockvale Granodiorite published by Kent (1994). As such, this age also constrains the age of the D5 event to ~300 Ma. 62

The 258-266 Ma Rb-Sr ages recorded in biotites from high-grade D7 mylonite zones represent a cooling age for these rocks (Landenberger et al. 1995). The fabrics and retrograde mineralogy developed in these mylonites (see Chapter 2) imply that temperatures remained above 450EC (c.f. Tullis 1983) at the time of shearing, placing these rocks well above the Rb-Sr blocking temperature for biotite. This conclusion is supported by the indistinguishable biotite ages recorded in the migmatites not affected by D7 shearing within the adjacent Wongwibinda Complex. However, the indication of rapid cooling within the complex, given by the similarity of biotite and muscovite Rb-Sr ages from the migmatite (sample W442), suggest that these cooling ages are a good approximation to the age of shearing and uplift. These data therefore constrain the age of D7 mylonitization and the associated uplift to the late Permian.

The lower temperatures inferred from the deformational textures and retrograde mineralogy in sheared rocks from the Kilburnie and Dingo fault zones are reflected in the wide-ranging, older biotite ages (271-290 Ma, Table 2). In addition, variation in the degree of dynamic recrystallization and annealing of quartz ribbons within these shear zones, implies local variation in temperature or strain rates during shearing, and correlates directly with the ages obtained from these rocks. The ages obtained suggest that shearing at this higher crustal level took place at, or just below the Rb-Sr closure temperature for biotite, and the degree of resetting is highly dependent on the degree of recrystallization and dislocation density. As such, the minimum biotite Rb-Sr ages from these fault zones have no geological significance (Landenberger et al. 1995).

The kinematics and geometry of these late Permian fault zones imply that the prime mechanism involved in the exhumation of higher grade terranes of the SNEFB was reverse faulting that caused crustal tilting, leading to the exposure of crust from a depth of ~10-12 km. This crustal tilting is likely to have played a major role in producing the large range of Rb-Sr biotite ages (257-297 Ma) for granitoids not directly affected by D7 mylonitization. Temperature estimates of >500EC for the Tia and Wongwibinda complexes at the initiation of uplift (Dirks et al. 1993, Farrell 1992) suggest that the extreme geothermal gradient (~70EC/km) responsible for amphibolite facies metamorphism and granitoid production at ~300 Ma only slowly waned by the time of uplift when the geothermal gradient was still in excess of 50EC/km. An extended thermal anomaly has published (Dirks granitoids which have not undergone mylonitization during D . Contours account for previously for account Contours . D during mylonitization undergone not have which granitoids Figure 1 0' S 00' 31° Armidale Age - contour map of the Rb-Sr ages 5. of biotite separates extracted from Hillgrove Suite 3 . hra resetting Thermal et al. 1992, Kent 1994) and unpublished data (Hensel 1982).

152° 00' E 0 hra resetting Thermal 20 a Teml resetting Thermal - Ma <250 abca Block Nambucca ae ema fut systems fault Permian Late basins Permian Early Ma 285 > Ma 285 - 275 275Ma - 265 Ma 265 - 253 7 km 0 0' S 00' 30° 50 Dorrigo N 63 64 previously been postulated by Dirks et al. (1992) and Farrell (1992). Such a protracted thermal event is supported by the widespread magmatic activity throughout the SNEFB during this time, namely intrusion of the voluminous Bundarra Plutonic Suite (Flood & Shaw 1977), along with widespread rifting and associated bimodal volcanism in the early Permian (Leitch 1988) and intrusion of the Barrington Tops Granodiorite into Tamworth Belt rocks. This slow thermal relaxation, combined with crustal tilting at ~260 Ma has produced a pattern of biotite ages which decreases from 296 Ma to 257 Ma, eastwards over a distance of ~50 km (Fig. 3A5), as the major late Permian shear zones are approached (Landenberger et al. 1995). Schematic cross sections of the Wollomombi Zone prior to the late Permian uplift, and at the present (Fig. 3A6) illustrate this effect.

The new Rb-Sr biotite ages from mylonite zones within granitoids of the Hillgrove Suite provide important chronological constraints on the tectonic evolution of the SNEFB. These results not only constrain the age of this deformation to the late Permian, but also explain the large range in Rb-Sr ages in Hillgrove Suite granites not mylonitized in D7. 65

(a)

Pre- late Permian surface 0 Nambucca Block

Wollomombi Zone

Wongwibinda Complex Approximate Vertical scale (km) 10

(b) Wongwibinda Gara Adamellite Complex WYFS Present surface

Nambucca Block Wollomombi Zone

253 - 265 Ma Early Permian basin (Nambucca Block)

265 - 275Ma Late Permian fault systems 275 - 285 Ma 0 50 Approximate > 285 Ma Horizontal scale km Figure 3. 6. Schematic cross-sections of the Wollomombi Zone. (a) Immediately prior to the ~260 Ma uplift. The slowly contracting geotherm has produced a vertical zonation of cooling ages within plutons of the Hillgrove Suite. At this stage, those granitoids depicted as recording ages in the range 263-265 Ma, together with rocks of the Wongwibinda Complex, remain above the biotite Rb-Sr blocking temperature. (b) Uplift along the WYFS at ~260 Ma has tilted the entire Wollomombi Zone westwards, rapidly cooling the higher grade rocks of the Wongwibinda Complex and associated plutons, and producing the present horizontal age zonation pattern depicted in Fig. 3. 5. 66

CHAPTER 4. PETROGENESIS OF THE HILLGROVE PLUTONIC SUITE

4A1 Introduction

Two distinct types of granite were originally recognized in the Kosciusko region of the Lachlan Fold Belt (LFB) by Dallwitz (in Ball et al. 1948), who noted that much of the batholith was composed of foliated granites which bore a red-brown to buff biotite, while other younger plutons were massive and contained chocolate brown to yellow biotite. He suggested that these two granite types may have a different genesis, with the foliated granites resulting from granitization of the enclosing metasediments. More detailed studies of granites of the Kosciusko region led to the subdivision of these granites into S-type and I-type (Chappell & White 1973). This subdivision was accompanied by the genetic inference that S-types, because of their peraluminous chemistry, were derived by the partial fusion of meta-sedimentary rocks; while the I-types, metaluminous in character, were sourced from a meta-igneous precursor. This subdivision has subsequently been extended, and encompasses most granites of the Lachlan Fold Belt (Chappell & White 1992a), with the exception of minor volumes of A-type granites (see discussion Chapter 5). The I-S classification has also since been applied to batholiths worldwide, including the granitoid suites of the New England Batholith. A summary of the major differences between S- and I-type granites is presented in Table 4A1.

Characteristic I-type S-type

Distinctive Biotite (Chocolate brown to Aluminous biotite (‘foxy’ red mineralogy yellow) brown to buff), ilmenite, magnetite ± ilmenite ±almandine, ±cordierite, ±hornblende, ±CPX, ±sphene ±muscovite Major element chemistry High Na, Ca; generally low K Low Na, Ca; high K Al<[Ca+Na+K] (metaluminous) Al>[Ca+Na+K] (peraluminous) <1% normative corundum >1% normative corundum Isotopic 87 86 87 86 composition Sr/ Sri 0A704060A7120 Sr/ Sri 0A709460A7180 gNd -8A96+0A4 gNd -9A26-5A8 Inferred source rocks Meta-igneous crustal basement Meta-sedimentary (infracrustal) (supracrustal)

Table 4A1. Characteristics of S- and I-type granites (data from Chappell & White 1993, McCulloch & Chappell 1982, White & Chappell 1983). 67

4A1A1 S-type granites in the SNEFB

The voluminous New England Batholith of the southern New England Fold Belt is comprised of numerous granitoid suites which span most of the granite compositional spectrum. Two granitoid suites of the batholith have traditionally been regarded as S-type in character. The Hillgrove Plutonic Suite (late Carboniferous) and the Bundarra Plutonic Suite (early Permian) were originally designated as S-type by Flood & Shaw (1977) and O’Neil et al. (1977). The rationale behind this designation lies primarily in the mineralogical makeup of the suites: they contain aluminous biotite, accompanied by occasional almandine garnet and/or cordierite, have ilmenite as the sole oxide phase, and are generally free of amphibole. Most members of these suites are also mildly corundum normative, which also infers a metasedimentary source.

Despite this classification, these two ‘S-type’ suites are isotopically primitive compared to the classical S-types granites of the Lachlan Fold Belt, and have a bulk rock chemistry which is somewhat transitional (see section 4A4) between the classical Lachlan Fold Belt I- and S- types (White & Chappell 1983). These differences have been attributed to the chemical immaturity of the sedimentary source rocks involved in magma generation (e.g. Flood & Shaw 1977). While this chapter deals primarily with the origin of the Hillgrove Suite granitoids, comparisons are made with the Bundarra Plutonic Suite and with the classical S- type granites of the Lachlan Fold Belt.

4A1A2 Previous petrogenetic studies of the Hillgrove Suite

The first published geochemical data (averaged major element compositions) and Sr isotopic data for Hillgrove Suite granitoids were presented by Flood & Shaw (1977). From the available data, these authors broadly identified the source rocks for both the Hillgrove and Bundarra plutonic suites, as the ‘pelitic’ volcanic-arc derived detritus which constitutes much of the accretionary prism of the SNEFB (Flood & Shaw 1977). This conclusion was reiterated in Shaw & Flood (1981), where the authors suggested that the differences between the Hillgrove and Bundarra suites (see table 4A2, below) were due to differences in exposure levels and the degree of partial melting. 68

The doctoral thesis of Hensel (1982, chapter 3) presented the first major mineralogical, geochemical and Sr isotopic study of the Hillgrove Suite (together with other suites of the batholith in separate chapters). This work, with the addition of the first Nd isotopic data for New England granites and host metasediments was published in 1985 (Hensel et al. 1985). These studies further restricted the range of source rocks suitable for producing the New England S-type granites, with the conclusion that these suites could be derived from ‘felsic’ greywackes of the accretionary prism (Hensel et al. 1985).

However, the conclusions drawn form these studies are largely based on the isotopic data alone, with little attention paid to major and trace element considerations, or to modelling of the processes involved. The specific source rocks involved remained poorly defined as ‘felsic greywackes’ or ‘pelitic’ rocks, with no analyses published other than the isotopic data. In addition, the geochemical, mineralogical and isotopic variation within the Hillgrove Suite (e.g. Flood & Shaw 1977), along with differences between the Hillgrove and Bundarra suites, demand that a range of source rocks and/or magma generation processes were involved in the generation of these S-type granites.

4A2 Defining the Hillgrove Suite

4A2A1 Previous definitions of the Hillgrove Suite

The belt of deformed granitoids cropping out to the southeast of Armidale was originally separated from the ‘massive’ plutons of the New England Batholith and named the Hillgrove Plutonic Suite by Binns et al. (1967), who made this distinction on the basis of the ‘stressed’ nature of the plutons involved. However, the definition of the Hillgrove Suite has changed several times since this initial distinction, and in the light of petrogenetic considerations presented herein, a new definition will be presented, along with detail of previous changes made to the definition of the suite.

The Hillgrove Suite definition of Binns et al. (1967) included most of the plutons in the definition presented herein, with the exception of plutons that were either not mapped, or not recognized as members of the suite (the Gostwyck, Murder Dog, Henry River and Glenifer adamellites). Pogson & Hitchins (1973) extended the Hillgrove Suite to include all but the 69

Gostwyck Adamellite, and in addition incorporated the unnamed group of gabbroic and dioritic rocks which were spatially associated with the suite. However, these mafic rocks were subsequently omitted from the definition because of their gross chemical dissimilarities with other members of the suite as pointed out in Flood & Shaw (1977). In addition, the Linden Hill Adamellite (near Walcha), included in the suite by Pogson & Hitchins (1973), more likely belongs to the Uralla Plutonic Suite, since it bears primary hornblende and has an early Triassic age (246 Ma, biotite Rb-Sr, Hensel 1982).

A further modification to the definition of the Hillgrove Suite was suggested by Hensel (1982), who excluded several plutons from the suite, placing them in a group of ‘transitional’ granitoids, which are described below. Hensel (1982) suggested that the mineralogy and chemistry of these granitoids, which have characteristics transitional between the Hillgrove Suite and the younger I-type granitoids (Uralla Suite), warranted their exclusion from the suite. These plutons included the Rockisle Granodiorite and Murder Dog Adamellite (transitional Group A), the Gara and Gostwyck adamellites (transitional Group B), the Billys Creek Tonalite (transitional Group C) (Hensel 1982). The exclusion of these plutons from the suite was reinforced, but somewhat modified by Hensel et al. (1985). These modifications are detailed and discussed as follows:

Group A transitional plutons: The Rockisle and Murder Dog plutons were included in the Nundle Plutonic Suite, a suite of low-K granitoids which crops out near the Peel-Manning Fault System. This was based on the primitive isotopic nature and moderately low-K composition of the Rockisle Granodiorite, together with the spatial relationship of these two plutons to other Nundle Suite members (Hensel et al. 1985). However, there are several inconsistencies between these plutons and other members of the Nundle Suite. Although some samples of the Rockisle Granodiorite are low in K2O, contents vary widely within the pluton (2A1863A67 - see geochemistry section below). The samples lowest in K2O are those bearing secondary amphibole, and they occur primarily along the mylonitized eastern margin of the pluton. There is also a large age contrast between the Rockisle Granodiorite from members of the Nundle Suite. The Nundle Suite varies in age from 262 Ma for the Barrington Tops granodiorite (Hensel et al. 1985) to 245 Ma for the Duncans Creek Trondhjemite (Shaw & Flood 1993), whereas the Rockisle Granodiorite has partially reset Rb-Sr biotite ages varying from 277-292 Ma (Landenberger et al. 1995), which is more 70 typical of the Hillgrove Suite. The Murder Dog Adamellite bears all of the characteristics of the Hillgrove Suite, and even though secondary amphibole is present in some specimens, particularly microgranitoid enclaves, garnet has also been observed (see section 4A3). This variation is typical of that observed within many Hillgrove suite plutons, and exclusion from the suite based on the presence of secondary amphibole is unwarranted.

Group B transitional plutons: The Gara and Gostwyck adamellites were retained as a group of ‘transitional’ granitoids based on their chemistry and mineralogy. Hensel (1982) described brown amphibole and fluorite bearing phases of the Gara Adamellite from near its western margin. However, these appear to be restricted in occurrence and are likely to be fractionated derivatives of more mafic phases of the pluton. The isotopic and chemical composition of these ‘transitional’ granitoids is indistinguishable from that of other Hillgrove Suite plutons (Hensel et al. 1985, and data herein - see below), and petrographic differences are minor and restricted.

Group C transitional plutons: The Billys Creek Tonalite (which crops out just to the north of the Sheep Station Creek Complex in the Coffs Harbour Block), has since been renamed the Billys Creek Quartz Monzonite (Hensel et al. 1985), and is unrelated to the Hillgrove Suite. This pluton was placed in a group of ‘Post-Orogenic Granitoids’ on the basis of a Rb- Sr biotite age of 230 Ma (Hensel et al. 1985), and is more likely related to the nearby Chaelundi Complex, which has a similar age (see chapter 5).

In summary, with the exception of group C, the criteria used by Hensel et al. (1985) to group these plutons as ‘transitional’, or to include them in other suites, are largely superficial. Many of the differences described for these rocks occur as variations within most plutons of the suite. Moreover, most of the ‘transitional’ plutons are geochemical and isotopically indistinguishable from other members of the suite (see below), with the exception of the Rockisle Granodiorite.

Gilligan et al. (1992) regrouped the Hillgrove Suite granitoids together with other intrusives, into their Deformed Granitoid Belt. On the basis of their spatial and/or temporal relationships, this belt included the Hillgrove Suite, the gabbroids and diorites spatially associated with the suite, a group of low-K granitoids (see 4A2A2 below), and several small 71 high-K granite stocks and lamprophyre dykes which crop out in the Rockvale area. The criteria used for the exclusion of the Gara and Gostwyck adamellites from the Hillgrove Suite (Hensel et al. 1985) were also questioned by Gilligan et al. (1992), who reassigned these plutons as Hillgrove Suite members. The Deformed Granitoid Belt as defined by Gilligan et al. (1992), closely resembles the definition of the Hillgrove Supersuite presented herein. However, the inclusion of high-K granitoids in this belt is not warranted, since, although they bear spatial relationship to the belt, they postdate other members of the group, having intrusive ages of <260 Ma (Kent 1994). On this basis, they should not be included in the Hillgrove Supersuite.

4A2A2 Definition of the Hillgrove Supersuite

The term suite has been applied to groups of granitoids (particularly within the Lachlan Fold Belt, e.g. the Jindabyne Suite, Hine et al. 1978) which have geochemical, mineralogical and isotopic coherence, thus implying that these suites are cogenetic. This has the implication that all members of a particular suite are derived from a specific source rock composition (Chappell et al. 1988). As such, suites are a first order subdivision, and are restricted to containing one particular granite type (i.e. they cannot contain both S- and I-type granites). A second order subdivision, called supersuites, has also been applied to granites of the Lachlan Fold Belt (e.g. the Boggy Plain Supersuite, Wyborn et al. 1987). Supersuites were vaguely defined by Chappell et al. (1988) as groups of suites/granitoids which share broad geochemical features, but differ in detail. As such, members of a supersuite need not necessarily be cogenetic, and may be derived from a combination of sources of similar chemical character. This close relationship also implies that members of a supersuite are coeval. The Boggy Plain Supersuite (Wyborn et al. 1987) contains both gabbroic rocks and more evolved I-type granitoids, implying that multiple sources (mantle and lower crustal) were involved in the generation of the supersuite, and that mixing occurred between coeval magmas of these contrasting sources (Wyborn et al. 1987).

A similar usage of the term supersuite is employed herein to define the Hillgrove Supersuite. Although no supersuites of the Lachlan Fold Belt contain both S- and I-type granites, the Hillgrove Supersuite contains a wide variety of rock-types which span the complete range of late Carboniferous granitoids. This range includes the S-type granites of the Hillgrove 72

Suite, and also a group of granitoids which share many of the characteristics of the Hillgrove Suite, but which are amphibole-bearing (the Rockisle Suite - new name). These two suites are here grouped together as the Hillgrove Supersuite. The group of mantle derived gabbros and diorites which bear a close temporal and spatial relationship with the Hillgrove and Rockisle suites (as noted in previous studies, e.g. Leitch et al. 1971), is here termed the Bakers Creek Suite (new name). Although this suite has also played a role in the petrogenesis of the Hillgrove Supersuite (see section 4A5A5), its predominantly mantle derived character precludes inclusion in this supersuite.

Although the name Hillgrove Supersuite has been used by Chappell (1994), no definition was provided, and so a new definition is presented here. The distribution of the Hillgrove Supersuite and its member suites is depicted in figure 4A1, and the broad characteristics of the three member suites are listed in table 4A2. These characteristics are discussed in more detail below.

SUPERSUITE HILLGROVE

SUITE Hillgrove Rockisle Bakers Creek

Member plutons or All Plutons designated Rockisle Granodiorite All mafic complexes complexes as Hillgrove Suite on & designated as Fig 4A1 amphibole-bearing Bakers Creek Suite on phases within Hillgrove Fig 4A1 Suite plutons.

Amphibole - rich granitic members of mafic complexes

Rock-types adamellites and granodiorites & minor Gabbros & diorites, granodiorites adamellites minor granophyres

Distinctive Biotite adamellites and Hillgrove-like Tholeiitic gabbros (high character granodiorites granodiorites and alumina tholeiites) and ± trace almandine adamellites bearing contaminated (amphibole free) trace secondary and/or fractionated sub-aluminous equivalents calcic amphibole.

Primary hornblende- bearing granitoids of mafic complexes

Table 4A2. Characteristics of Hillgrove Supersuite and Bakers Creek Suite members. 73

1090 00 10 20 30 40 50 10

0 20 40 00 3 00 Distribution of the Hillgrove Supersuite km 90 90

N 60 70 80 90 80 80 4

70 70 27 26 Major 60 60 roads 28 5 25 Guyra 2 50 50

Dorrigo 40 40 29 24 1 7 Ebor 6 30 30 Bellingen

Armidale 60 70 80 90 20 8 20 Hillgrove 10 9 10 30 10 Uralla 11 00 00 31 12 33 LEGEND 90 32 90 13 14 80 80 Post-Triassic faults 15 Walcha 70 34 70 Late Permian faults 16 60 19 35 60 Tertiary Basalt 17 Triassic Granites & volcanics 18 50 21 50 20 Early Permian Sediments and volcanics

40 40 Bakers Creek Suite Late 23 Carboniferous 30 30 Rockisle Suite Hillgrove Supersuite Hillgrove Suite 20 20 Accretion complex 22 Nowendoc Pre- late Carboniferous metasediments 10 10 and metavolcanics

00 00 40 50 60 70 80 90 00

Figure 4. 1. Member plutons of the Hillgrove Supersuite (see text for extended definitions for each suite) Hillgrove Suite Bakers Creek Suite 1 Glenifer Adamellite 13 Eastlake Adamellite 24 Dorrigo Mountain Complex 2 Dundurrabin Granodiorite 14 Winterbourne Granodiorite 25 Charon Creek Diorite 3 Henry River Adamellite 15 Argyll Granodiorite 26 Sheep Station Creek Complex 4 Kookabookra Adamellite 17 Garibaldi Adamellite 27 Mornington Diorite 5 Tobermory Adamellite 18 Tia Granodiorite 28 Day's Creek Gabbro 6 Abroi Granodiorite 19 Kilburnie Adamellite 29 Camperdown Complex 7 Rockvale Adamellite 20 Campfire Adamellite 30 Baker's Creek Diorite 8 Hillgrove Adamellite 21 Ingleba Adamellite 31 Barney House Gabbro 9 Gara Adamellite 22 Murder Dog Adamellite 32 Woodburn Diorite 10 Gostwyck Adamellite Rockisle Suite 33 Cheyenne Complex 11 Blue Knobby Adamellite 16 Kimberley Park Adamellite 34 Moona Plains Complex 12 Enmore Adamellite 23 Rockisle Granodiorite 35 Apsley River Complex 74

(i) Hillgrove Suite: The Hillgrove Suite as defined here, contains all of the 17 plutons included in the initial definition by Binns et al. (1967), with the exception of the Rockisle Granodiorite. The Rockisle Granodiorite is excluded on geochemical, mineralogical and isotopic grounds, and is placed in the Rockisle Suite (see below). In addition, six plutons have been added to the suite. The Glenifer Adamellite, Murder Dog Adamellite, Ingleba Adamellite, Campfire Adamellite (all added to the suite in Pogson & Hitchins (1973)), the Henry River Adamellite (included in the suite by Shaw & Flood (1981)) and the Gostwyck Adamellite (tentatively added to the suite by Gilligan et al. (1992)), are also included in the definition herein, bringing the total number of plutons in the suite to 22. These additions are largely the result of detailed mapping (e.g. Flood 1971, Leitch 1972), with most of these additional plutons cropping out at the spatial extremities of the suite. Both the Henry River Adamellite and the Murder Dog Adamellite, the most northerly and southerly plutons of the suite respectively, remain poorly mapped. The expanse of the Murder Dog Adamellite has been extended in an easterly direction by personal mapping (see Fig. 4A1), although the depiction of truncation by the Dingo Fault remains tentative. In addition, the aerial extent of the Henry River Adamellite has been reduced from that indicated in Shaw & Flood (1981), as much of the northeastern portion of this intrusive mass has been identified (personal mapping) as massive I-type granitoid, with characteristics of the Uralla Plutonic Suite. The Gostwyck Adamellite, originally included in the Uralla Plutonic suite (O’Neil et al. 1977), was placed in the ‘transitional’ group of plutons by Hensel (1982). However, due to the chemical, mineralogical and isotopic similarity of this pluton to other members of the Hillgrove Suite (as discussed above), it is incorporated here into the Hillgrove Suite, as suggested by Gilligan et al. (1992). In addition, the Gostwyck Adamellite is temporally related to the Hillgrove Suite (see Chapter 3) with the Rb-Sr biotite age of 290 Ma (Landenberger et al. 1995), rather than to the Uralla Suite which has biotite Rb-Sr ages ranging from 245-252 Ma (Shaw & Flood 1993).

Lithologically, the Hillgrove Suite is characterized by biotite adamellites and subordinate granodiorites, which occasionally bear trace almandine garnet (e.g. the Winterbourne Granodiorite and Murder Dog Adamellite). The primary distinction from the Rockisle Suite is the absence of amphibole. 75

(ii) Rockisle Suite: The distinctive feature of the Rockisle Suite is the presence of trace amphibole and the absence of garnet. Otherwise, members of the Rockisle suite share many of the characteristics of the Hillgrove Suite, such as the ubiquitous deformation features and the absence of magnetite. Most members of this suite are granodioritic in composition, although amphibole-bearing adamellites do occur (e.g. parts of the northern margin of the Hillgrove Adamellite). Amphibole bearing phases, which occupy minor parts of many Hillgrove Suite members, including the Dundurrabin, Hillgrove, Gara, Gostwyck, Kilburnie and Murder Dog plutons, are here placed in the Rockisle Suite. The abundance and precise extent of these phases within each pluton is undefined, and therefore, on a larger scale, these plutons must remain defined as Hillgrove Suite. Amphibole is ubiquitous throughout the Rockisle Granodiorite and Kimberley Park Adamellite, with all samples from plutons defining linear geochemical trends separate from the Hillgrove Suite, and hence the division of these granitoids into a separate suite is warranted. In addition, discrete granitoid bodies that occur within the mafic complexes also share the characteristics of these plutons. For example, low- K amphibole-bearing granodiorites occur in the Moona Plains, Woodburn, Dorrigo Mountain and Sheep Station Creek complexes. These granodiorites are in many respects similar to the Rockisle Granodiorite, and indeed the low-K composition of granodiorites from the Sheep Station Creek Complex (SSCC), led Paul (1984) to place these in the Clarence River Plutonic Suite. The Clarence River Plutonic Suite (Shaw & Flood 1981) has similar characteristics to the Nundle Suite, and is of similar age (250-255 Ma, S.E. Shaw pers. comm. 1995), and the placement of the granodiorites of the SSCC into the earliest Triassic Clarence River Plutonic Suite, is based on the similar reasoning as placement of the Rockisle Granodiorite into the Nundle Suite by Hensel et al. (1985). However, as discussed by Gilligan et al. (1992), these complexes are likely to have been coeval with the Hillgrove Suite, and are therefore unrelated to these younger granitoid suites.

Although the presence of amphibole is the distinguishing feature of the Rockisle Suite, mineralogy within the suite varies considerably. Two distinct types of granodiorite occur within the suite. The dominant type is distinguished by the presence of secondary actinolite, or actinolitic hornblende, which occurs as decussate aggregates, probably replacing primary clinopyroxene (as discussed below - section 4A3A3). This is the most common form of amphibole within the Hillgrove Supersuite granitoids, and is the sole mode of occurrence in 76 the Rockisle Granodiorite and Kimberly Park Adamellite. The second minor group is distinguished by the presence of primary (magmatic) hornblende, which is a common phase within the granodiorites of the mafic complexes. Adamellites from the western margin of the Kilburnie pluton also contain primary igneous hornblende, which occurs as overgrowths on the decussate aggregates of actinolite described above. The mineralogy of these two groups will be discussed in more detail below (section 4A3A3).

(iii) Bakers Creek Suite: A group of small intrusive complexes which are dominated by gabbros and diorites, comprise the Bakers Creek Suite. Although granitic phases do occur within these complexes, these have been assigned to the Rockisle Suite as discussed above. The 11 intrusive complexes of the Bakers Creek Suite (see Fig. 4A1) are spatially associated with the Hillgrove Suite, and although some form isolated plutons (e.g. Apsley River Complex), others are in direct contact with Hillgrove Suite members (e.g. the Mornington Diorite). In addition, these complexes are truncated by D7 mylonite zones, which are contiguous with those that cut Hillgrove Suite plutons (e.g. the Mornington Diorite, Farrell 1992). Most of these complexes also display the deformation style characteristic of the Hillgrove Suite granitoids. Those rocks which do not display mesoscopic deformation (i.e. foliation development), show abundant microscopic deformation features (e.g. deformation bands and lamellae in quartz, together with bent mica grains, and kinked and/or fractured feldspars). The only rock-types which appear to be free of deformation are the quartz-free gabbros, which have an exceptionally rigid matrix of interlocking (doleritic texture) plagioclase, and are unlikely to easily develop foliation or deformational microstructure (c.f. Vernon & Flood 1988).

Hensel (1982) regarded the mafic complexes as early Permian in age (and therefore younger than the Hillgrove Suite), and suggested that field relationships demonstrated that these mafic plutons intruded Hillgrove Suite granitoids. However, the field relationships of the mafic complexes of the Bakers Creek Suite are equivocal. In most cases, contacts are obscured, and megascopic outcrop patterns (the convexity of pluton margins) are also equivocal. The Mornington Diorite displays a convex relationship with the Abroi Granodiorite suggesting that its intrusion postdated the granitoid (Farrell 1992), yet in detail, the pluton margins show an intimate relationship, with intermingling of mafic and felsic magmas (Mason 1968, Farrell 1992). The Sheep Station Creek Complex, which contains both members of the Hillgrove 77

Supersuite in addition to Bakers Creek Suite gabbros, also shows equivocal field relationships (Paul 1984). However, the large-scale outcrop pattern (convexity of phase boundaries) suggest that the gabbroic phase of this complex is the oldest.

Since most rocktypes of these complexes are exceptionally low in K2O and Rb, and most do not contain zircon, dating has proved difficult. Rb-Sr whole rock isochrons from within some of these complexes, yielded poorly constrained ages with large errors (Hensel 1982). The only reliable ages obtained were Rb-Sr analyses from biotite - whole-rock pairs from biotite bearing diorites within these complexes. The ages obtained range from 2616283 Ma for the Moona Plains Complex, 2686277 Ma for the Mornington Diorite, and 284 Ma for the Bakers Creek Diorite. In addition a sample of the Bakers Creek Diorite was dated by K-Ar (biotite), yielding an age of 291 Ma (Pogson & Hilyard 1980). These ages bear a remarkable similarity to the range of biotite Rb-Sr biotite ages obtained from Hillgrove suite granitoids (Landenberger et al. 1995). Hence, the younger biotite Rb-Sr ages obtained from these mafic complexes, most likely represent cooling ages, as suggested for the Hillgrove granitoids (Landenberger et al. 1995). This explanation was also suggested by Gilligan et al. (1992). Furthermore, although Hensel (1982) suggested that these complexes were younger than the Hillgrove Suite, he conceded that the older mineral ages obtained (e.g. Bakers Creek Diorite) implied that a close temporal relationship probably existed between those complexes that are in direct contact with Hillgrove Suite plutons.

In summary, the available data suggest that intrusion of the mafic complexes of the Bakers Creek Suite are coeval with that of Hillgrove Supersuite plutons. As listed in table 4A2, the major rock types here included in the Bakers Creek Suite, are gabbros and diorites. Although most complexes of this suite are dominated by these mafic rocks, more felsic rocks include granophyre dykes and vein networks (Hensel 1982), as well as larger volumes of granitic material (hornblende-bearing granodiorites), which as stated previously are included in the Rockisle Suite. More detail on rocktypes will not be discussed here as these complexes have already been described in detail by Hensel (1982 - chapter 7), although relevant mineralogical detail will be discussed in section 4A3A9. 78

4A3 Petrography & Mineral Chemistry

4A3A1 General Petrography of Hillgrove Supersuite granitoids

Most granitoids of the Hillgrove Supersuite are adamellites and granodiorites, based on the classification of Streckheisen (1973) (Fig. 4A2a). Although the term ‘monzogranite’ is preferred to ‘adamellite’ in the Streckheisen classification, the latter term is favoured here, as it forms an integral part of Australian stratigraphic nomenclature, and is still used widely in Australia (note that the term ‘granite’ (sensu stricto) as used here, is the equivalent of the ‘syenogranite’ of Streckheisen (1973)). Members of the Hillgrove Supersuite and the Bundarra Suite exhibit a horizontal trend on the Streckheisen plot (Fig. 4A2a), resulting from the high modal quartz of both granodiorites and adamellites, a feature which is typical of S- type granitoids (e.g. Hine et al. 1978). The following petrographic descriptions concentrate on the occurrence of major mineral phases, the original igneous relationships between minerals (as preserved in weakly deformed samples), and the chemical mineralogy of Hillgrove Supersuite granitoids. Brief comparisons will be made with other granitoid suites of the New England Batholith and the Lachlan Fold Belt (LFB). The deformational textures developed within Hillgrove Supersuite granitoids have already been described in detail in chapter 2, and will not be discussed in detail here.

Hillgrove Supersuite granitoids are typically medium to fine grained and mildly porphyritic, being phenocrystic in plagioclase (and alkali feldspar in adamellite varieties). Fresh hand- specimens are normally pale bluish-grey to white, a colour which is interspersed with abundant reddish brown biotite. In addition to feldspars and quartz, the only other major phase occurring in the Hillgrove Supersuite granitoids is biotite, which forms up to 20% in the more mafic samples. Modal proportions of these major phases, with samples in order of increasing silica of the host, are summarized in figure 4A2b. Ilmenite is the sole opaque phase throughout the Hillgrove Supersuite, and other accessories in the granitoids include zircon, apatite, monazite and xenotime. Undeformed samples have a typical equigranular granitic texture with early crystallized euhedral plagioclase and biotite, joined by later quartz, alkali feldspar and minor phases, which form the interstices (Plate 4A1b). The distinctive mineralogies of the Hillgrove Suite, Rockisle Suite and Bundarra Suite are summarized in table 4A3 and comparisons are made with the classical S- and I-type granites of the LFB. 79 Quartz

Hillgrove Supersuite (a) Bundarra Suite

Adamellite Granodiorite

Alkali Feldspar Plagioclase 100 (b) Accessories 90 Biotite 80 70 Quartz 60

Mode 50 40 Alkali feldspar 30

20 Plagioclase 10 0 67 68 69 70 71 72 73 74 75 76 77

SiO2% Figure 4. 2. Modal plots for Hillgrove Supersuite Bundarra Suite granitoids: (a) Streckheisen plot and (b) Combined modal variation within the Hillgrove Suite over the observed silica range. 80

Hillgrove Rockisle Bundarra LFB S-type LFB I-type

Aluminous Biotite Aluminous Aluminous Biotite biotite ±actinolite biotite biotite ±hornblende ±almandine (after CPX) ±cordierite ±cordierite, ±CPX (rare) ±almandine ±almandine, ±ALS, ±muscovite ±OPX1 ±OPX1 ± OPX ±OPX ilmenite ilmenite ilmenite ilmenite magnetite

monazite monazite monazite ± monazite sphene xenotime xenotime xenotime ± xenotime ± allanite

Table 4A3 Distinctive mineralogy of the Hillgrove, Rockisle and Bundarra plutonic suites, as compared with LFB S- and I-type granites (Chappell & White 1972). ALS = aluminosilicate (andalusite, sillimanite). 1Trace OPX was noted in these suites by Hensel (1982).

Enclaves and other inclusions are relatively common in Hillgrove Supersuite granitoids. While microgranitoid enclaves are the most common form of enclave present, accidental xenoliths of country rock are not uncommon, and are particularly abundant in the deeper level granitoids such as the Tia Granodiorite (Dirks et al. 1993) and Abroi Granodiorite (Farrell 1992). These intrusions are enveloped by a migmatitic complexes, and contain significant inclusions of high grade metamorphic material which are identical to the enclosing metasediments. Refractory components of the surrounding wallrock are particularly abundant in the Tia Granodiorite, where large lumps of isolated vein quartz (up to 30 cm) are not uncommon. These inclusions are concentrated near the pluton margins, and their occurrence suggests limited in situ assimilation of country rock by the magma. However, these inclusions are rare outside the Wongwibinda and Tia complexes, and microgranitoid enclaves predominate elsewhere in the Hillgrove Supersuite. Like microgranitoid enclaves in most granitoids (Vernon 1990), these inclusions are generally more mafic and finer grained than their hosts, and most bear the same mineralogy. However, some enclave samples analysed contain primary pyroxene, and significantly both clinopyroxene and orthopyroxene are present (e.g. sample A186X from the Gara Adamellite) - a point which will be discussed below.

The most felsic rocks of the Hillgrove Suite are late-stage microgranite dykes, which occur sporadically throughout the suite. They are highly felsic (containing up to 78% SiO2) and are commonly accompanied by late pegmatitic phases containing tourmaline. 81

Plate 4A1. Photomicrographs of Hillgrove Supersuite granitoids and associated rocks.

(a) Typical ‘fertile’ greywacke (low-grade), dominated by volcanogenic lithic clasts and quartz, and minor detrital amphibole and clinopyroxene. Plane polarized light, field of view = 5 mm.

(b) Mildly deformed Hillgrove Suite Adamellite, with aluminous biotite, plagioclase, microcline and deformed quartz. Crossed polars, field of view = 5 mm.

(c) Garnetiferous Hillgrove Suite member showing corroded almandine garnet with reaction rims of low-Ti biotite (green) and muscovite. Plane polarized light, field of view = 5 mm.

(d) Amphibole-bearing Hillgrove Suite member, showing aggregates of fibrous actinolite, pseudomorphous after clinopyroxene. Crossed polars, field of view = 2 mm. 03

(a) (b) 14

13 82

4A3A2 Biotite

Aluminous biotite (α=pale straw, β=γ= ‘foxy’ red-brown) is generally the sole mafic silicate occurring in Hillgrove Suite granitoids. Biotite contents generally exceed 10% in all but late- stage felsic phases, with the most mafic granodiorites containing up to 20%, giving an average mode of 14% (see Fig. 4A2b, appendix C). In most specimens, biotite forms large (up to 5 mm) discrete flakes of subhedral character, and along with plagioclase, appears to be a near-liquidus phase. However, since biotite, together with quartz, are the first minerals to exhibit the effects of deformation, the original igneous relationships are obscured in most cases. Samples from the western margin of the Gara Adamellite (in which undulose extinction in quartz is the only deformation feature present) display perfectly euhedral prisms (principal section) of biotite ~4 mm in length (Plate 4A1b).

Inclusions of other major phases are rare in biotite flakes, but inclusions of accessory minerals are ubiquitous. Apatite, zircon, monazite and xenotime (in decreasing order of abundance), all occur as small, euhedral inclusions within biotite. Pleochroic haloes are well developed around all of these, excluding apatite. Ilmenite also occurs as inclusions in biotite forming somewhat larger crystals. However, these are generally not euhedral and frequently exhibit corrosion textures.

Although Hillgrove Suite samples show degrees of deformation varying up to mylonitic, and biotite flakes are bent and kinked in most specimens, the majority of biotite observed is fresh and unaltered. Intense deformation and annealing in high temperature mylonites produces foliated aggregates of recrystallized biotite of similar composition to the primary mica, together with minor muscovite/ilmenite intergrowths (Landenberger et al. 1995). Contact metamorphism by younger I- and A-type granitoids produces similar aggregates, which, in this case are decussate. Lower temperature mylonites contain neo-crystallized low-Ti green biotite, associated with exsolved ilmenite or rutile needles (Landenberger et al. 1995). Although Hensel et al. (1985) described the red-brown and green varieties of biotite as ‘coexisting’, the green biotite is unequivocally secondary, and has developed in response to deformation. Lower temperature alteration into chlorite is limited and incipient in most cases. However, a more common low temperature alteration product (which I have also observed in S-type granites of the Kosciusko Batholith) are small aggregates of 83

(trioctahedral) mixed-layer clays (illite-smectite interlayers - identified by microprobe), which are common even in ‘undeformed’ samples. These decussate aggregates usually occur as boudin-like shapes within biotite {001} cleavage planes, their growth appearing to cause opening of the cleavage planes, and lattice kinking. Occasional alteration (hydrothermal?) to epidote also apparent in a minority of samples, and only appears to occur in the least deformed granitoids (e.g. Gara Adamellite).

Compared to biotites from the Triassic I-type granitoids of the New England Batholith, biotites from the Hillgrove and Rockisle suites have a relatively restricted compositional range (see appendix D). Mg-numbers for most samples of the Hillgrove Suite (excluding samples from the western margin of the Kilburnie Adamellite) are restricted to the range

IV Mg30-40 and Al contents between 2A3 and 2A6 cations (Fig. 4A3a). By comparison, biotites from the Bundarra Suite are higher in [AlIV] and are generally lower in Mg-number. Biotites from the Rockisle Suite are distinguished from those of the Hillgrove Suite by being slightly more magnesian and less aluminous, reflecting the chemistry of the host rocks (see section 4A4). Although the fields for biotite from the Hillgrove and Rockisle suites overlap with the more felsic end-members of the younger I-type granitoids in terms of Mg-number and AlIV, they are better distinguished in terms of Ti and AlIV contents (Fig. 4A3b). Hillgrove Suite biotites show distinctly higher Ti contents (Fig. 4A3b) and are, overall, more aluminous than the younger I-type suites. Biotites from the Rockisle Suite, however, plot in the field of overlap between the I-types and Hillgrove Suite (Fig. 4A3b). Bundarra Suite biotites are also more distinguished in terms of AlVI vs Ti, since they have distinctly higher AlVI contents (figure 4A3b). Likewise, secondary green biotites from the Hillgrove Suite are also strongly distinguished on figure 4A3b, having both higher AlVI and much lower Ti contents. The degree of alumina saturation in biotites from the Bundarra, Hillgrove and Rockisle suites together with the presence or absence of other aluminous phases (cordierite/garnet), is controlled by the alumina saturation of the host rock. This is demonstrated by an igneous ACF plot of mineral and whole rock chemistry (see later, Fig. 4A9), where tie lines from plagioclase, projected through the whole-rock compositions for the respective suites, intersect the right-hand-side of the diagram at the points corresponding to the range of biotite compositions occurring in each suite. Whole-rock compositions for the Bundarra Suite which plot above the tie-line between plagioclase and the most aluminous biotite, correspond to those samples which bear cordierite. 84 3.0 Siderophyllite Suite: Hillgrove (a) Rockisle Bundarra Hillgrove 2° biotite I-types

2.5 IV Al

Kilburnie Adamellite (western margin)

Annite Mg# Phlogopite 2.0 0 20 40 60 80 100 0.7

0.6 (b) 0.5

0.4 Ti 0.3

0.2

0.1

0.0 0.0 0.2 0.4VI 0.6 0.8 1.0 Al Figure 4. 3. Biotite compositions for the Hillgrove and Rockisle (open diamonds) and Bundarra (squares) suites plotted in terms of (a) Mg# vs. AlIV atoms, and (b) Ti vs. AlVI atoms. Biotite compositions from the I- type granites of the Moonbi Plutonic Suite, the Barrington Tops Granodiorite (upublished data, Collins et al. 1994) and the Chaelundi Complex (Chapter 5) are plotted as a field ("I" fill pattern) for comparison. Secondary green biotites from the Hillgrove Suite (solid diamonds) are plotted for comparison. 85

4A3A3 Other ferromagnesian phases: Amphiboles & pyroxenes

Although orthopyroxene is a relatively common ferromagnesian phase in S-type granitoids (Clemens & Wall 1988), and is often partly or wholly replaced by secondary cummingtonite (or more commonly replaced by biotite), clinopyroxenes and calciferous amphiboles are by definition, absent (see table 4A1). Although Hensel (1982) recorded primary orthopyroxene in the Enmore Adamellite, primary orthopyroxene not observed in this study (except in microgranitoid enclaves). Decussate aggregates of cummingtonite secondary after orthopyroxene are more common. Secondary cummingtonite occurs more commonly in the granodiorites of the Hillgrove Suite, as noted by Shaw & Flood (1981).

Decussate aggregates of sub-aluminous calcic amphibole are more common than secondary Fe-Mg amphiboles, and have been previously noted by both Hensel (1982) and Shaw & Flood (1981). Although the occurrence of these calcic amphiboles in the Hillgrove Suite was noted by these authors as unusual for an S-type granitoid suite, this problem has been resolved here by placing these calcic amphibole bearing granitoids into the Rockisle Suite, which is regarded here as transitional in character. These aggregates are usually <1 mm in diameter (see Plate 4A1d), often have pyroxene-like outlines, and never exceed 0A5% of the mode. They are composed of a pale green fibrous amphibole (α=colourless, β=pale brown γ=pale green), which classify as actinolite or actinolitic hornblende (Fig. 4A4a) on the classification diagrams of Hawthorne (1983). In addition to their occurrence in the Rockisle Suite, these actinolite aggregates are also common in microgranitoid enclaves of both the Hillgrove and Rockisle suites.

Actinolites of the Rockisle Suite are Mg-rich, with Mg-numbers ranging from 55-65 (Fig. 4A4a), which contrasts with biotites of Hillgrove Supersuite, which have Mg-numbers in the range 30-45. This chemical disequilibrium, together with the secondary nature of these amphiboles, implies that they are not cognate with the host granitoids, and their presence requires explanation. Flood & Shaw (1977) and Hensel (1982) suggested that they originated as disaggregated remnants of calcareous nodules, similar to those occurring as lenses in the surrounding sediments. However, a simpler explanation for the origin of these aggregates, which is evaluated below, is by the alteration (uralitization) of primary igneous clinopyroxene. 86

(a) ANa+AK<0.5 Ti<0.5 1 Tremolite Tr Hb

Tsch Actinolite Act Magnesio-Hbl Tschermakite Hbl Hbl

Mg/(Mg+Fe) Fe- Fe- Ferro- Ferro- Act Ferro-Hbl Tsch Actinolite Tschermakite Hbl Hbl

0 8.0 7.5 7.0 6.5 6.0 5.5 TSi Wo (b)

En Fs Figure 4. 4. (a) Classification of secondary amphibole compositions from the Rockisle Suite (hollow diamonds) and primary hornblende (filled diamonds) from the Kilburnie Adamellite (Hawthorne 1984), Secondary actinolites from the Bakers Creek Suite are plotted as hollow circles (Landenberger 1988) and the green shaded area (Hensel 1982). (b) Amphiboles plotted on the pyroxene quadrilateral together with secondary amphiboles (hollow circles and green shaded area) and primary pyroxene compositions from the Bakers Creek Suite (red circles and yellow shaded areas). Data taken from Hensel (1982) and Landenberger (1988). 87

Although primary augite does not occur in the granitoids of the Hillgrove and Rockisle suites, the morphology of these aggregates often mimics that of pyroxene. In addition, both primary augite and hypersthene occur in the microgranitoid enclaves of both suites, where both minerals are partly altered to secondary fibrous aggregates of amphibole (actinolite and cummingtonite respectively). Furthermore, decussate aggregates of secondary amphibole occur in large quantities as partial alteration (uralitization) products of augite and hypersthene in the gabbros and diorites of the Bakers Creek Suite. The secondary actinolites occurring in these mafic rocks, have compositions identical to the actinolite aggregates of the Rockisle Suite (Fig. 4A4a,b). Hence, the available evidence suggests that the fibrous secondary amphiboles of the Rockisle Suite originated as primary igneous augite derived by contamination of these granitoids from a more mafic magma, a point which will be discussed in more detail later (see section 4A5A5).

4A3A4 Peraluminous phases other than biotite: garnet, cordierite, muscovite.

Almandine garnet is an uncommon accessory in Hillgrove Suite granitoids, and is even rarer in the Bundarra Suite, where cordierite is a more common aluminous phase. Garnet is absent in granitoids the Rockisle Suite. In the Hillgrove Suite, almandine appears to be an early- formed phase and is always strongly corroded, displaying disequilibrium reaction coronas of green biotite (see plate 4A1c). The sample analysed from the Winterbourne Granodiorite has a uniform composition across the grain. The garnet present in Hillgrove Suite granitoids is always more Fe and Mn-rich (>74% almandine & up to 9% spessartine - see appendix D) than biotites from the same sample (Fig. 4A5). Hensel (1982) described three types of garnet occurring within the Hillgrove Suite, all of which displayed moderate core6rim zonations, primarily showing increases in Mn contents, although grossular-rich varieties exhibit a reverse zonation. Based on the evidence available, Hensel (1982) concluded that the grossular-rich varieties (type ‘C’) were derived from accidental calcareous xenoliths, while the other two groups (types ‘A’ and ‘B’) originated as restite and magmatic crystallization products, respectively.

In contrast to the variety of garnet types occurring in the Hillgrove Suite, garnet from the Bundarra Suite is relatively restricted in type and occurrence, and has a zonation pattern (appendix D, Fig. 4A5) that is consistent with crystallization from a granitic melt (Flood et al. 88

1991). Cordierite from the Bundarra Suite also display chemical zonation (appendix D, Fig. 4A5) that is consistent with magmatic crystallization (Flood et al. 1991). However, the compositions of biotite, garnet and cordierite from these granitoids cannot be used to estimate P/T conditions of crystallization, as both garnet and cordierite appear to be early crystallizing phases, and both exhibit disequilibrium reaction coronas consisting of green biotite ± muscovite, and hence are not in equilibrium with the primary magmatic biotite.

Muscovite does not occur as a primary magmatic phase within granitoids of the Hillgrove Suite. This is primarily due to moderately peraluminous magma compositions, with most Hillgrove Suite granitoids plotting below the aluminous biotite - plagioclase tie line on the igneous ACF diagram (see later, Fig. 4A9). Hence, biotite is the only peraluminous phase required to account for the alumina oversaturation in these granitoids. Additionally, the >3 kb pressures, low temperatures and high f conditions required for the crystallization of H2O primary magmatic muscovite (Clemens & Wall 1988), were highly unlikely in the case of the Hillgrove Suite, with emplacement levels of <3 kb inferred for the majority of plutons (Landenberger et al. 1995). However, secondary muscovite, together with ilmenite, form as subsolidus reaction products of biotite, which is common in the deeper-level plutons of the Hillgrove Suite. It generally occurs as foliated aggregates formed by dynamic recrystallization during deformation. Secondary, very fine-grained white mica (‘sericite’) is also a common alteration product occurring in the cores of plagioclase crystals and in alkali feldspars, a feature common to most granitoids.

The coarser grained muscovite (#2 mm), which occurs in granitoids of the Bundarra Suite has been described as ‘primary’ (Flood & Shaw 1975). However, the P/T requirements for the crystallization of primary muscovite (>3 kb) is at odds with the P/T conditions required for the crystallization of cordierite (<3 kb, Clemens & Wall 1988), a phase common in the Bundarra Suite. It is more likely that this muscovite is a high temperature (subsolidus) reaction product of cordierite, since it often occurs as aggregates rimming corroded cores of cordierite. The problem of a primary versus secondary origin for muscovite in granitic rocks is discussed in detail by Clemens & Wall (1988). These authors concluded that the only genuine primary magmatic muscovite is that occurring in the two-mica granites of high grade regional metamorphic belts - the regional-aureole granites. The only example of this mode of occurrence in the SNEFB is the two-mica granites of the Wongwibinda Complex, which 89

Garnet Ilmenite Hillgrove Bundarra

Mg Fe2+ Cordierite (Bundarra Suite) Biotite

70 60 50 40 30 20 10 XMg Figure 4. 5. Mg-Fe2+ -Mn molecular projection of biotite, ilmenite, garnet and cordierite from the Hillgrove and Bundarra suites. Ilmenite and biotite fields emcompass analyses from both suites

Hillgrove Rockisle

Andesine Bundarra

Oligoclase

Orthoclase + Microcline

Ab Or Figure 4. 6. An-Ab-Or ternary plot of feldspar compositions from the Hillgrove Supersuite and Bundarra Suite 90 are probably products of low degrees of partial melting of metasediments (at upper amphibolite facies conditions), forming either by muscovite breakdown or under fluid- present conditions (Farrell 1992). The two-mica granites of the Wongwibinda Complex, crystallized under high f conditions at ~3 kb (Farrell 1992), allowing the crystallization H2O of primary magmatic muscovite.

4A3A5 Oxides and other opaque phases

Ilmenite is the sole oxide phase occurring in the plutonic rocks of the Hillgrove Supersuite, and the gabbros and diorites of the Bakers Creek Suite. Within the granitoids, ilmenite forms small subhedral prisms up to 0A5 mm in length, occurring almost exclusively as inclusions within biotite flakes, although sometimes as inclusions in plagioclase crystals. Modal abundances are generally low, not exceeding 0A25% in most cases (appendix C). The mode of occurrence is similar in the Rockisle and Bundarra Suites, and in biotite bearing diorites of the Bakers Creek Suite. However, ilmenite in these diorites tends to show extensive resorption. In gabbroids of the Bakers Creek Suite, ilmenite occurs almost exclusively interstitially.

Ilmenite from the granitoids of the Hillgrove Supersuite shows little variation in chemistry apart from varying replacement of Fe by Mn, with contents varying from ~5611 wt% MnO (Fig. 4A5, appendix D). Mn contents of ilmenite from the Bundarra, Hillgrove and Rockisle suites show no systematic variation. However, samples from the western margin of the Kilburnie Adamellite (Rockisle Suite), have lower contents of ~3 wt% MnO (appendix D, Fig. 4A5). Ilmenite from diorites and gabbros of the Bakers Creek Suite have similarly low MnO contents.

Other opaque phases occurring within Hillgrove Supersuite granitoids are occasional sulphides (generally pyrrhotite) and ‘ubiquitous’ graphite (Hensel et al. 1982), although the latter phase has not been recognized in the current study. 91

4A3A6 Plagioclase

Plagioclase generally occurs as euhedral prisms up to 10 mm in length in all Hillgrove Supersuite granitoids, with modal variations of 30-36% (Fig. 4A2b). Plagioclase phenocrysts are isolated and not interlocking, and together with biotite phenocrysts, form a granitic texture, with other mineral phases occurring interstitially (Plate 4A1b). Therefore, it is considered to be a primary liquidus phase. However, unlike biotite, plagioclase phenocrysts retain their euhedral appearance with the onset of deformation, only bending and/or fracturing with mylonitization. Inclusions of other major phases in plagioclase are rare, with the exception of biotite flakes which may be partly enclosed in the rims of phenocrysts. Inclusions of accessory phases in plagioclase are more common, but less so than in biotite. Apatite is the most abundant accessory in plagioclase, along with minor zircon and rare-earth phosphates, however in contrast to biotite, inclusions of ilmenite are rare. Alteration of plagioclase is limited to sericite in most samples, with the most extensive alteration occurring in the cores of phenocrysts, although some samples have prehnite as an additional alteration product (also more predominant in cores).

Plagioclase zoning in Hillgrove Suite granitoids is generally normal, with oscillatory zoning uncommon, and where present, weakly defined. Large cores of uniform composition, mottled cores, and compositional spikes are also rare. In contrast, plagioclase phenocrysts from the Rockisle Suite, commonly have rounded cores of uniform composition which show a weak mottling, and commonly show a greater degree of sericitic alteration than the remainder of the crystals. These cores are usually overgrown by normally zoned plagioclase, with oscillatory zoning observed in a minority of samples. These rounded cores, together with the zoned plagioclase overgrowths, are the equivalent of the ‘two generations’ of plagioclase described by Hensel et al. (1985).

Plagioclase compositions are relatively uniform across the Hillgrove Suite. In most samples, plagioclase rim compositions terminate at ~An20, and only rarely extend down to ~An12. Plagioclase core compositions (appendix D) are also remarkably consistent throughout the suite, restricted to the range An33638, even in the more felsic leucogranite dykes (Fig. 4A6). This range of core compositions is also common in the Bundarra Suite, although some felsic variants of this suite have plagioclase cores not exceeding An25. The zoned portions of 92 plagioclase from the Rockisle Suite also conform to the restricted range exhibited by Hillgrove Suite plagioclases. However, the distinct mottled, uniform plagioclase cores peculiar to this suite, have compositions of An40650, distinctly more calcic than the zoned portions of plagioclase from either the Rockisle Suite or Hillgrove Suite (Fig. 4A6). The degree of cloudy alteration and sericitization present in these cores (as well as secondary prehnite), also suggest that they were originally more calcic than their present composition. Plagioclase cores in the hornblende-bearing western portion of the Kilburnie Adamellite are also rather calcic, with compositions of An50, which is comparable to plagioclase compositions from the Bakers Creek Suite.

4A3A7 Alkali Feldspar

Unlike plagioclase, alkali feldspar is not an early crystallizing phase in Hillgrove Supersuite granitoids, and occurs either interstitially or as poikilitic phenocrysts. Total alkali feldspar contents for the Hillgrove Suite vary from 12% to 22% (Fig. 4A2b), whereas Rockisle Suite granitoids are restricted to <16%. Phenocrysts only occur in those samples with totals of >18% alkali feldspar, and their presence distinguishes adamellites from granodiorites. Phenocrysts of alkali feldspar are commonly poikilitic, enclosing early formed plagioclase and biotite phenocrysts, and occasionally quartz. Interstitial alkali feldspar and phenocryst margins are commonly associated with quartz in a micrographic intergrowth. Alkali feldspar in fresh Hillgrove Supersuite granitoids is either white or bluish-grey in colour, a feature typical of S-type granites (White & Chappell 1983), reflecting the lower oxidation states of S-type magmas relative to I-types. The colour of alkali feldspars of the Hillgrove Supersuite also appears to relate to the degree of deformation, with the orthoclase-rich varieties of the least-deformed samples being white, while microcline development is accompanied by a change to a bluish-grey colour in hand specimen.

Perthitic texture is well developed in most specimens except microgranite dykes. Exsolution patterns are coarser (i.e. lamellae are both larger and more widely spaced) in phenocrysts than in interstitial feldspar. The total proportion of exsolved feldspar in phenocrysts is also greater, particularly in cores, presumably reflecting an primary igneous zoning of sodic cores with more potassic rims. In contrast to alkali-feldspar from some I-type granitoids (e.g. that of the Chaelundi Complex - see chapter 5), in which exsolution lamellae are oligoclase in 93 composition, those of the Hillgrove Suite are exclusively albite (not exceeding An3). The proportion of exsolved albite in individual samples also appears to relate to the structural state of the feldspar; microcline has a far greater proportion of exsolved albite than the less structured orthoclase. These microcline microperthites commonly show two stages of perthitic exsolution, with an early coarse set of exsolution lamellae, which is overprinted by a second set of finer lamellae at a high angle to the former. Additionally, the composition of the host alkali feldspar in microcline samples is more potassic than their orthoclase counterparts.

The structural state of alkali feldspars in the Hillgrove Supersuite varies widely, with the least deformed specimens, such as those from the western margin of the Gara Adamellite, having simply twinned orthoclase. In contrast, strongly deformed samples are dominated by tartan- twinned microcline. The relationship of triclinicity to deformation was demonstrated by Flood (1971), who showed that with increasing proximity to the Kilburnie Fault, triclinicity of alkali feldspar in the Kilburnie Adamellite systematically increased from 0A5660A93 as the fault was approached.

The varying original composition of magmatic orthoclase (phenocryst cores vs. rims and interstitial feldspar), together with the variable structural state, has resulted in a wide range in alkali feldspar compositions for the Hillgrove Supersuite. Orthoclase-rich varieties vary in composition form Or55 to Or80, while host microcline in strongly deformed samples is generally >Or90 in composition (Fig. 4A6). Bundarra Suite samples (all of which are phenocrysts) show a similar spread, but none are more sodic than Or70, with microcline samples again exceeding Or90 (Fig. 4A6).

4A3A8 Quartz

All Hillgrove Suite granitoids are quartz-rich (a typical feature of S-types), with contents always exceeding 30% and approaching 36% in the most felsic variants (Fig. 4A2b). It occurs both as subhedral phenocrysts up to 5 mm in diameter and as finer grained interstitial aggregates. Phenocrysts also form aggregates in the most felsic variants, exhibiting a glomeroporphyritic texture. With the onset of deformation, these aggregates are strung out into elongate quartz ribbons (see chapter 2). Although quartz forms phenocrysts in most 94 samples, it is unlikely to be an early liquidus phase, as it rarely occurs as inclusions in plagioclase or biotite, but does itself poikilitically enclose these liquidus phases.

Microscopically, quartz shows ubiquitous deformation features, varying from undulose extinction in the least deformed samples, to deformation bands, deformation lamellae, subgrain development, and finally recrystallization in mylonitic samples (see chapter 2 for details). ‘Blue’ quartz in Hillgrove Supersuite granitoids has been described by many authors (Binns 1966, Hensel 1982, Shaw & Flood 1981). Although the hand-specimen colour of the quartz is typically an opalescent blue, this varies from samples in which all quartz is colourless to those in which all quartz is blue, with both varieties frequently occurring in the same specimen, as noted by Hensel (1982).

Various explanations have been suggested for the differing colours of quartz. Binns (1966) and Flood (1971) suggested that the colour was due to inclusions of rutile, which is a commonly suggested cause of bluish quartz. Hensel (1982), suggested that the blue quartz represented refractory restite entrained from the source region, while the colourless quartz was that crystallized from the magma. However, quartz is unlikely to occur in large quantities as a refractory phase of partial melting, and inclusions of rutile have not been observed in this study.

It is suggested that the opalescence of quartz in Hillgrove Supersuite granitoids directly relates to the structural dislocations caused by deformation. The degree of opalescence and the intensity of the blue-grey colour, has a direct relationship with the degree of deformation. Samples in which quartz only exhibits undulose extinction, have colourless quartz in hand specimen. With increasing deformation, the formation of deformation bands, deformation lamellae, and subgrains, the opalescence of quartz in hand specimen increases systematically. However, this opalescence is in turn replaced by colourless quartz in intensely mylonitized samples, in which complete dynamic or static recrystallization has occurred. Those samples in which recrystallized quartz coexists with highly strained magmatic quartz, usually along dislocation boundaries, are those samples in which both colourless and opalescent quartz is observed in hand specimen. So the occurrence of opalescent blue quartz, at least in this case, appears to directly relate to the density of crystal lattice distortions. 95

4A3A9 General petrography of the Bakers Creek Suite

The Bakers Creek Suite is composed of a range of rock types from gabbros to quartz diorites. Hensel (1982) subdivided these mafic complexes into ‘tholeiitic’ (most of the gabbroic rocks) and ‘calc-alkaline’ (some gabbros + dioritic rocks). Some gabbros were erroneously designated as ‘calc-alkaline’ in this subdivision, based on the association with more felsic diorites and the degree of subsolidus alteration (e.g. uralitization), which masks the original mineral assemblage of some gabbros. A simple division into gabbros and diorites will also be applied here, with the exclusion of the terms tholeiitic and calc-alkaline. The following descriptions are generalized and brief, and the reader is referred to Hensel (1982, chapter 7) for more details of the petrography and chemical mineralogy.

(I) Gabbros: The gabbros of the Bakers Creek Suite are plagioclase, two-pyroxene ±olivine gabbros. Olivine is present in most gabbros, but is often pseudomorphed by low temperature minerals in altered specimens, or may be completely absent in fractionated varieties. Calcic plagioclase (up to An80) dominates these gabbros and aside from olivine, was the earliest crystallizing phase. Plagioclase is euhedral and interlocking, and imparts a typical ophitic texture to the rocks. Pyroxenes are interstitial to plagioclase in most specimens, with augite dominant over hypersthene in most gabbros, particularly the more mafic types. Additionally, exsolved pigeonite was noted by Hensel (1982) in some samples from the Days Creek Gabbro. An important feature of these gabbroic rocks is the occurrence of ilmenite as the sole oxide phase, although rare magnetite was noted by Hensel (1982). Additional accessory phases include ubiquitous apatite, together with minor quantities of interstitial phlogopite and brown magnesian hornblende in some specimens. Subsolidus hydrothermal alteration is extensive in some examples, indicated by serpentinization of olivines and uralitization of pyroxenes.

(ii) Diorites and quartz diorites: Like the gabbros, the Bakers Creek Suite diorites are also dominated by plagioclase, and characteristically have ophitic textures. The plagioclase in these diorites is not as calcic as that in the gabbros, and often shows complicated zoning. The ferromagnesian minerals present are dominantly hydrous, with both primary hornblende and biotite present in most 96 specimens. Coexisting with these phases are decussate aggregates of actinolite and cummingtonite occurring as the alteration products of primary augite and hypersthene respectively. Unaltered pyroxenes are absent in some rocks, but frequently occur as relict cores within reaction coronas. Quartz is a common interstitial phase, comprising up to 10% of the more felsic varieties. As in the gabbros, ilmenite is usually the only oxide phase present, but sphene and zircon appear as accessories in addition to apatite.

4A4 Geochemistry

4A4A1 Defining the geochemical characteristics - variations within the Hillgrove Supersuite, and comparisons with other granitoids.

The majority of Hillgrove Suite granitoids sampled are restricted to a silica range of 68672%

SiO2, while the relatively few samples with <68% SiO2, and those that are strongly differentiated, extend the range to 67677% SiO2. Within this range, the increase in silica is accompanied by an increase in K2O, Rb, Pb, Th, Y, Nb, and a decrease in all other elements except Na2O, which remains relatively constant. Some trace elements such as Y and Nb are highly variable (Figs. 4A7a,b, 4A8). Major element ratios also vary, with Mg# (molecular

Mg/(Mg+Fe) or XMg) and An# (molecular Ca/(Ca+Na) or XCa) decreasing, Or# (molecular

K/(K+Na) or XK) increasing, and ASI (alumina saturation index) decreasing slightly.

Averages for ‘mafic’ (Rb<150 ppm) and ‘felsic’ members (190

Legend

Hillgrove Supersuite (all samples - this study)

+ Accretion complex metasediments (this study)

Bundarra Suite (all samples - this study)

Hillgrove Suite (excluding Rockisle Suite - uncontaminated)

Metasediments (Tablelands Complex)

Average fertile (intermediate) greywacke

Bundarra Suite

Accretion complex metabasalts (circles represent Tia complex metabasalts - this study)

Hillgrove Suite microgranitoid enclaves.

Diorites Bakers Creek Suite Gabbros}

Main Hillgrove Supersuite contamination trend(s) 1.0 8 ao lmn akrposfrSO s i lO, E O , Al vs. TiO 7a. Major element Harker plots for SiO Figure 4

TiO2

2 0.8

. 6

1 0.6 4

0.4 EFeO 10 EFeO TiO 2 0.2 2 5

50 60 70 60 65 70 75 80 60 65 70 75 80

SiO2 SiO2 3 2

Al2O3 17 20 Al2 O 3 3 2 2 MgO 16 15 2

FeO, and MgO. 15

MgO 14 10 1

13 5 98 60 65 70 75 80 50 60 70 60 65 70 75 80

SiO2 SiO2 99 6 5 4 3 2 1 0 0.2 0.1 0.0 2 5 PO 2 2 SiO SiO 2 KO 60 65 70 75 80 60 65 70 75 80 5 O 2 P CaO O 2 50 60 70 50 60 70 K O 2 Na 5 4 3 2 1 5 4 3 2 5 10 0.4 0.3 0.2 0.1 2 CaO Na O 2 2 SiO SiO 60 65 70 75 80 60 65 70 75 80 6 5 4 3 2 1 0 0 1 2 3 4 . Figure 4 7b. Major element Harker plots for SiO2 vs. CaO, Na2 O, K 2 O, and P2 O 5 . 100 0 80 60 40 20 1.6 1.4 1.2 1.0 0.8 An# 2 2 SiO SiO ASI corundum 1% normative 60 65 70 75 80 60 65 70 75 80 Mg# ASI 50 60 70 50 60 70 Or# An# 80 60 40 20 1.4 1.2 1.0 0.8 0.6 60 40 20 80 60 40 20 2 2 SiO SiO Mg# Or# 60 65 70 75 80 60 65 70 75 80 0 0 60 50 40 30 20 10 20 40 60 80 . Figure 4 7c. Major element Harker plots for SiO2 vs. Mg#, An#, Or#, and ASI. 101 40 30 20 10 0 0 600 500 400 300 200 100 Sr 2 2 SiO SiO Pb 60 65 70 75 80 60 65 70 75 80 Sr Ba Rb 50 60 70 50 60 70 Pb 300 200 100 500 400 300 200 100 800 600 400 200 30 20 10 2 2 SiO SiO Rb Ba 60 65 70 75 80 60 65 70 75 80 0 0 100 200 300 200 400 600 800 1000 . Figure 4 8a. Trace element Harker plots for SiO2 vs. Ba, Sr, Rb, and Pb. 102 5 0 20 15 10 0 60 50 40 30 20 10 2 2 SiO SiO Y Nb 60 65 70 75 80 60 65 70 75 80 fractionated Hillgrove Nb 50 60 70 Th 50 60 70 Zr Y 20 10 50 40 30 20 10 400 300 200 100 30 20 10 2 2 SiO SiO Th Zr 60 65 70 75 80 60 65 70 75 80 0 0 10 20 30 100 200 300 400 . Figure 4 8b. Trace element Harker plots for SiO2 vs. Zr, Y, Th, and Nb. 103 0 80 60 40 20 100 60 50 40 30 20 10 0 Zn Cr 2 2 SiO SiO 60 65 70 75 80 60 65 70 75 80 Zn Cr V 50 60 70 50 60 70 Cl 500 400 300 200 100 300 200 100 300 200 100 50 100 V 2 2 SiO SiO Cl 60 65 70 75 80 60 65 70 75 80 0 0 50 50 100 150 100 150 200 . Figure 4 8c. Trace element Harker plots for SiO2 vs. V, Cr, Cl, and Zn. 104 Ca

Grey- wackes

Quartz-rich psammites

Pelites

Na K A l - N a - K Al-Na-K

Igneous ACF (atomic)

Plagioclase Cordierite

Garnet

Biotite

CPX (Bakers Creek Actinolite (Rockisle Suite gabbro) Suite - after CPX) Ca Fe+Mg . FigureC 4 a 9. Molecular Ca-Na-K and atomic igneous ACF plots of the HillgroveF e + MSupersuite, g Bundarra Suite and accretion complex lithologies. Ferromagnesian mineral assemblages from granitoids and gabbros are also included on the ACF plot. 105 1.2 1.0 0.8 0.6 0.4 0.2 0.0 0 50 200 150 100 2 2 SiO SiO K/Ba Rb/Ba 60 65 70 75 80 60 65 70 75 80 K/Rb Ca/Sr K/Ba Rb/Ba 50 60 70 50 60 70 1.0 0.8 0.6 0.4 0.2 600 400 200 50 200 150 100 600 400 200 K/Rb 2 2 SiO SiO Ca/Sr 60 65 70 75 80 60 65 70 75 80 0 0 50 100 200 300 400 100 150 200 . Figure 4 10a. Inter-element ratio Harker plots for SiO2 vs. Ca/Sr, K/Ba, K/Rb, and Rb/Ba. 106 3 2 1 0 80 60 40 20 100 2 2 SiO SiO Fe/Mn 10000Ga/Al 60 65 70 75 80 60 65 70 75 80 Mg/V Ga/Al 50 60 70 Rb/Sr 50 60 70 Fe/Mn 4 3 2 1 80 60 40 20 2.5 2.0 1.5 300 200 100 2 2 SiO SiO Rb/Sr Mg/V 60 65 70 75 80 60 65 70 75 80 0 0 1 2 3 4 5 100 200 300 . Figure 4 10b. Inter-element ratio Harker plots for SiO2 vs. Rb/Sr, Ga/Al, Mg/V, and Fe/Mn. 107

Major LFB LFB Bullen- Strath- ‘Mafic’ ‘Felsic’ Rockisle NEB Bundarra Elements I-type S-type balong bogie Hillgrove Hillgrove Suite I-types Suite n 1074 704 86 25 1027712

SiO2 69A50 70A91 69A14 72A71 69A11 70A83 69A13 68A51 72A81

TiO2 0A41 0A44 0A55 0A34 0A63 0A52 0A49 0A45 0A31

Al2O3 14A21 14A00 14A26 13A62 14A35 14A03 15A04 14A61 13A93

Fe2O3 1A01 0A52 0A69 0A13 0A53 0A31 0A48 1A18 0A20 FeO 2A22 2A59 3A31 2A28 3A36 2A64 2A55 2A23 1A78 MnO 0A07 0A06 0A06 0A05 0A11 0A06 0A07 0A09 0A05 MgO 1A38 1A24 1A87 0A71 1A22 1A01 1A26 1A71 0A45 CaO 3A07 1A88 2A32 1A34 2A31 2A12 2A80 3A16 1A59

Na2O 3A16 2A51 2A19 2A69 3A28 2A96 3A84 3A61 3A15

K2O 3A48 4A09 3A69 4A42 3A65 4A40 3A02 3A51 4A40

P2O5 0A11 0A15 0A14 0A15 0A14 0A10 0A12 0A12 0A13 C -2A44 2A85 2A38 1A16 0A80 0A66 - 1A43 Di 0A17------1A32 - ASI 1A00 1A21 1A24 1A21 1A09 1A06 1A04 0A96 1A11 Fe3+/3Fe 0A29 0A15 0A16 0A05 0A12 0A10 0A14 0A32 0A09 Trace elements (ppm) Ba 519 440 480 605 505 591 474 - 492 Rb 164 245 181 236 68 198 114 - 203 Sr 235 112 128 88 468 124 251 - 120 Pb 19 27 28 24 14 26 16 - 22 Th 20 19 18 12 8 22 14 - 19 Zr 150 157 172 144 176 214 186 - 173 Y 31 32 32 31 22 38 25 - 36 La 31 27 31 18 26 23 26 - 23 V 57 49 72 28 91 42 47 - 21 Cr 20 30 46 17 23 22 20 - 12 Ni 811177 3 - 3 - - Zn 48 59 - - 67 48 51 - 39 Table 4A4. Comparative major and trace element chemistry for LFB I-type, LFB S-type and New England Batholith I-type granites, with the Hillgrove, Rockisle and Bundarra Suite granites. Average analyses of LFB I-type granites (LFB I) and S-type granites (LFB S) are from Chappell & White (1992a), averages for Bullenbalong Suite (S-type) and Strathbogie Suite (S-type) are from White & Chappell (1988), and average New England I-type is from Shaw & Flood (1981) (average of granitoids with>63% SiO2). ‘C’ and ‘Di’ are the % CIPW normative corundum and diopside respectively, ASI is the Alumina Saturation Index (Zen 1986a), and Fe3+/3Fe is the molecular proportion of Fe3+ in total Fe. For the Hillgrove Suite averages, ‘Mafic’ refers to samples with <150 ppm Rb, and ‘Felsic’ refers to samples with Rb contents between 190 and 200. boundary (normative C=0A8). However, this lower value is a common effect of differentiation in S-type granitoids (Chappell & White 1992a).

All sampled Rockisle Suite granitoids fall into the range 67671% SiO2, and for similar silica values, have lower TiO2, 3FeO, K2O, Ba, Sr, V, Zn and higher Al2O3, CaO, Na2O, Rb, than the Hillgrove Suite. Higher CaO and Na2O contents lower the ASI considerably, and hence normative corundum values for the suite (0A66) are below the division between S- and I- types. This is consistent with the amphibole-bearing character of the suite and with earlier 108 suggestions that members of this suite are transitional between I- and S-types (Shaw & Flood 1981) .

In comparison to the Hillgrove and Rockisle suites, sampled members of the Bundarra Suite are generally more felsic and more strongly peraluminous, with normative corundum contents of 1A43. This value falls well within the range for S-types, even though this suite is considerably higher in silica than granitoids of the Hillgrove Supersuite. However, a comparison with averages for the Strathbogie Suite of the LFB (table 4A4) shows that for similar silica values, the Bundarra Suite is (like the Hillgrove Suite) less peraluminous than its LFB counterpart. The lower ASI for the Bundarra Suite is (c.f. Strathbogie Suite) primarily a result of higher Na2O contents, although CaO contents of the Bundarra Suite also exceed those of the Strathbogie Suite. At similar silica contents, Bundarra Suite granitoids are similar to the Hillgrove Suite granitoids in most respects (Figs. 4A7,4A8) except for slightly higher Al2O3, P2O5, Nb, and a marked difference in the behaviour of some elements with increasing silica. For example, both K2O and Pb decrease with increasing silica, which is the opposite to the trend exhibited by the Hillgrove Suite, an effect which is probably produced by alkali feldspar fractionation (see section 4A5A1).

Rare-earth-element (REE) contents for the Hillgrove Suite granitoids are generally low, not exceeding 100×chondrite. Although relatively few samples were analysed, there is remarkable consistency across the suite (Fig. 4A11), with La/Lun values restricted to the range 4A95-5A84. All samples have moderate negative europium anomalies, with Eu/Eu* ranging from 0A33 to 0A53. In contrast the only sample analysed from the Rockisle Suite has higher overall REE contents (La .150×chondrite), and a larger negative Eu anomaly (Eu/Eu* =

0A21), but similar La/Lun( 5A36). The more felsic Bundarra Suite has lower total REE contents than either the Hillgrove or Rockisle suites, with values less than 70×chondrite.

La/Lun values are also lower (3A24-5A64), but negative Eu anomalies are larger (Eu/Eu* 0A21- 0A34). Lower total REE contents for these more felsic granitoids are unusual, as differentiation (which is indicated by the larger negative Eu anomaly) is normally accompanied by an increase in total REE levels. This paradox is strong evidence that REE contents in these S-type suites are largely controlled by the REE accessory phases monazite and xenotime, resulting in lower REE levels in the daughter products of any differentiation process. 109 Yb Hillgrove Suite Bundarra Suite Ho Lu Tb Eu Sm Nd Ce G37 La BI2 N45 A180 CB2 A284

10 100 Chondrite / Rock . 4 (dashed Suite granitoids and Bundarra (solid lines), Suite granitoids 11. for Hillgrove REE plot (Nakamura 1977) Chondrite-normalized lines). The single dotted line represents a primary hornblende bearing adamellite from the western margin of the Kilburnie Adamellite. Kilburnie the of margin western the from adamellite Figure bearing hornblende primary a represents line dotted single The lines). 110

4A4A2 The Bakers Creek Suite

In addition to the two gabbros and seven diorites analysed in this study, the fields for gabbros and diorites of the Bakers Creek Suite (Figs. 4A764A10) incorporate all analyses from Hensel

(1982, chapter 7). While gabbros are restricted to a range of ~46653% SiO2, the range for diorites includes quartz diorites, and extends from ~52663% SiO2.

(I) Gabbros: Gabbros of the Bakers Creek Suite show a marked variation for many elements over the restricted silica range. In particular, compatible major elements (Al2O3, MgO, CaO) and related trace elements (Cr, V, Sr) all decrease dramatically with increasing silica, with contents of these elements at least halving over the silica range (Figs. 4A7, 4A8). A concomitant increase in TiO2, 3FeO, Na2O and incompatible trace elements (Ba, Pb, Th, Zn) occurs over the same interval. Although Hensel (1982) treated gabbros and diorites together, and considered crustal contamination had played an important role in the diorites of the Bakers Creek Suite, he also demonstrated that fractional crystallization and accumulation processes were largely responsible for the range in compatible and incompatible elements within the gabbros of the suite. The dramatic changes in inter-element ratios displayed in figure 4A10, are a testament to these process.

Olivine fractionation within the gabbros of the Bakers Creek Suite typifies a tholeiitic (iron- enrichment) trend on the AFM plot (Fig. 4A12c), with most samples plotting within the field for tholeiites. Mafic diorites of the suite also display the same iron enrichment pattern as gabbros, but fall within the calc-alkaline field (Fig. 4A12c).

The tectonic association of gabbroids of the Bakers Creek Suite has primarily been ascribed as that of mid-ocean-ridge (e.g. Hensel 1982, Hensel et al. 1981), with this interpretation lending support to the idea of subduction of a spreading centre underneath the New England Fold Belt in the late Carboniferous (e.g. Murray et al. 1987). However, major and trace element comparison with various basalt types (Table 4A5), shows that the chemistry of Bakers Creek Suite gabbros is somewhat transitional, but more closely matches that of primitive island arc tholeiites than MORB, with depleted contents of high-field-strength elements (Y,

Nb, Zr) and elevated Sr, Ba and Al2O3 (c.f. high-alumina basalts) contents. This 111 interpretation is also supported by the various tectonic discrimination diagrams presented in figure 4A12. Although the fields for arc tholeiites and MORB overlap on some plots (e.g. Fig. 4A12d), the gabbros plot in the volcanic arc basalt field where the two are distinguished (e.g. fig. 4A12a).

Basalt type MORB Island arc Island arc Bakers Creek Calc-alkaline tholeiites Suite gabbros % n=27

SiO2 50A67 49A24 50A49 49A95

TiO2 1A28 0A55 0A70 0A89

Al2O3 15A45 20A16 19A44 17A43 FeO 9A67 8A45 8A85 7A21 MnO 0A15 0A17 0A17 0A14 MgO 9A05 4A80 4A26 8A59 CaO 11A72 11A61 11A66 11A92

Na2O2A51 2A71 2A53 2A45

K2O0A15 0A27 0A14 0A24

P2O5 0A20 0A31 0A11 0A09 ppm Ba 12 58 110 37 Rb 1A044A69 Sr 124 198 200 178 Y 29141219 Zr 85 23 22 58 Nb 3A11A40A72A2

Table 4A5. Comparison of the average analysis of Bakers Creek Suite gabbros, with mid-ocean -ridge (major elements from Best (1982) trace elements from Sun (1980)), island arc calc-alkaline and island arc tholeiitic (South Sandwich arc, Luff 1982) basalts. 112

Hf/3 (a) A = N-type MORB B = E-type MORB and tholeiitic WPB and differentiates (b) CAB = Calc-Alkaline Basalts C = Alkaline WPB and WPB and IAT = Island Arc Tholeiites differentiates MORB = Mid-Ocean Ridge Basalts D= Destructive plate-margin basalts OIA = Ocean Island Andesites and differentiates OIT = Ocean Island Tholeiites

A TiO2/10

B D

C OIT MORB

Th Nb/16 IAT OIA EFeO

MnO´10 P O ´10 Gabbros 2 5 Tholeiitic Diorites Ti/100

Calc-Alkaline

Na2O+K2O MgO D A ´ Nb 2 B C

Zr (d) A,B = LKT Low Potassium Tholeiites Y´3 AI B = OFB Ocean Floor Basalts B,C = CAB Calc-Alkaline Basalts D = WPB Within plate basalts AII B (e) AI-AII = WPA (within plate Alkaline Basalts) AII-C = WPT (within plate Tholeiites) B = P MORB (Mid-Ocean Ridge Basalts) C D = N MORB (Mid-Ocean Ridge Basalts) D C-D = VAB (Volcanic Arc Basalts)

Zr/4 Y

Figure 4. 12. Discrimination plots for Bakers Creek Suite gabbros. (a) Th-Nb/16-Hf/3 (Wood 1980), (b) MnO´ 10- ´ P2 O 5 10-TiO2 /10 (Mullen 1983), (c) AFM plot, (d) Zr-Y-Ti (Pearce & Cann 1973), (e) Zr/4-Y-Nb/2 (Meschede 1986). Solid triangles are gabbro analyses from this study, hollow triangles are diorites (c only). Fields incorporate data from Hensel (1982). 113

(ii) Diorites and quartz diorites: As noted above, mafic diorites of the suite display the same iron enrichment ‘tholeiitic’ pattern as gabbros, but fall within the calc-alkaline field on the AFM plot (Fig. 4A12c). However, more felsic diorites of the suite display a more typical calc-alkaline trend of alkali enrichment with relatively little change in Fe-Mg ratios. This calc-alkaline trend for diorites in figure 4A12c, is further extended by hornblende-bearing granitoids and granophyres of these mafic complexes.

Over a range of 52663% silica, diorites of the Bakers Creek Suite display a large range in major and trace element abundances (Figs. 4A7, 4A8). Diorites at the mafic end of this spectrum generally mirror the linear fractionation trends of the gabbros, and extend this range. However, the combination of these mafic diorites with more felsic diorites into a single field on Harker plots (Figs. 4A7, 4A8), defines a geochemical range that is distinctly two-dimensional rather than linear, covering an area that is in many cases, triangular in shape

(e.g. SiO2 vs. Al2O3, 3FeO, MgO, Na2O, K2O, Cr), but is in other cases more complicated than this (e.g. for many trace elements).

Although a simple process of fractionation easily accounts for the compositional range of gabbros of the suite (Hensel 1982), the broad range of compositions of the diorites, which is frequently non-linear, implies that processes other than fractionation are involved. The triangular shaped field covered by diorites on many Harker plots (Figs. 4A7, 4A8) must be explained by at least two separate differentiation vectors, with each of these vectors representing separate processes. Hensel (1982), recognized this necessity, and proposed the combined processes of fractional crystallization and contamination by metasediments of the accretion complex, to explain the geochemical (and isotopic) range displayed by the diorites. This hypothesis is strongly supported by the observation that the felsic extreme of the triangular diorite field (Figs. 4A7, 4A8), is proximal to the field for metasediments on all Harker plots, a point will be expanded on in section 4A5A5ii. 114

4A4A3 Rocks of accretion complex

The accretion complex of the SNEFB is comprised of a wide variety of rock types, including ocean floor basalts (both MORB and OIB types - Cawood 1984), pelagic material (dominantly cherts), and a range of arc-derived detritus (Cawood & Leitch 1985). Of these three subdivisions, arc detritus is by far the most voluminous component, and varies considerably in rock type. Lithic arenites and greywackes are the most abundant rock types, varying from ‘mafic’ varieties (58-62% SiO2) through intermediate types (62-68% SiO2 -

Plate 4A1a) to more felsic varieties (>68% SiO2). The range of volcanogenic detritus is further extended to higher silica contents by more quartz-rich arenites and wackes (>70%

SiO2), and to lower silica contents by pelites. A total of 21 samples were collected, spanning the lithological range present in the accretion complex. Only cherts and quartzites, which are highly unlikely to be capable of partial melting, were not collected. Most samples were collected from low metamorphic grade sequences. However, analyses of high grade metasediments from the Wongwibinda Complex (Farrell 1992) and the Moona Plains Complex (Hensel 1982) show little, if any, differences to their low-grade counterparts, and are included within the ‘metasediment’ field on all geochemical plots (Figs. 4A764A12).

Sampled accretion complex metasediments have a wide compositional range (Figs. 4A7,4A8, appendix E), varying from 63673% SiO2, with the exclusion of one particularly mafic greywacke at 59% SiO2. While this range often displays a sub-linear trend on Harker plots

(Figs. 4A7, 4A8, e.g. SiO2 vs. 3FeO), differences between pelites and mafic greywackes result in a large spread of compositions at the low SiO2 end of this spectrum. For example, pelites have lower CaO, Na2O, Sr and higher K2O, Al2O3 Ba, Rb, Pb than greywackes at similar silica contents, and hence are also more strongly peraluminous. Additionally, trends for these elements (and many others) on Harker plots (Figs 4A7,4A8), frequently cross-cut the trend of the Hillgrove Suite granitoids at a high angle, with the two trends intersecting at the composition of ‘intermediate’ greywackes. The SiO2 vs K2O and Rb Harker plots (Figs 4A7,4A8) provide the best examples of this feature.

Two amphibolites from the Tia Complex were analysed in this study and are highlighted as individual symbols within the ‘greenstones’ field in Harker plots (Figs. 4A7,4A8). This field also includes data from Cawood (1984) and Stephenson & Hensel (1982). The metabasalts 115 sampled from the Tia Complex represent the alkaline end-member of a compositional spectrum of metabasalts from the accretion complex, which ranges from ocean floor basalts (MORB) to alkaline basalts of OIB affinity (Cawood 1984, Stephenson & Hensel 1982). The alkaline basalts of the Tia Complex have a chemistry typical of OIB magmas, with high alkalies, LILE (e.g. Rb) and HFSE (e.g. Ti, Zr). The high-Ti character of the Tia Complex greenstones, helps to distinguish the metabasalts from the gabbros of the Bakers Creek Suite, which are more tholeiitic in character.

4A5 Petrogenetic Constraints

4A5A1 Parental magmas of the Hillgrove Suite

Before assessing possible source rock types and magma generation processes for Hillgrove Suite granitoids, it is necessary to identify which, if any, of the granitoids sampled, represent parental magmas (i.e. undifferentiated primary magmas). Although most granitoids of the suite fall within the limited range of 68672% SiO2, the causes of geochemical variation within this range need to be elucidated.

The major possible causes of geochemical variation within granitoid suites include fractional crystallization, magma mixing, restite unmixing, and sequential (batch) partial melting. Of these processes, fractional crystallization is most easily distinguished geochemically, since it is essentially a disequilibrium process, and has distinctive effects on the behaviour of compatible trace elements (Arth 1976). Additionally, evidence for the presence of restitic solids (and therefore restite removal), or for magma mixing processes, may be found petrographically. Alleged petrographic evidence for restite includes features such as uniform plagioclase cores, and in peraluminous granites, mineral phases such as sillimanite, garnet or cordierite which display disequilibrium textures, or their secondary equivalents, such as decussate aggregates of biotite and/or muscovite (Chappell et al. 1987). Evidence for magma mixing also includes uniform plagioclase cores, which in this case are inferred to have crystallized in a more mafic magma (Hibbard 1991, Huppert & Sparks 1988). Disequilibrium textures surrounding ferromagnesian phases (e.g. pyroxenes) are also implied to be early magmatic crystallization products (as opposed to entrained restite). Such 116 petrographic evidence (for either the presence of restite or occurrence of magma mixing) is lacking within granitoids of the Hillgrove Suite. However, the features described above are more apparent in the Rockisle Suite members. As discussed earlier, plagioclase cores of uniform composition are rare in the Hillgrove Suite, cordierite does not occur, and garnet is uncommon, as are decussate aggregates of ferromagnesian phases. With a lack of petrographic evidence for magma mixing or restite unmixing, the intrasuite geochemical variation in the Hillgrove Suite granitoids, must be explained by fractional or equilibrium crystallization, or by sequential (batch) partial melting.

Variation within the Hillgrove Suite - fractionation or partial melting? Mass-balance calculations (using least-squares mixing techniques) may be applied in the major element modelling of granitoid suites in order to ascertain if intrasuite variations may be attributed to fractional crystallization. If removal if primary (early-crystallizing) phases is supported by mass-balance, then trace-element methods must be applied to determine whether equilibrium or Rayleigh (fractional) processes were operating during differentiation (Landenberger & Collins 1996). Major-element modelling of the intrasuite variation within the Hillgrove Suite (from 68672% SiO2) was attempted using the average mafic and felsic end-members of this range, together with the compositions of early crystallizing mineral phases from the mafic end-member (plagioclase, biotite, ilmenite). However, the modelled removal of this mineral assemblage from the mafic end-member, failed to reproduce the composition of the felsic end-member. The increasing K2O contents with increasing silica within the suite, caused the rejection of biotite in all models attempted, and replacement of biotite with alkali feldspar in the modelling was also rejected. In turn, modelling with the absence of biotite, failed to match the decreasing contents of FeO and MgO with increasing silica in the suite. Therefore, fractionation (either Rayleigh or equilibrium) is not supported as a mechanism of explaining the variation of major elements within the Hillgrove Suite.

In contrast to the increasing K2O contents with increasing silica within the Hillgrove Suite,

K2O contents of Bundarra Suite granitoids decrease with increasing silica. Consequently, major-element modelling of the limited intrasuite variation within the Bundarra Suite, was successful. The most felsic granitoids of the suite (~75% SiO2) may be modelled as derived from the mafic end-members (~72% SiO2) by a modest 9% fractionation, with removal of 2A6% biotite, 2A9% orthoclase, 3A4% plagioclase and 0A1% ilmenite (residual sum of squares 117

= 0A0538). The moderate increase in incompatible trace-elements (particularly Rb), and the decrease in compatible trace-elements (particularly Ba, Sr, Pb, V, Zn) over this range, are consistent with ~9% disequilibrium (Rayleigh) fractionation.

Having established fractional crystallization as the major process causing the intrasuite variation within the Bundarra Suite, the stark contrast in behaviour of major and trace elements between the Bundarra and Hillgrove suites, also lends support to the rejection of fractional crystallization as a major process within the Hillgrove Suite. Accompanying the contrast in K2O behaviour with increasing silica, are other elements which increase in Hillgrove, but decrease in Bundarra. These include Pb and Y, which probably reflects removal of orthoclase and xenotime, respectively. Furthermore, the sharp changes in inter- element ratios (Rb/Sr, Rb/Ba, K/Ba, Ca/Sr, Mg/V, Fe/Mn) apparent for the Bundarra Suite (Fig. 4A10), which are consistent with fractionation processes, are either not developed in the

Hillgrove Suite, or only occur at the more extreme felsic end of the suite (>74% SiO2). This suggests that if fractionation has occurred within the Hillgrove Suite, it is responsible only for the late, very felsic microgranite dykes.

The only major possibility remaining to explain the chemical variation of the Hillgrove Suite, is the process of batch (equilibrium) partial melting, with the more felsic granitoids derived by lower degrees of partial melting than their mafic counterparts. Consequently, the differentiation trends exhibited by Hillgrove Suite granitoids on Harker and other geochemical plots, are considered to be the result of variable partial melting. This implication, may in turn be used to aid the identification of potential source rocks involved in partial melting (see below). Batch partial melting also solves the problem of increasing

K2O, but decreasing MgO and 3FeO with increasing SiO2, which cannot be accounted for by fractional crystallization, since the only ferromagnesian mineral present in the Hillgrove Suite is biotite. Modelling of the differentiation process as batch partial melting (see below), introduces the possibility of orthopyroxene as a residual phase. Although orthopyroxene is exceedingly rare in the granitoids, is likely to be a common mineral in the source region during partial melting, as it constitutes one of the major breakdown products of biotite dehydration partial melting (Vielzeuf & Holloway 1988). 118

4A5A2 Identifying the source rocks

(I) Constraints on the fertility of source rocks

Experimental work has shown that granitic magmas may be produced from the anatexis of a wide range of crustal protoliths, and that solidus temperatures for water-saturated (wet) partial melting lie in the range 650-700EC (Clemens & Vielzeuf 1987). However, most petrologists now consider that water-saturated melting, or dehydration melting involving muscovite breakdown under amphibolite facies conditions, produce only low volumes of melt (<10%), which are unlikely to efficiently segregate from their source (Clemens & Vielzeuf 1987, Johannes & Holtz 1990). Magmas produced by the breakdown of muscovite, are probably represented in nature by migmatites, such as those present in the Wongwibinda Complex of the SNEFB, which were produced at upper amphibolite facies conditions (Farrell 1992). However, experimental work has demonstrated that the production of large (batholithic) volumes of granitic magma, for which large degrees of partial melting is essential, can only be achieved under water undersaturated conditions, via the breakdown of hydrous ferromagnesian phases at granulite facies conditions of >800EC (e.g. White & Chappell 1977, Clemens & Vielzeuf 1987). In the case of peraluminous source rocks and their partial melting products, this process involves the breakdown of biotite.

Four basic physical parameters determine the melt fertility (ability to produce granitic magma) of common crustal materials during dehydration partial melting: 1. Composition of the source rock. 2. The pressure at which partial melting takes place. 3. Temperature. 4. The hydrous mineral content of the source (Clemens & Vielzeuf 1987)

Most recent theoretical considerations of water undersaturated (dehydration) partial melting (Clemens & Vielzeuf 1987, Johannes & Holtz 1990, Whitney 1988) contend that the water content of any given source rock is the overriding factor in calculating the melt fertility, with solidus temperatures reduced dramatically at any given pressure by increasing a (e.g. H2O Johannes & Holtz 1990). Additionally, for most conditions where a <1, theoretical H2O 119 considerations (Clemens & Vielzeuf 1987, Johannes & Holtz 1990) and experimental work on both synthetic (Patiño-Douce & Beard 1995) and natural starting compositions (Montel & Vielzeuf 1996, Vielzeuf & Holloway 1988), show that above ~2 kb, water undersaturated solidii have positive dP/dT slopes in P/T space (i.e. solidus temperatures increase with increasing pressure). This observation is in contrast to early experimental data on water- present melting, for which the solidus was determined to have a negative dP/dT slope for all pressures. Again, the reason for the observed positive dP/dT slope is directly is related to water activity, with a lowering of a caused by increasing pressure (e.g. Johannes & Holtz H2O 1990).

All of the above theoretical considerations consider that the melt fertility of potential source rocks at given pressures and temperatures, is dependant primarily on the water content of the source rock. Since water undersaturated conditions prevail, then this is directly linked to the percentage of hydrous mineral phases in the source. However, if any potential source rock is deficient in one of the primary low melting constituents (quartz-albite-orthoclase), then its fertility at given P-T-a conditions will be critically restricted. For example, compared to H2O typical granite compositions (and the ternary minimum) in the Q-Ab-Or ternary system, pelites are deficient in Ab, greywackes may be deficient in either the Q or Or component, whereas a quartz rich arenite will be deficient in both the Or and Ab components. As soon as one or more of these components are consumed by partial melting reactions, any further melting must cause the melt composition to deviate from the ternary Q-Ab-Or minimum, and hence an increase in temperature is required.

Melt fertility in terms of Q-Ab-Or composition was also discussed by White & Chappell

(1988). Particular emphasis was placed on Na2O, accounting for the Ab component in the ternary Q-Ab-Or system, and 3FeO, representing the ferromagnesian mineral content, and therefore the amount of water available for dehydration partial melting. This has particular significance for pelites, which are regarded by many researchers as prime candidates for partial melting to produce peraluminous granitoids (e.g. Johannes & Holtz 1990). Although pelites generally have higher FeO and MgO contents than most greywackes, and therefore have higher water availability through the breakdown of biotite, they also contain excess refractory components, and are generally deficient in plagioclase (the Ab component of the Q-Ab-Or ternary system), and will therefore be poor candidates as magma sources (Clemens 120

& Wall 1981, Patiño-Douce & Johnson 1991, White & Chappell 1988). As pointed out by White & Chappell (1988), melt fertility based on water contents alone (e.g. Clemens & Vielzeuf 1987) is model dependant, and does not take into account other geochemical factors.

Melt fertility is thus viewed here as a balance between the ferromagnesian mineral content, and therefore the amount of available water available for melting (a ), and the total content H2O and proportions of the low melting constituents of potential source rocks. For any given ferromagnesian mineral content, the fertility of potential granitoid source rocks may be considered as a function of the total Q+Ab+Or normative content, and the percentage of this content which corresponds to the ternary Q-Ab-Or minimum for a given pressure. The further any given source rock composition is removed from this ternary minimum, the lower the melt fertility will be. This balance corresponds to the ‘fertile window’ described for Lachlan Fold Belt S-type granite source rocks (White & Chappell 1988), which are greywackes with 3-4A5% 3FeO. These greywackes have sufficient water available for partial melting as they have sufficiently high biotite content, but unlike pelites, are not hampered by low Na2O (normative Ab) contents.

Before discussing the melt fertility of potential source rocks from the SNEFB accretion complex, it is necessary to consider the conditions under which partial melting may have taken place. For aluminous metasedimentary sources (both pelites and metagreywackes), melt production during biotite breakdown may be described in general terms by:

Bt + Pl + Qtz (±Als) º melt + Grt/Crd/Spl + OPX ± Ca-Plag/Kfs/Qtz (1) (adapted from Vielzeuf & Holloway 1988, Vielzeuf & Montel 1994) where: Als = aluminosilicate phases (primarily sillimanite in metapelites)

Whether garnet, cordierite or spinel occur as the reaction products, is primarily dependant on the pressure at which the reaction occurs, with cordierite occurring as a reaction product at 2 kb, spinel at 3 kb and 5 kb and garnet at pressures >5 kb (Vielzeuf & Montel 1994). Although most fluid absent melting equations, particularly for pelites, only indicate Kfs as a reaction product in addition to OPX and a peraluminous phase, this equation has been extended here to include both plagioclase and quartz as potential reaction products. Although pelites, which have excess K2O, are likely to produce residual orthoclase, metagreywackes 121 are likely to produce residual Ca-rich plagioclase.

The temperature interval over which this melt reaction occurs is constrained by the OPX-in and biotite-out curves (Vielzeuf & Montel 1994), and occurs in the range 810-860EC at 1 kb, systematically increasing to 890-990EC at 10 kb. These temperatures correspond to granulite facies conditions, and are likely to have been developed within the SNEFB accretion complex at relatively low pressures. Metamorphic mineral equilibria studies in both the Tia and Wongwibinda metamorphic complexes (Dirks et al. 1992, Farrell 1992 - see discussion Chapter 3), have demonstrated that during peak metamorphism in the late Carboniferous, upper amphibolite facies conditions existed at 3-3A5 kb, corresponding to a geothermal gradient in the order of 60EC/km. At this gradient, granulite facies conditions, with temperatures in the range ~800-850EC, would have been attained at a pressures of 4-5 kb. Correspondingly, the temperature interval for melting reactions (equation 1) at 5 kb, is 820- 880EC (Vielzeuf & Montel 1994).

Having established the likely P/T conditions at which partial melting took place within the SNEFB accretion complex, it is now possible to estimate the relative melt fertility of various granitoid source-rock candidates. For this purpose, a function combining the total contents of low melting (minimum-melt) components (Q+Ab+Or) with the percentage of these constituents which correspond to the 5 kb ternary eutectic was devised, and values were calculated for individual rock compositions. This function is referred to herein as the melt fertility index or MFI. The MFI may be defined as follows:

(2)

The first part of equation 2 is the sum of the normative Q, Ab and Or components (expressed as a percentage), while the second part of the equation is a matrix division calculation. This equation calculates the percentage of the 5 kb eutectic (minimum melt) composition in potential source rocks. The ternary eutectic composition used in this equation (Q:Ab:Or = 28:50:22) is that determined for water saturated conditions at 5 kb (Luth et al. 1968). Although it has already been concluded that water saturated conditions did not apply, the ternary eutectic compositions for a <1 are not well constrained experimentally, in either H2O haplogranite systems or for natural starting materials. Most recent experimental data indicate 122 that lowering of a causes a shift in the Q-Ab-Or ternary eutectic towards the Or apex. H2O Preliminary data for the haplogranite system (Johannes & Holtz 1990) at 5 kb, suggests that lowering a from 1 to 0A5 causes a shift in the Q-Ab-Or eutectic from 28-50-22 to 30-42- H2O 28. Experimental data using natural starting materials indicate a similar shift in the eutectic compositions (of similar magnitude) towards the Or apex (Conrad et al. 1988, Holtz & Johannes 1991). Although this shift is significant in terms of melt compositions, available data suggest that the direction and magnitude of the shift varies for different starting materials, and additionally, the precise a conditions during generation of the Hillgrove H2O Suite magmas are unknown. Furthermore, shifts in the eutectic composition of the magnitudes measured, will not significantly affect calculation of the MFI, except for source rock compositions that lie close to the eutectic. These rocks will have high fertility regardless of what eutectic composition is used in the equation.

The MFI calculation above, indirectly accounts for the biotite contents of potential metasedimentary sources, and therefore available water during partial melting. A simple measure of the likely biotite contents of potential metasedimentary sources prior to biotite breakdown reactions is the sum of FeO and MgO in each sample (White & Chappell 1988). For mafic greywackes of the SNEFB, combined 3FeO and MgO contents approach 10%, whereas both pelites and intermediate greywackes have ~6% 3FeO+MgO. The value for 3FeO+MgO only drops below 5% for felsic greywackes and quartz-rich psammites, which often exceed 70% SiO2. These are unlikely candidates for the generation of granitoids, as most Hillgrove Suite granitoids are more mafic than this. The low fertility of these more silicic sediments is also reflected in low values for MFI (see below), due to low Na2O and

K2O contents. Hence, although the MFI calculation does not contain variables directly accounting for water content of the source (such as 3FeO), the low alkali contents of rock- types which have low 3FeO+MgO, consequently give low MFI values for these silicic rock types. Water availability during dehydration partial melting, is therefore not considered as a property that distinguishes the fertility between pelites and greywackes, since all of these source rock types are likely to contain ample contents of biotite (most would exceed 18% biotite) before melting occurs. 3FeO+MgO contents that span the range pelites and mafic to intermediate greywackes from the accretion complex of the SNEFB, correspond to the ‘fertile window’ identified by White & Chappell (1988). 123

The calculation of MFI (equation 2) can thus be used to estimate the melt fertility of all potential metasedimentary source rocks from the accretion complex. As mentioned above, siliceous psammites and felsic greywackes have low fertility, with MFI values not exceeding 30%. Although these silicic rocks contain high 3(Q+Ab+Or) contents, the Ab and Or portions of this sum are very low, and hence their composition is far removed from the Q-Ab- Or ternary eutectic (Fig. 4A13). At the opposing end of the metasedimentary spectrum of the accretion complex, the mafic greywackes (~58-63%SiO2), also have low fertility, with MFI values not exceeding ~30%, with most <20%. Unlike their silicic counterparts, these greywackes have compositions which plot relatively close to the ternary eutectics (Fig. 4A13), but they have low 3(Q+Ab+Or) contents.

Most pelitic rocks (particularly the more mafic types) also have low calculated MFI values (<30%). These low values are a result of low to moderate 3(Q+Ab+Or), combined with low Ab contents, and hence plot remote to the Q-Ab-Or eutectics (Fig. 4A13). Hence, although pelites are regarded by many experimental petrologists as fertile source rocks due to high water availability, other factors of their composition (i.e. Q-Ab-Or contents) generally rule them out as fertile sources. This conclusion also supported by the isotopic studies presented herein (see section 4A5A3iii).

Greywackes of intermediate composition (63-68% SiO2) have the highest MFI values of the accretion complex metasediments, with values always exceeding 50%, and approaching 90% for the more felsic members of this range. However, for greywackes more felsic than 68%

SiO2, fertility is decreased due to lower contents of alkalis, as discussed above. Although 3(Q+Ab+Or) contents for this group are only moderate compared to their more felsic counterparts, close proximity to the Q-Ab-Or eutectics (see Fig 4A13) results in high MFI values. 124 Q

Hillgrove Bundarra Metasediments Average ‘fertile’ greywacke Siliceous New England Psammites Calcareous sediments Average ‘unfractionated’ Greenstones Hillgrove Suite FG = felsic greywacke Greywacke, TVZ IG = intermediate greywacke (fertile source) Initial melt composition MG = mafic greywacke @ a = 025. FG Pelites HO2

IG 1 3 5 10 kb dry eutectic MG 10

30 Water-saturated eutectics (kb).

Ab An An Or

7kb (dry) cotectic

Ab Or Figure 4. 13. Normative Q-Ab-Or and An-Ab-Or plots of Hillgrove Suite and Bundarra Suite granitoids and potential source rocks. Ternary water-saturated eutectic compositions for the haplogranite system are taken from Luth et al. (1964), Tuttle & Bowen (1958) and 10 kb dry eutectic from Huang & Wyllie (1975). Position of the An-Ab-Or 7 kb cotectic trough is from Winkler (1979). Compositions for the average fertile greywacke from the Taupo Volcanic Zone and experimental melt compostions are from Conrad et al. (1988). Additional data for sediments and the calcareous sedimentary field are taken from Farrell (1992) and Hensel (1982). 125

The strong contrast in MFI values between these fertile intermediate greywackes (MFI>50%) and other rock types, generally with MFI<40%, allows the division between fertile versus relatively infertile sources to be set at 50% MFI. Although this is an arbitrary value, it does serve as a distinction between highly fertile sources, and those which are less likely to produce large melt volumes, at given P, T and a conditions. The average composition of H2O these fertile intermediate greywackes is plotted on all geochemical plots as a star symbol (Figs. 4A07-4A10) and as a blue star in the Q-Ab-Or plot (Fig. 4A13). The average major and trace element values of this average are also presented as part of the melt modelling in Table 4A6. Thus, the average fertile greywacke is intermediate in terms of silica content (65A46%

SiO2), is rich in 3FeO (4A47), Na2O (3A62%), CaO (3A31%), but poorer in K2O (2A27%) and MgO (1A77%), and is mildly peraluminous (1A11 ASI). In particular, the calcium and alkali contents of this composition contrast strongly with the fertile Ordovician metasediments of the LFB, which are poorer in CaO (0-1%) and Na2O (1-2%) contents, but richer in K2O (2- 5%), and are more strongly peraluminous (ASI 1A5-3A7, White & Chappell 1988). This difference is a function of the low chemical maturity of greywackes of the SNEFB accretion complex, compared with their LFB counterparts, and is reflected in their textural immaturity.

At high metamorphic grade prior to biotite breakdown (upper amphibolite facies), this average fertile greywacke composition corresponds to a quartzofeldspathic gneiss with modal proportions of 32% quartz, 21A5% biotite (Mg# 38), 46% plagioclase (An35), and 0A5% ilmenite. Such rock types are relatively common in upper amphibolite facies sequences within the Tia and Wongwibinda metamorphic complexes (Dirks et al. 1992, Farrell 1992), and are typically the rock types which show incipient partial melting (migmatization) at this grade. In contrast, rocks of pelitic composition in these complexes are relatively inert, and appear to contribute little to leucosome formation (Farrell 1992). 126

(ii) Constraints from major and trace elements: Batch partial melting trends in granitoid suites - an example of restite removal at the source? As discussed earlier, models of fractional crystallization, magma mixing and restite unmixing, have been rejected as major causes of the linear variation displayed by the Hillgrove Suite on Harker plots. Rather, it was concluded that these trends are produced by processes of batch (equilibrium) partial melting. A model explaining the geochemical trends within the Hillgrove Suite are discussed in more detail here. Importantly, batch partial melting trends, which involve the removal of differing proportions of magma from residual source, and therefore involve restite removal. The model to be considered here, is thus based on the restite model, but with several critical modifications.

In the restite model, Chappell et al. (1987) explained the linear geochemical variation displayed on Harker diagrams in many granitoid suites, by progressive removal of restite (unmelted residuum from the source) which is entrained in the magma as it rises en mass from the source region. This theory implies that the linear geochemical trends, when projected to more mafic compositions, will intersect the compositions of both the source rock and the entrained restite. The model views all restite-controlled suites as mixtures of melt, which is of minimum melt composition (>75% SiO2), and entrained restite, with more mafic magmas containing successively greater proportions of restite (Chappell et al. 1987). Figure 4A14a (adapted from Chappell et al. 1987) depicts the important geochemical concepts of the restite model. Importantly, a specific source composition is inferred for each suite, and in particular, the composition of the restite is fixed. In turn, since the composition of restite in each suite is fixed, then extrapolation of the geochemical trends must intersect both source and restite compositions for all elements plotted on Harker plots. However, as noted by Wall et al. (1987), although these extrapolations are correct for most elements on Harker plots within LFB granitoid suites, it never works for all elements within a given suite, with widely varying proportions of restite required for individual elements in some cases.

For the Hillgrove Suite, although extrapolation of differentiation trends intersect the source composition (as calculated above) for many elements (Al2O3, MgO, CaO, Na2O, K2O, Rb, Pb, Y, Nb, Th, see Figs. 4A7,4A8), it does not intersect for all. Notable exceptions include

TiO2, 3FeO, P2O5, Ba, Sr and Zr (see Figs. 4A7,4A8). This problem is resolved if the linear trends on Harker plots are considered as batch partial melting trends. 127

R (entrained Restite) Figure 4. 14 (a). Linear geochemical variation in a granitoid suite according to the restite theory. ‘L’ (a) represents the melt composition (minimum melt i S component) in all magmas, with more mafic magmas (1 to 4) containing successively greater proportions of entrained restite (R) of fixed composition. The 4 progressive removal of entrained restite (from L) in more mafic magma compositions is the mechanism of In % element c 3 rea sing differentiation to more felsic magmas and the linear en trained prop 2 geochemical variation. Note that the projection of the o r rt differentiation trend intersects both source (S) and restite es ion of tite 1 (R) compositions. Adapted from Chappell et al. (1987).

%SiO2 F

R (Residuum remaining R (Residuum remaining in source region) in source region) In 4 cr ea (b) si (d) ng S d

I in residuum e 3 gr ee 4 of S pa Inc r t re ia as l % element i in 3 2 m g % element i el (compatible in residuum) d t eg in re g e 2 o i decreases rapidly in residuum with f p a (compatability in residuum diminishes rt 1 ia 1 with increasing degree of partial melting) l m el element t the disappearance of mineral %SiO ing 2 F %SiO2 F

R (Residuum remaining F in source region) 1 (c) (e) 2 g

3 tial meltin I in residuum ar f p e o re S 4 eg sing d S ea i increases with increasing

Incr % element i % element i

(incompatible in residuum) I ncreasi 4 element component of mineral n (compatability in residuum increases g degr 3 with increasing degree of partial melting) ee of 2 R (Residuum remaining part 1 in source region) ial m %SiO2 elting F %SiO2

Figure 4. 14. Cases (b) to (e) represent different instances of linear geochemical variation in a granitoid suite as generated by batch (equilibrium) partial melting. The most felsic magma (F) is the minimum melt composition. The remainder of the suite (magmas 1 to 4) are progressively more mafic non-minimum melts, resulting from increased degrees of partial melting. In contrast to the restite model, the ‘restite’ (R, termed residuum here) is not necessarily of fixed composition, and is not entrained in the magma, but rather, remains in the source region. (b) Element i is equally compatible in the residuum for all partial melting cases. As in case (a), projection of the differentiation trend intersects both the source (S) and residuum (R) compositions. (c) Element i is incompatible in the residuum for all partial melting cases. As in case (a) and (b), the differentiation trend intersects both the source (S) and residuum (R) compositions. (d) Element i is compatible in the residuum for low degrees of partial melting, but rapidly becomes incompatible with the disappearance of mineral phase I (the major phase containing element i), from the residuum. As such, the concentration of element i in the residuum (R) rapidly decreases with increasing degrees of partial melting. The resulting differentiation trend is still linear, but intersects neither the source nor residuum compositions. (e) Element i is moderately compatible in the residuum for low degrees of partial melting, and compatibility increases as the proportion of mineral phase I (the major phase containing element i) increases in the residuum. As such, the concentration of element i in the residuum (R) increases with increasing degrees of partial melting. As in case (d) the differentiation trend is still linear, but intersects neither the source nor residuum compositions. 128

This model may be applied to all suites in which internal variation is demonstrably not generated by magma mixing or fractional crystallization. In the case of batch partial melting, these linear trends are still technically generated by the removal of restite (or residuum) from magma, but the process involved has several crucial differences to the restite model of Chappell et al. (1987). Firstly, the varying magma compositions within individual suites are not viewed as proportional mixtures of minimum melt composition and entrained restite, but rather as successively more mafic non-minimum melts (Fig. 4A14). Secondly, the residuum (restite) remains in the source region after melt extraction, and is not entrained in the magma. Thirdly, since magma batches of different compositions are generated by extraction from a residuum at the source, through different degrees of partial melting, then the residuum itself cannot be of fixed composition (Fig. 4A14), as in the restite model. Although silica variation in the residuum may change little (see section 4A5A4ii), contents of other elements may change markedly, since varying degrees of partial melting will result different mineral assemblages as the residuum. Finally, the highly felsic end-members of suites (>75% SiO2) which are near to the inferred ‘minimum melt’ composition, are likely to have been generated by fractional crystallization, rather than existing as primary minimum melts generated from the source. Theoretical considerations (Wall et al. 1988) suggest that minimum melts are probably the exception rather than the rule, and experimental observations (Conrad et al. 1987) demonstrate that minimum melts cannot be produced by partial melting of natural materials, and hence are more likely to be fractionated products of more mafic magmas (a point conceded by Chappell 1996).

Four separate theoretical examples of linear trends on Harker plots as generated by batch partial melting, are presented in Figure 4A14, where they are compared and contrasted with the restite model of Chappell et al. (1987). These examples can be used to explain all examples of linear geochemical variation the Hillgrove Suite. Elements which are retained in the residuum in similar contents for varying degrees of partial melting, through equal compatibility (Fig. 4A14b) generate linear geochemical trends which do intersect the source and residuum compositions. Examples of this behaviour in the case of the Hillgrove Suite are Al2O3, MgO, CaO and Na2O (Fig. 4A7). Likewise, elements which are equally incompatible for varying degrees of partial melting also generate linear geochemical trends which intersect the source and residuum compositions (Fig. 4A14c). Examples of this behaviour in the case of the Hillgrove Suite are K2O, Rb, Pb, Y and Th. Although these 129 instances are superficially similar to instance ‘a’ (restite unmixing), the residuum may vary slightly in composition, but is still restricted to a composition somewhere on the extrapolation of the differentiation trend.

In contrast, the concentration of some elements in the residuum changes dramatically with variation in the degree of partial melting. In Fig. 4A14d, element I is compatible in the residuum for small degrees of partial melting, but as the degree of partial melting increases, and mineral I (the major phase containing element I) is completely consumed by the partial melt reaction, element I rapidly becomes incompatible, and is thus partitioned into the magma. Although this instance produces a linear geochemical variation for the granite suite, the source and residuum compositions both lie below the extrapolation of this trend. This situation applies to TiO2, 3FeO, P2O5, Ba and Zr in the case of the Hillgrove Suite (Fig 4A7,

4A8). For the case of P2O5 the phase in question would be apatite, which would rapidly decline in the residuum, with increasing degrees of partial melting. The opposite situation is depicted in Fig. 4A14e, in which element I increases in concentration in the residuum with increasing degrees of partial melting. In this instance, the differentiation trend is also linear, but the source and residuum compositions lie above the extrapolation of this differentiation trend. The major example of this behaviour in the Hillgrove Suite is that of Sr. Although plagioclase is likely to exist in the residuum for all degrees of partial melting, the modal proportion of plagioclase would increase as other phases are consumed by the partial melting reactions, and hence the concentration of Sr would also increase.

The behaviour exhibited in both instance ‘d’ and ‘e’ (Fig. 4A14) typifies the likely behaviour of trace elements in particular, since their partitioning into residual phases is likely to be more sensitive to small changes in the modal mineralogy of the residuum. Conversely, most major elements (and trace elements which are incompatible for all cases), will display linear trends which do intersect source compositions. As such, although extrapolation of the differentiation trends for trace elements on Harker plots are unlikely to be indicative of source compositions, extrapolation of major element trends is likely to prove more useful.

With this in mind, a review of the major and trace element Harker plots (Figs. 4A7, 4A8), reveals that, for most major elements, and incompatible trace elements, the composition of the average fertile greywacke calculated above (star symbol), is intersected by extrapolation 130 of the differentiation trends. Thus, application of the MFI for calculation of the average fertile metasediment of the SNEFB accretion complex, is strongly supported by the source composition, as inferred by extrapolation of these differentiation trends. However, the behaviour of those compatible trace elements (in addition to TiO2 and 3FeO) for which the differentiation trends do not intersect this source composition, is satisfactorily explained by the theoretical considerations detailed above.

(iii) Isotopic Constraints

An isotopic study of both the granitoids and potential source rocks from the accretion complex sequences was undertaken to further constrain the potential source rocks for the Hillgrove Supersuite granitoids. Nd and Sr isotopic ratios were measured from nine samples of the Hillgrove Supersuite (seven from the Hillgrove Suite and two from the Rockisle Suite), along with six samples that spanned the range of exposed metasedimentary rock types. In addition, a single sample of high-Al gabbro from the Bakers Creek Suite was analysed, in order to ascertain if these mantle-derived intrusives had played any role in granitoid petrogenesis.

Initial Sr and Nd ratios, and gNd values were calculated for all samples at 300 Ma (the age of Hillgrove Suite intrusion - Collins et al. 1993, Kent 1994, also see Chapter 3). These results are presented on a conventional initial 87Sr/86Sr - gNd plot in figure 4A15, and the data is tabled in appendix F. For comparison, the field for Lachlan Fold Belt S-type granites is included (McCulloch & Chappell 1982). Isotopic data from Hensel et al. (1985) for Hillgrove Suite granites and accretion - complex metasediments were recalculated at 300 Ma and are also plotted in figure 4A15, along with two samples of the Bundarra Suite, recalculated at 285 Ma (the age of Bundarra Suite intrusion - Collins et al. 1993).

Metasediments from the accretion complex sequences are spread over an elongate, narrow field in figure 4A15, ranging from mafic greywackes at the isotopically primitive extreme, 131 . 0 718

Hillgrove Supersuite } . 0 716 LFB S-types . 0 714 Hillgrove Hillgrove Suite Rockisle Suite Bakers Creek Suite (gabbro) Metasediments ‘Fertile’ (intermediate) greywackes LFB Cambrian (Tablelands (Tablelands Complex) greenstones (tholeiites) Pelites MG = mafic greywacke FG = felsic greywacke . i trends - line 1 has calculated isotopic mixing Lines labelled 1 and 2 are contamination Initial ratios calculated @ 300 Ma except Initial ratios calculated @ 300 Ma compositions at 10% intervals - see text for explanation. Bundarra Suite samples @ 285 Ma. Bundarra Suite samples @ 285 0 712 Sr/ Sr . 87 86 0 710 . 0 708 87 86 FG 2 (Burke et al. 1982) . Hillgrove Suite Metasediments Hillgrove Suite Metasediments Bundarra Suite 0 706 Cambrian - Carboniferous { { Maximum seawater Sr/ Sr MG . seawater alteration 0 704 al. (1985).

1 et Nd Current study MORB e Hensel . 8 6 4 2 0 0 702 -8 -6 -2 -4 Values Values recalculated from 10 -10

Figure 4. 15. Initial 87 Sr/ 86 Sr and eNd results for New England metasediments and Hillgrove Supersuite granitoids. (Field for LFB S-type granites - Mc Culloch & Chappell 1982, field for LFB greenstones - Nelson et al. 1984). 132 through intermediate and felsic greywackes to pelites at the most evolved end of the spectrum. The large spread of 87Sr/86Sr ratios (at 300 Ma) in this range is strongly influenced by high 87Rb/86Sr ratios in pelitic samples (e.g. W443=6A04), while most greywackes have ratios of <1. Granitoids from the Hillgrove Suite are more restricted, defining a range which only partly overlaps with the field spanned by metasediments. The two Bundarra Suite samples plot at the most isotopically evolved extreme of the Hillgrove Suite range (Fig. 4A15). The single analysed Bakers Creek Suite gabbro sample from the Sheep Station Creek Complex has the highest gNd value (+9@55) and the lowest initial 87Sr/86Sr (0@70276) for any analysed sample from eastern Australia. Such depleted isotopic ratios are consistent with geochemical evidence which suggests that these gabbros are primitive arc magmas, of high alumina tholeiite (HAT) affinity (see previous discussion).

The Hillgrove Suite granitoids overlap with the accretion complex rock-types within a restricted field (Fig. 4A15). This narrow field corresponds to greywackes of intermediate composition (~65% SiO2). The overlap can be interpreted in two ways: (1) the source rocks are the inferred intermediate metagreywackes, and hence the isotopic data supports the geochemical considerations above; or (2) the granitoids are result of mixing of gabbro and pelite. Although the isotopic composition of Hillgrove Suite granitoids may theoretically be produced by mixing of isotopically distinct sources, if the isotopic data were considered in isolation, the low melt fertility of pelites (as discussed above) renders such a scenario unlikely. Additionally, the required mix of ~80% gabbro and ~20% pelite predicted by isotopic mixing calculations, cannot be matched by major and trace element calculations predictions. Major element modelling using these proportions result in a dioritic composition.

The hatched area denoted ‘fertile greywackes’ (Fig. 4A15) includes all metasedimentary rocks which have calculated MFI values of greater than 50%. This ‘melt-fertile’ field corresponds to the intermediate greywackes (63-68% SiO2 - as discussed above) which isotopically overlap with the granitoids. Hence, for the production of those Hillgrove Supersuite granitoids which isotopically overlap with the field for metasediments, the evidence is consistent with simple partial melting of these intermediate greywackes. The lithologies with the lowest MFI values (<25% - as calculated above) are mafic greywackes and pelites, which correspondingly form the two isotopic extremes of the metasedimentary field, which are most remote from the area of Hillgrove Suite granitoids (Fig. 4A15). 133

The range of isotopic compositions for some analysed Hillgrove Supersuite samples extends beyond the field of fertile metasediments, and beyond the field of all analysed accretion complex metasediments, suggesting that other processes, or other magma sources were involved in the generation of these granitoids. Possible magma sources and mechanisms for this contamination will be discussed below (section 4A5A5).

Unlike the Hillgrove Suite, the two samples of Bundarra Suite granitoids plotted on figure 4A15 are more restricted in isotopic composition, and plot at the most isotopically evolved end of the Hillgrove Supersuite spectrum, coinciding with the field for ‘fertile’ greywackes. Although this lack of variation in isotopic composition may simply be a result of lack of sampling, the mineralogical and geochemical homogeneity throughout this suite (Shaw & Flood 1981), suggests that the Bundarra Suite granitoids are essentially direct partial melts from metasedimentary source-rocks, with no contamination from other crustal or mantle sources.

4A5A3 Quantification the partial melting process

(I) Incompatible trace-elements: inferences on the degree of partial melting

Empirical calculation of the average fertile source rock involved in the generation of Hillgrove Suite granitoids allows modelling of the partial melting processes involved in magma generation. Since all modelling of partial melting processes requires knowledge of the degree of partial melting, in addition to source and product compositions, then a major prerequisite for modelling is the estimation or calculation of this value.

The behaviour of any given trace element I during partial melting may be described by:

(Arth 1976)

i i Where: Cmelt and Csource are the concentrations of element I in the melt and source respectively, F is the fraction of partial melt (as a fraction of the source), and D is the bulk distribution coefficient for element I in the residual mineralogy. 134

For trace elements which are incompatible in the residuum, the value for D approaches zero, and if complete incompatibility is assumed (i.e. D=0), then the above equation becomes:

i i This, in turn may be used to estimate the degree of partial melting, since if Cmelt and Csource are known, then rearranging the above expression gives:

This equation was applied to the Hillgrove Suite and its primary source, using the average analysis for fertile greywackes (section 4A5A3i) and the average ‘mafic’ Hillgrove Suite granitoid (those samples with <150 ppm Rb - see table 4A4). The incompatible elements used in this instance were chosen on the basis of the likely mineralogy of the residuum, which would comprise OPX+Ca-plag±Kfs/Qtz±Grt/Crd/Spl±monazite, and hence the trace elements applied, were Rb, Th and Y. The results are summarized as follows:

element Concentration in Concentration in Calculated source (ppm) magma (ppm) F Rb 67A9 141 0A48 Th 7A64 15A60A49 Y22A246A10A48

Although incompatibility of Rb is dependent on the assumption that little or no biotite is left as residuum, the remarkable consistency of results for these three elements is supports this assumption. Additionally, although garnet may be theoretically present in the residuum, the ratio calculated for Y is also consistent with other ratios, suggesting that garnet is not a residual phase, which given that the source is only weakly peraluminous, is hardly surprising. Ratios calculated for all other elements (which are more compatible in the residuum) are expectedly higher, since D is no longer near zero for these elements.

The above results therefore imply that production of the Hillgrove Suite magmas involved a degree of partial melting of ~50%, and an average value for F of 0A48 will be used in subsequent modelling. The full trace element modelling in section 4A5A4iii will be used as a cross-check on the assumptions applied above. 135

This calculated degree of partial melting is in accord with recent experimental work. Experimental melting of metasediments under water-undersaturated conditions, indicates that complete biotite breakdown in a given source rock yields melt percentages in the order of 30- 70% by volume (Conrad et al. 1988, Patiño-Douce & Johnston 1991, Vielzeuf & Holloway 1988, Vielzeuf & Montel 1994), depending primarily on source composition, a , P and T. H2O Similar degrees of partial melting are also supported by thermodynamic calculations (Clemens & Vielzeuf 1988).

In addition to the above calculation of the degree of partial melting for ‘mafic’ Hillgrove Suite granitoids, ratios were also calculated for the ‘felsic’ Hillgrove Suite granitoids (200>Rb>190ppm, see Table 4A4), since if the differentiation trend exhibited by the Hillgrove Suite is inferred to be a partial melting trend, then these more felsic magmas must have been generated by lower degrees of partial melting. The results obtained are summarized as follows:

element Concentration in Concentration in Calculated source (ppm) magma (ppm) F Rb 67A9 198 0A34 Th 7A64 22A00A35 Y22A237A90A59

Although the expected lower value for F is supported using the ratios for Rb and Th (~35%), the ratio for Y is unexpectedly high. This suggests that the production by lower degrees of partial melting of these more felsic granitoids produced a residuum more complex in mineralogy, with Y no longer behaving incompatibly. Although garnet is a possible candidate for retaining Y in the source, as discussed above, it is unlikely to have occurred in sufficient quantities in such a weakly peraluminous source. Xenotime is a widespread accessory mineral in the Hillgrove Suite granitoids, and is likely to occur in the residuum, particularly for lower degrees of partial melting. Since Y is essentially a major element in xenotime, little of this phase would be required in the residuum to markedly decrease contents of Y in the melt, which might explain the high value for F implied by the ratios. Hence, although the calculated value for F in this case is higher for Y, the good agreement between the ratios for Rb and Th suggests that a degree of partial melting of ~35% could be applied to modelling these more felsic Hillgrove Suite variants. 136

(ii) Major element modelling: calculation of the residuum

Before any extended trace element modelling of equilibrium partial melting can be conducted, estimation of the composition and modal mineralogy of the residuum must be made. Combination of the degree of partial melting (F, as calculated above) combined with the known compositions of the source and derived magmas, a simple mass balance calculation of the major element composition of the residuum may be made. For 48% (F=0A48) partial melting then:

source = 0A48×magma+0A52×residuum

Hence, the concentration of each major element I (in oxide weight %) may be expressed as:

Similarly, for 35% (F=0A35) partial melting:

Using the two formulas above, major element residuum compositions were calculated for the production of both the mafic and felsic end-members of the Hillgrove Suite. The results are presented in table 4A6. 137

Table 4A6. Major element modelling. Average Average Average Residuum Residuum ‘Mafic’ ‘Felsic’ Fertile after 48% after 35% Hillgrove Hillgrove Greywacke melt removal melt removal (=48% melt) (=35% melt)

SiO2 69A11 70A83 65A46 62A09 62A56

TiO2 0A63 0A52 0A65 0A67 0A72

Al2O3 14A35 14A03 15A51 16A58 16A31

Fe2O3 0A53 0A31 0A84 1A13 1A13 FeO 3A36 2A64 3A72 4A05 4A30 MnO 0A10 0A06 0A07 0A04 0A08 MgO 1A22 1A01 1A77 2A28 2A18 CaO 2A31 2A12 3A31 4A23 3A95

Na2O 3A28 2A96 3A62 3A93 3A98

K2O 3A65 4A40 2A27 1A00 1A12

P2O5 0A14 0A10 0A13 0A12 0A15 CIPW norm quartz 28A02 29A72 24A80 21A84 22A40 corundum 1A16 0A80 1A43 1A70 1A78 orthoclase 22A01 26A41 13A86 6A37 7A07 albite 28A28 25A44 31A69 34A76 34A79 anorthite 10A76 10A02 16A13 21A04 19A22 hypersthene 8A09 6A46 9A93 11A68 10A78 magnetite 0A77 A46 1A26 1A71 1A68 ilmenite 1A22 1A00 1A28 1A32 1A41 apatite 0A31 0A22 0A29 0A26 0A35

In order to estimate the modal mineralogy of the residuum, CIPW normative calculations were applied to the resulting residuum compositions. If the previous assumption that biotite is completely consumed by melt reactions is valid, then these normative calculations provide a good approximation of the actual modal mineralogy of the residuum. Thus, after extraction of the average ‘mafic’ Hillgrove Suite granitoid from its fertile greywacke source (~65A5%

SiO2), the residuum remaining in the source region, would be a granulite of intermediate composition (~62% SiO2), with the approximate modal mineralogy of: 22% quartz, 55% plagioclase (~An40), 7% orthoclase (estimated at Or90), 12% orthopyroxene, and 3% oxide phases. The normative corundum value calculated for the residuum (~2%) is not included in this modal estimation, and may be present as either garnet, cordierite or spinel. Calculations for a 35% partial melt to produce the more felsic granitoids, produces a residual composition (~62A5% SiO2) which is very similar to that produced by 48% partial melting.

The similar calculated residual compositions for production of both mafic and felsic end- members of the Hillgrove Suite, suggest that the differences in resulting magma composition are primarily a product of the degree of partial melting (rather than 138

FR MR 0.7 4 MR FR S S CaO TiO2 MH 0.6 3

FH FH MH 0.5 2 60 62 64 66 68 70 72 60 62 64 66 68 70 72 SiO2 SiO2 17 MR FR FR 4 16 S MR S

MH 15 Al2 O 3 Na2 O

3 FH 14 MH FH 60 62 64 66 68 70 72 60 62 64 66 68 70 72

SiO2 SiO2

FR 4 FH 5 MR MH S 3 EFeO S 4 MH 2 KO2

FH 1 3 MR FR

60 62 64 66 68 70 72 60 62 64 66 68 70 72 SiO SiO2 2

FR MR FR MH 2 0.14 S PO2 5 MgO 0.12 S MH MR FH FH 1 0.10

60 62 64 66 68 70 72 60 62 64 66 68 70 72

SiO2 SiO2 Figure 4. 16. Harker plots of residuum compositions calculated by major element modelling. FH = ‘felsic’ Hillgrove, MH = ‘mafic’ Hillgrove, S = fertile greywacke source, FR = calculated residuum from 35% partial melting to produce of FH, MR = calculated residuum from 48% partial melting to produce . MH. A comparison with Fig. 4 14 shows that MgO, Na2 O and Al2 O 3 closely approximate instance ‘b’, K2 O is instance ‘c’, TiO2 and P 2 O 5 are instance ‘d’, and CaO is instance ‘e’. 139 differences in the residual mineralogy). Such similar residuum compositions seem at odds with the earlier inferences that the differentiation trends in the Hillgrove Suite are partial melting trends. However, if the inferred residual modal mineralogy is calculated as a percentage of residuum + magma (rather than just the residuum), then this apparent contradiction is resolved. For example, the inferred modal proportions of orthopyroxene in the residuum for the case of 48% partial melting is ~12%, and is actually lower in the case of 35% partial melting (~11%), but nonetheless produces a more felsic granitoid. Recasting of these values as percentages of melt+residuum gives values of ~6% modal OPX for 48% partial melting and ~7% for 35% partial melting. This ‘modal’ mineralogy, expressed as a percentage of residuum+melt is also presented in table 4A6. Additionally, Harker plots demonstrate that the calculated residua compositions (Fig. 4A16) follow the behaviour predicted earlier by different instances of partial melting (Fig. 4A14).

(iii) Trace element modelling of the granitoid products of equilibrium partial melting - cross-checking the assumptions.

In addition to the application of incompatible trace elements to estimate the degree of partial malting, modelling the behaviour of all trace elements during partial melting may be applied. Models were calculated using the modal proportions of residuum predicted by major element modelling, and the degree of partial melting predicted earlier, in order to validate earlier assumptions. As stated earlier discussed the behaviour of any given trace element I during (equilibrium) partial melting may be described by:

Trace element concentrations were thus calculated for the ‘mafic’ Hillgrove Suite granitoid average, using the above equation for equilibrium partial melting. The value of 0A48 for F (calculated above) was used, together with the composition of the fertile greywacke source, and published partition coefficients (Arth 1976, Henderson 1982). Predicted melt compositions are compared with the actual composition in Table 4A7. The mineral 140 proportions used in these calculations are as fractions of the total removed mineralogy (as required by the above equation), and not as percentages of the parent melt as calculated in the major element modelling.

Table 4A7. Trace Element Modelling of 48 % partial melting of fertile greywacke to produce the ‘mafic’ Hillgrove Suite composition.

Mineral OPX Biotite Plag- Alkali Mt Il Ap Quartz Actual Predicted Predicted Actual Resid- ioclase F’spar Source Melt (1) Melt (2) Melt uum Ba 0.0036.360.366.120000505626624634205.27 Rb 0.0033.260.0480.340000681341271417.18 Sr 0.009 0.12 6.00 3.87 0 0 10 0 468 202 202 181 308.76 Pb 000.4510000142222233.42 Th 0 0.5 0 0.02 0 0 1.1 0 8 16 16 16 0.11 U 00.100.02002.9012230.01 Nb 0.86.370.030.0062.5400111816122.81 Y 1 1.23 0.06 0.006 0.92 2.09 44 0 22 34 34 46 5.82 La 0.11 0.33 0.3 0.053 0.67 3.6 23.5 0 26 41 41 24 6.50 Ce 0.15 0.32 0.27 0.044 0.74 3.46 34.7 0 34 53 52 43 8.82 V 4 50 0.02 0.55 50 35 0 0 91 63 47 54 68.57 Cr 1350.150.56106002316152415.68 Ni 8130.081.610100032221.59 Zn 0.9200.370.075120006789706533.41 11.68 0.00 55.17 7.00 1.71 1.32 0.26 22.86 Weight% used in Model 1 11.68 2.00 55.17 5.00 1.71 1.32 0.26 22.86 Weight% used in Model 2

The predicted trace element contents (model 1) generally compare well with the actual granitoid composition, with the exception of La, Y and Zn. As discussed earlier, the presence of monazite and xenotime in the Hillgrove Suite granitoids, and its likely presence in the source region, will strongly control the behaviour of the rare-earth elements and Y. As a result, these elements should be regarded as major elements in these phases, and hence are not governed by the trace element equations above. The over-prediction of Zn contents in the melt suggest that biotite may have constituted a minor component of the residual mineralogy. To test this possibility, a second model was attempted, involving 2% biotite as a residual phase. Although the predicted value for Zn in model 2 more closely approximates the observed value, the predicted value for Rb also decreases, further removing it from the actual value. This suggests that if biotite forms part of the residual mineral assemblage, it can only occur as accessory quantities. 141

4A5A4 Contamination within the Hillgrove Supersuite - the possibilities.

(I) The Rockisle Suite - contaminated Hillgrove Suite granitoids?

The Rockisle Suite is isotopically distinct from the Hillgrove Suite (Fig. 4A15), and also has a distinctive chemistry, as detailed above. Samples from the suite plot on an isotopic mixing curve between uncontaminated Hillgrove Suite samples and the gabbro analysed from the Bakers Creek Suite (Fig. 4A15). Isotopic mixing calculations show that this isotopic signature can be reproduced by mixing ‘uncontaminated’ Hillgrove Suite granitoids with the gabbro. The calculated mixing curve depicted in figure 4A15 is graduated in steps of 10% between the gabbro and ‘uncontaminated’ Hillgrove Suite (sample A7). Although this 60:40 gabbro/granite mix is difficult to reproduce by mixing Hillgrove Suite averages with a gabbro of ~50% SiO2, such a mechanism of direct mixing is unlikely, given the likely viscosity contrasts. Rather, mixing is more likely to occur between Hillgrove Suite magmas, with more chemically evolved members of the Bakers Creek Suite (e.g. quartz diorites or tonalites), which share much of the primitive isotopic nature of the gabbros.

Major and trace element contents of the Rockisle Suite provide supporting evidence that the Bakers Creeks Suite gabbros are the major contaminant involved in production of this suite. Contamination vectors (large arrow) shown on all Harker plot insets (Figs. 4A7, 4A8) depict the major trends for the Rockisle Suite. The majority of these trends are distinct for those of the Hillgrove Suite, and in all cases, define a vector which is directed at the field for Bakers Creek Suite gabbros. Hence although modelling of this mixing process is difficult since the precise composition of the end-members are unknown, this graphical evidence supports the isotopic data.

In addition to the contamination trend with the Bakers Creek Suite, several samples show a trend towards higher gNd and/or higher initial 87Sr/86Sr (contamination trend 2 in figure 4A15), suggesting that a secondary source of contamination was involved in magma generation. Another important component of the accretion complex sequences of the southern New England Fold Belt are ocean floor basalts (both MORB and OIB - Cawood 1984). Although no samples of metabasalt from the southern New England Fold Belt have yet been analysed for Nd and Sr isotopes, a field for the Lachlan Fold Belt Cambrian tholeiitic greenstones 142

(Nelson et al. 1984) is plotted on figure 1. The contamination trend exhibited by the Hillgrove Suite granitoids towards higher gNd and higher initial 87Sr/86Sr, suggests that similar altered seafloor basalts may have played some role in the generation of the granitoid magmas. A possible mechanism for this contamination may be small degrees of partial melting of amphibolite to yield tonalitic magmas (e.g. Rushmer 1991), which subsequently mix with the more abundant partial melts from metasedimentary sources.

(ii) Concomitant contamination processes within the Bakers Creek Suite

The contamination of Hillgrove Suite magmas by Bakers Creek Suite magmas detailed above, suggests that mixing may have also occurred at the opposing end of the spectrum. Therefore, contamination of gabbroic magmas by granitic material should be observed within the Bakers Creek Suite. Contamination of the Hillgrove Suite magmas by these mantle derived magmas was not considered by Hensel (1982), since he regarded the Bakers Creek Suite gabbros as younger (early Permian), and therefore they could not be genetically linked. However, contamination of the gabbros by accretion complex metasediments was implicated, and was detailed in a Sr isotopic study of the gabbros and more evolved diorites of the suite (Hensel 1982). Hensel (1982), concluded that the ‘calc-alkaline’ diorites of the suite were the result of contamination of the ‘tholeiitic’ gabbros by ‘felsic’ greywackes, similar to those involved in Hillgrove Suite generation. This conclusion was largely based on the Sr isotopic data, with most gabbros having initial 87Sr/86Sr ratios (0A7025-0A7030) similar to the sample analysed in this study, while the diorites spanned a larger range in initial 87Sr/86Sr (0A7030- 0A7045).

Major and trace element contents also support contamination of the gabbros by assimilation of fertile metasedimentary materials. The range of compositions of diorites on Harker plots (insets on Fig. 4A7, 4A8) cannot be explained by fractional crystallization processes alone, as discussed earlier. Rather, the array of compositions must be explained by a combination of fractional crystallization and assimilation. The field for diorites on Harker plots also lends support to earlier discussions of melt fertility, as the secondary vectors defining these arrays always implicate the fertile greywacke source composition as the major contaminant. Further discussion of this contamination process will not be presented here, since it has already been discussed at length by Hensel (1982). 143

4A6 Conclusions

Geochemical characteristics, combined with Sr and Nd isotopic compositions of granitoids and various potential source rocks, provide tight constraints on possible source-rocks for Hillgrove Supersuite granitoids. Only volcanogenic greywackes of intermediate composition

(~65% SiO2) overlap with the isotopic composition of the granitoids. Calculated melt fertilities of various potential source rocks, based on the proportional component of the ternary Q-Ab-Or minimum melt composition at 5 kb, also indicate that these intermediate greywackes are the most likely sources to produce large volumes of partial melt. Significantly, the isotopic characteristics and calculated melt fertilities preclude the involvement of pelites and felsic greywackes (~70%SiO2), which have previously been inferred as granitoid sources. The isotopic and chemical immaturity of these sediments

87 86 ( Sr/ Sr 0A7048 to 0A7070, gNd +2 to -1, high Na2O and low ASI), explains the unusual character of Hillgrove Supersuite granitoids, which are isotopically primitive (87Sr/86Sr 0A7040 to 0A7065, gNd -1 to +4 ), only mildly peraluminous (ASI 1A00 - 1A15), and relatively high in Na2O (3 - 4%) compared to most S-types. Major and trace element modelling indicate that the more mafic magmas (68-70% SiO2) of the suite were produced by ~48% partial melting of the intermediate greywacke source, under water-undersaturated conditions involving biotite breakdown at granulite facies conditions and mid crustal depths (~5 kb).

Isotopic and chemical variability within the Hillgrove Supersuite demands that two additional sources have contributed to magma formation. The more isotopically and chemically primitive granites of the Rockisle Suite (87Sr/86Sr 0A7040, gNd +4.0), which form ~5% of the Hillgrove Supersuite, have a bulk chemistry which deviates from the main Hillgrove Suite, with higher CaO, Al2O3, TiO2 and lower K2O, ΣFeO contents. These granites plot on an isotopic mixing curve between intermediate greywacke (87Sr/86Sr 0A7048 to 0A7070, gNd +2 to -1) and coeval gabbros of the Bakers Creek Suite (87Sr/86Sr 0A7027, gNd +9.5). Accordingly, mantle-derived magmas are considered to have been a contributor to the most primitive granites. Another possible minor contaminating source are seawater-altered metabasalts, which are common in the deeper parts of the accretion complex. These metabasalts are likely to have undergone small degrees of partial melting (via amphibole breakdown), contributing a minor melt component to the primary S-types, causing an isotopic shift towards higher gNd and higher initial 87Sr/86Sr. 144

CHAPTER 5. PETROGENESIS OF THE CHAELUNDI COMPLEX A-TYPE GRANITOID SUITE: DERIVATION BY PARTIAL MELTING OF A DEHYDRATED CHARNOCKITIC LOWER CRUST.

5A1 Introduction

The term ‘A-type granite’ was introduced by Loiselle and Wones (1979) for granites which are alkaline and ‘anhydrous’ in composition, although this classification, unlike the I- and S- type classification (Chappell & White 1974), does not imply a specific source rock type. In general, A-type granites are geochemically characterized by high contents of large ion lithophile elements (LILE), high field strength elements (HFSE) and rare-earth elements (REE), together with low contents of CaO, MgO and most compatible trace elements. A-type granites are also easily distinguished in hand specimen, since they are also mineralogically and texturally distinct from other granite types. Compared to I-type and S-type granites they are feldspar-rich, generally leucocratic, contain hastingsitic hornblende, occasional fayalite, and late-crystallizing annitic biotite, which is often interstitial to all other minerals. They also contain abundant accessory minerals such as allanite, zircon, fluorite.

Several petrogenetic schemes have been proposed for the origin of A-type granites, including: (1) fractionation products of alkaline basalts (Loiselle & Wones 1979, Turner et al. 1992); (2) melting of lower crustal source rocks under fluxing of mantle derived volatiles (Bailey 1978); (3) low degrees of partial melting of F and/or Cl enriched dry, granulitic residue from which a granitoid melt was previously extracted (Collins et al. 1982, Clemens et al. 1986) and (4) melting of a tonalitic I-type granite (Cullers et al. 1981, Creaser et al. 1991, Skjerlie & Johnston 1993). Recently, the last model has gained considerable favour (e.g. Skjerlie & Johnston 1993, Pitcher 1993), which casts doubt on the integrity of A-types as a group distinguishable from I-types (Creaser et al. 1991, p.166). However, the distinctive features of euhedral orthoclase and interstitial biotite allow unequivocal field identification of A-types, which, combined with their characteristic geochemical features (e.g. Collins et al. 1982 Whalen et al. 1987), indicates that they should be regarded as a separate granite type. The Chaelundi example, described below, demonstrates the diagnostic petrographic 145 and geochemical character of an A-type granite suite intimately associated with an I-type suite, and places considerable doubt on the validity of the tonalite source model, particularly in ‘post-collisional’ plate margin tectonic environments where granitic continental basement is lacking.

The A-type suite from the Chaelundi Complex, New England Batholith of eastern Australia, falls into the A-type subgroup with characteristics analogous to arc magmas, as defined by

Eby (1990). The suite is of particular significance, as it has an extended SiO2 range (66-76%) with the mafic end-member being similar in composition to the I-type granite which it intruded. Comparison of these two granite suites with charnockitic (C-type) magmas (Kilpatrick & Ellis 1992), provides significant insights into the origin of A-type granites. We propose a petrogenetic model involving partial melting of a mafic-intermediate lower crust of similar composition to the I-type source, which was dehydrated, but not melt depleted, at the time of production of the slightly earlier I-type granites, leaving alkali feldspar, rather than biotite, as the dominant K-bearing phase in the source region. Subsequent partial melting of this source, at elevated temperatures, produced a hot, dry magma of slightly different composition to the I-type magma, and these differences were magnified by fractional crystallization.

5A2 Geological Setting

The southern New England Fold Belt (Fig. 5A1a) consists of a Devonian-Carboniferous arc-forearc sequence (Tamworth Belt) in the west and an accretion-subduction assemblage (Tablelands Complex) in the east (Leitch 1974, Cawood & Leitch 1985). In the late Carboniferous, parts of the accretion-subduction complex underwent a rapid change from a low-T, high-P to a high-T, low-P metamorphic regime, accompanied by intrusion of S-type granites and minor mantle derived mafic magmas (Flood & Shaw 1977, Farrell 1988, Dirks et al. 1992). Metamorphism and granite emplacement occurred during a continued compressive stress regime, and represent a ‘stepping out’ of the subduction zone in an easterly direction, hence forming a new magmatic arc within the older subduction-accretion complex (Collins et al. 1993). The new arc underwent rifting in the early Permian, accompanied by basin formation and extrusion of mafic and felsic volcanic rocks (e.g. Leitch 1988). Rifting was followed by late Permian deformation which disrupted the terrane and 146

Leucoadamellites and related rocks

Map area Granodiorite / Adamellite Older granites and volcanics (>240Ma) of the New England Batholith New England Fold Belt

0 50 km

Great Australian Clarence- Basin Moreton Basin

30° 00' S

Peel -

Figure 5. 1b

southern New England Fold Belt Tablelands Complex

Manning

Tamworth Pacific

00' E Ocean Belt Sydney Fault 152° Basin System

Figure 5. 1a. Geological map of the southern New England Fold Belt showing the outcrop area of the New England Batholith, highlighting the units younger than 240 Ma. 147 caused uplift along major crustal sutures, forming the several fault-bound crustal blocks which now comprise the Tablelands Complex (Landenberger et al. 1995). Ensuing this deformation, the voluminous ‘post-tectonic’ I- and A-type granites of the New England Batholith intruded the Carboniferous subduction/accretion assemblage (Shaw & Flood 1981, Collins et al. 1993) during the Triassic.

In contrast to the southern New England Fold Belt, granitoid magmatism appears to have been semi-continuous throughout the northern New England Fold Belt from the Carboniferous through to the late Triassic, with subduction being the underlying cause (Gust et al. 1993). The Gympie Terrane which forms the eastern margin of the northern New England Fold Belt, records Permian arc-volcanism, dominated by basaltic and basaltic andesite tuff-breccias of island-arc-tholeiite affinity, along with subordinate dacitic tuffs (Sivell & Waterhouse 1988). The Gympie Terrane is considered to have formed as an outboard island-arc during the Permian and was subsequently accreted to the continental margin during the early Triassic (Coney et al. 1990). Although magmatic-arc activity is not recorded in the Permian of the southern New England Fold Belt, correlation of the Gympie Terrane with New Caledonia and the Nelson-Eglinton-Tikitimu region of New Zealand (Waterhouse & Sivell 1987), attest to an extensive volcanic arc situated to the east of the southern New England Fold Belt during the Permo-Triassic. Therefore, although the Triassic I- and A-type plutonism in the southern New England Fold Belt may not directly constitute a volcanic arc, it is likely to be ultimately subduction-related. This Triassic magmatism probably records the terminal stages of arc-related magmatic activity in the New England Fold Belt, after which Cretaceous calc-alkaline and Tertiary alkaline igneous activity heralded rifting and the opening of the Tasman Basin (Coney et al. 1990).

The Chaelundi Complex (Binns et al. 1967), a large (~110 km2), high-level intrusion in the Tablelands Complex (Leitch 1974), contains both I- and A-type granite suites that have indistinguishable Rb-Sr biotite-whole rock ages (233-235 Ma, Shaw & Flood 1993) and initial 87Sr/86Sr ratios (0A7042 - 0A7044, S.E. Shaw - unpublished data). The complex forms part of the most easterly and youngest group of intrusives of the New England Batholith (Fig. 5A1a), a belt dominated by leucoadamellites with A-type affinities, together with subordinate I-type granodiorites (Shaw & Flood 1981, Shaw & Flood 1993). 148

A-type suite Triassic

152° 20' E Granodiorite I-type Adamellite }

Carboniferous & Permian

metasediments River Dalmorton

30° 00' S Demon 130 0 5 km 131 132 Creek 133 129 57 134 64 135 Chandler’s 136 66 N Chaelundi 128 Mtn. kes

w Fault

18 100 12 119 10 9 20

8 106 101 7

System 114 Dundurrabin

103 Guy Ebor

Figure 5. 1b. Geological map of the Chaelundi Complex showing the three main phases of intrusion. Sample sites are indicated with numbers corresponding to sample numbers in Appendix E, with the exclusion of the ‘CC’ prefix. 149

The Chaelundi Complex has three main phases of intrusion (Fig. 5A1b); the first two are I-type hornblende-biotite adamellite and granodiorite, with the former distinguished by the presence of orthoclase phenocrysts. The youngest phase is A-type in character, with a felsic end-member which is similar to other Triassic leucoadamellites in the southern New England Fold Belt. However, unlike most A-types in the region, the Chaelundi Complex suite is compositionally extended, ranging from quartz monzonite to leucoadamellite. The pluton is regularly zoned from a ~1 km rim of quartz monzonite on the eastern and northern boundaries where the pluton contact dips shallowly (~40E) northward, to leucoadamellite at the core and southwestern boundary. Since the complex has been deeply weathered, sampling was restricted to forestry road cuttings and as a result two gaps in the silica range are apparent on the geochemical data plots (see later) for the A-type suite (at ~67-70% and

72-74% SiO2). These apparent gaps are considered to be an artefact of sampling rather than a reflection of any internal intrusive contacts within the pluton, because of the chemical continuity between the low- and high-SiO2 phases.

5A3 Petrography of the Suites

5A3A1 The I-type suite

The granodiorites are medium to light greenish-grey coloured rocks, with an even, medium grained granitic texture (Plate 5A1a), varying to porphyritic in the adamellite. Euhedral to subhedral hornblende (α=straw yellow, β=brownish green, γ=green, Mg# 55-45, Fig. 5A4) and plagioclase (An55-An20, Fig. 5A5) form the larger phenocrysts and were early crystallizing phases. Plagioclase phenocrysts are normally zoned, but regularly contain complexly zoned and/or mottled cores, and commonly contain abundant fine grained inclusions of hornblende. Subhedral - anhedral biotite (α=straw yellow, β=γ=dark reddish brown, Mg# 45-40, Fig. 5A3) is interstitial to plagioclase and hornblende. Quartz is generally interstitial to all minerals, but does occur as phenocrysts in the β-form in the more felsic adamellite varieties. Perthitic orthoclase is always interstitial in the granodiorites, but forms euhedral phenocrysts (Or65-

Or90, Fig 5A5) varying up to 20 mm in length in adamellite varieties, where it commonly contains inclusions of hornblende and biotite. Ilmenite and magnetite comprise the opaque mineralogy, accompanied by minor interstitial pyrrhotite and other sulfides. Accessory 150

Quartz

1 2

Alkali feldspar Plagioclase

Figure 5. 2. Streckheisen (1973) plot of representative modes from the two suites of the Chaelundi Complex. Symbols: Squares = I-type suite, Triangles = A-type suite. Fields are: 1 = adamellite, 2 = granodiorite, 3 = quartz monzonite. 151

Plate 5A1. Photomicrographs comparing the microstructure of the mafic end-members of the two suites: (a) I-type suite granodiorite in plane polarized light, (b) As (a) with crossed polars, (c) A-type suite - quartz monzonite in plane polarized light, (d) As (c) with crossed polars. Base of photographs = 5 mm. 03

(a) (b) 14

13 152

3 Siderophyllite

Bundarra

IV Al Hillgrove Chaelundi I-types

Annite Phlogopite 2 0 Mg/(Fe+Mg) 1 0.6

0.5

Ti 0.4

0.3

AlVI 0.2 0 1 2 Figure 5. 3. Mg/(Mg+Fe) vs AlVI and AlVI vs Ti plots for biotite compositions from the Chaelundi A- type suite. The compositions of biotite from the Chaelundi I-type suite are also included (Landenberger 1988). Biotite compositions from the Hillgrove (green) and Bundarra (purple) S-type suites are included for comparison. 153 1.0 Tremolite Tr. Hbl. . (ANa +A K )<0 5

Magnesio- Tsch Act Hornblende Tschermakite Actinolite Hbl Hbl

2+ I-types

Fe- Fe- Ferro-

Mg/(Mg+Fe ) Ferro- Act Tsch Actinolite Ferro-Hornblende Tschermakite Hbl Hbl

0.0 1.0 Mg (ANa +A K )>0.5 Fe3 >Al Magnesio- Has Silicic Ed Hastingsite Edenite Hbl Edenite Hbl 2+

I-types Mga Magnesian Has Hastingsite Hbl Fe Mg/(Mg+Fe ) Silicic Ferro- Ed Ferro-Edenite Edenite Has Hbl Hastingsite Hbl 0.0 8.0 7.5 7.0TSi 6.5 6.0 5.5

Figure 5. 4. Amphibole compositions for the Chaelundi A-type suite. Fields for the I-type suite are taken from Landenberger (1988). Amphibole classification scheme after Hawthorn (1981). 154

Ab Or

Figure 5. 5. Feldspar compositions from Chaelundi A-type granitoids. The pink shaded areas represent feldspar compositions from the I-type suite (Landenberger 1988). 155 minerals include sphene, apatite and zircon in decreasing order of abundance. Microgranitoid enclaves, varying from 10 mm to 100 mm in diameter, are common. These enclaves bear essentially the same mineralogy as their host, but are always more mafic. Pegmatitic and aplitic phases are rare, but mariolitic cavities are quite common, which together with the presence of β-form quartz, suggests a subvolcanic emplacement.

5A3A2 The A-type suite

The A-type granites are leucocratic, feldspar-rich rocks (Fig. 5A2) with a coarse- to medium- grained granitic texture. The mafic end-member (quartz monzonite - Plate 5A1b), despite having a lower silica content than the I-type suite, is considerably lighter in colour, due to the higher feldspar and lower ferromagnesian mineral contents (Table 1). Although A-type granites with such low silica contents are relatively rare, cumulate textures are not evident in these quartz monzonites (Plate 5A1b) and so a cumulate origin for these rocks is unlikely. A common phenocryst phase in the quartz monzonite is Fe-rich hypersthene (Mg# 35), which is commonly altered to fibrous aggregates of cummingtonite, but the rock is dominated by phenocrysts of perthitic orthoclase (Or70 - Or75, Fig. 5A5) and normally zoned plagioclase

(An30 - An5, Fig. 5A5). These feldspars are free of any inclusions, and plagioclase ‘cores’ are very rare, contrasting with those in the I-type suite. Quartz occurs both as β-form phenocrysts and as interstitial grains. Ferro-hornblende (α=pale straw yellow, β=pale brownish green, γ=olive green, Mg# 35, Fig. 5A4) is typically interstitial to both feldspars and quartz, but less commonly occurs as subhedral crystals. Biotite (α=pale straw to colourless, β=γ=brown, Mg# 30, Fig. 5A3), commonly occurs interstitially, but also forms minor subhedral ragged flakes. Ilmenite is the main opaque phase, along with subordinate late pyrrhotite. Accessory minerals include zircon, apatite and allanite in decreasing order of abundance.

In contrast to the I-type suite, aplite and microgranite dykes are common, and microgranitoid enclaves are absent. However, occasional small mafic ‘clots’ or single crystals of augite, hypersthene and calcic plagioclase do occur in the quartz monzonites, displaying gross disequilibrium textures, with augite rimmed by secondary actinolitic hornblende, hypersthene rimmed by cummingtonite, and calcic plagioclase rimmed by sodic plagioclase. 156

The felsic end member of the suite (leucoadamellite) is also dominated by feldspars and quartz, which comprise more than 95% of the rock. Plagioclase is normally zoned, but more sodic (An10 - An5, Fig 5A5), and orthoclase more potassic, than in the quartz monzonite. Quartz commonly forms aggregates of large β-form phenocrysts. Interstitial annitic biotite (α=colourless, β=γ=dark brown, Mg# 20, Fig 5A3), and minor magnetite are the only mafic phases. Accessory phases are uncommon, except for F-rich apatite.

A-type Suite I-type Mafic Felsic Early crystallizing plagioclase, hornblende, plagioclase, quartz, plagioclase, quartz, phases biotite orthoclase, OPX orthoclase Late crystallizing quartz, orthoclase biotite, hornblende biotite phases Plagioclase An# range 55620 30651065 Mafic phases Mg# 55645 35630 20610 range Opaque phases present ilmenite, magnetite ilmenite magnetite Accessories (in order of sphene, apatite, zircon zircon, apatite, allanite fluor-apatite abundance) Microgranitoid >10cm common rare mafic ‘clots’ <5mm absent enclaves Leucogranite dykes rare common common

SiO2% range 68671 66676 Higher MgO, CaO, Sr, Higher K O Ba, Zr, Hf, Higher Rb, F, Pb, Th, Comparative chemistry 2 V, Cr, Ni, Fe3+/ΣFe LREE Nb, Y, Ga, HREE

Table 5A1. Comparative mineralogy and geochemistry of the I- and A-type suites.

5A4 Geochemistry

The first notable geochemical difference between the two suites is their compositional range: the granodiorites and adamellites of the I-type suite are restricted to the range 68% to 71%

SiO2, whereas the A-type suite varies from 66% to 76% SiO2 (Appendix E). At the low SiO2 end of this compositional spectrum, the two suites are indistinguishable in terms of Al2O3, total FeO, Rb, Pb, Ga, Nb, Y, Th, U, Cl and F (Figs. 5A6 & 5A7). However, the A-type suite has lower MgO, CaO, Sr, V, Cr, Ni and Fe3+/ΣFe, and higher total alkalis, Ba, Zr, Hf, 157

0.6 3.0 CaO TiO2 0.5

0.4 2.0

0.3

0.2 1.0

0.1

16 Al O 4.6 2 3 Na2 O 4.4 15 4.2

4.0 14 3.8

13 3.6

EFeO KO2 3.0 4.5

2.0 4.0

1.0 3.5

1.4 3+ E MgO 0.4 Fe / Fe 1.2

1.0 0.3 0.8 0.6 0.2 0.4 0.1 0.2

66 68 70 72 74 76 66 68 70 72 74 76 SiO2 SiO2

Figure 5. 6. Selected major elements plots. Symbols as for figure 5. 2. 158

Ba 300 1000 Sr 250 800 200 600 150 400 100

200 50

Rb 400 400 Zr

300 300

200 200 100

2000 60 Pb F

1500 40 1000

20 500

70 40 Th Zn 60 30 50 20 40

10 30

66 68 70 72 74 76 66 68 70 72 74 76 SiO2 SiO2

Figure 5. 7. Selected trace elements plots. Symbols as for figure 5. 2. 159

Zn, and LREE (Figs. 5A6, 5A7 & 5A8a). The REE patterns of the low SiO2 members of each suite are remarkably similar (Fig. 5A8a): they are relatively unfractionated and show minor negative Eu anomalies of 0A72 and 0A83 for the I- and A-type suites, respectively, with the latter distinguished by slight LREE enrichment (La/LuN = 9A22) relative to the I-type (La/LuN = 3A97).

The more felsic rocks of the A-type suite (leucoadamellite and microgranite dykes) show extreme depletion of all compatible elements, including TiO2, Al2O3, FeO, CaO, MgO, Ba, Sr, Zr and the LREE. On the other hand Rb, F, Pb, U, Th, Nb, Ga, Y, and HREE all increase substantially, producing the characteristic chemistry of the New England leucoadamellites. Ferric iron also increases, resulting in the leucoadamellites having high Fe3+/ΣFe. Also, the characteristically high Ga/Al ratios of A-type granites (and of all strongly fractionated granites) is only evident in the more felsic rocks. The development of a large negative Eu anomaly (Eu/Eu* . 0A06), combined with the fall in LREE and increase in HREE, has produced a flat or ‘seagull’ REE pattern, which is characteristic of the New England A-type granites (Fig. 5A8b). The elevation of HREE contents, and the drop in LREE contents, is coincident with the disappearance of zircon and the appearance of allanite as accessory phases.

The geochemical variation for the A-type suite is reflected by changes in mineral chemistry.

Plagioclase in the quartz monzonites have core to rim variations of An40 to An10, whereas in the leucoadamellites, the compositional range is restricted to An10 - An5. Biotite varies widely in composition with MgO and TiO2 lower, and FeO and MnO contents higher in the more felsic rocks. Orthoclase, hornblende and ilmenite compositions remain relatively unchanged throughout the suite, although the latter two only occur in the more mafic rocks. Magnetite is restricted to the leucoadamellites, and its appearance causes a decrease in Fe/V ratios at ~74% SiO2 (Fig. 5A10g). In contrast, mineral compositions in the I-type suite exhibit little intra-suite variation, but tend to differ from those in the A-type suite, which reflects the whole-rock geochemical differences. For example, biotite and hornblende from the I-type suite have an Mg# values of 45-55, compared to 30-35 for the mafic A-types.

A-type granites have been subdivided into two groups on a geochemical basis; those associated with rift zone magmatism and those associated with the waning stages of arc 160

“Parent magmas” (a)

100 A-type suite

I-type suite

10 1982). al. et Rock / Chondrite

1 Model 1 predicted (b) Gabo 100 Mumbulla

Model 2 predicted 10 Quartz monzonite A-type suite Rock / Chondrite Granite Leucogranite

1 (c) I-type suite 100

10 Granodiorite Rock / Chondrite Granite Predicted . REE - chondrite normalized (Nakamura 1974) plots of: plots 1974) (Nakamura normalized chondrite 5- REE 8. 1 La Eu Tb Ho Lu Figure (a) ‘Parent’ magmas for both suites. both for magmas ‘Parent’ (a) (b) Intra-suite variation for the A-type suite. The dashed line represents the composition of the leucoadamellite as predicted by REE modelling. REE by predicted as leucoadamellite fields the Shaded of composition suite. the A-type represents line (b) dashed The the for variation Intra-suite represent the ranges of composition observed for the Gabo Suite (darker shading) and Mumbulla Suite (lighter shading) (Collins shading) (lighter Suite Mumbulla and shading) (darker Suite Gabo the for observed composition of ranges the represent Ce Nd Sm Yb modelling. REE by predicted as end-member felsic the of composition the represents line dashed The suite. I-type the for variation Intra-suite (c) 161 magmatism (Eby 1990). The tectonic discrimination diagrams used by Pearce et al. (1984) and Eby (1990) to discriminate tectonic setting (Fig. 5A11 & 5A12) imply an arc setting for both the I-type suite and most members of the A-type suite. Subduction-related granitoid magmatism appears to have been semi-continuous throughout the northern New England Fold Belt from the Carboniferous through to the Triassic (Gust et al. 1993), and Triassic I- and A-type plutonism in the southern New England Fold Belt probably records the terminal stages of this arc-related magmatic activity.

The more felsic members of the A-type suite plot in the within plate granite field (Fig. 5A11) or well outside the volcanic arc field and shifted towards the field of OIB (Fig. 5A12). The distribution of the A-type suite data over two fields on the Y-Nb tectonic discrimination diagram demonstrates one hazard of using these diagrams. Pearce et al. (1984) stated that the diagrams should not be used for rocks which show evidence of crystal accumulation. Similarly, they should not be used for rocks which have undergone significant magmatic differentiation.

5A5 Petrogenetic Constraints

5A5A1 Fractionation

Major-element modelling

Mass-balance calculations using least-squares-mixing (Le Maitre 1981) were applied to each suite to determine if the chemical variation for the analysed samples could be ascribed to fractional crystallization, to determine the minerals involved, and to quantify this fractionation. Although geochemical modelling of fractionation processes using this method and trace element models (see next section) has limitations, the consistency of modelling results using least-squares-mixing for major elements and results derived from Rayleigh fractionation models (using trace elements) show that these methods are applicable where fractionation has been a rapid and efficient process (as in the case of A-type magmas).

Initially, analyses of all minerals occurring in the mafic end-member for each suite were used in calculations to quantify the assumption that: 162

mafic end-member ! fractionating minerals = felsic end-member

Mineral compositions not used in the initial mixing calculations were then rejected.

Calculations for the I-type suite (Table 5A2) reveal that only a moderate degree of fractionation (~16%) is necessary to generate the restricted compositional range. Fractionation is dominated by removal of plagioclase and hornblende, with minor involvement of biotite and ilmenite.

Different mineral phases dominate the fractionation within the A-type suite, and is indicated by subdividing the mixing calculations into two separate stages (Table 5A2). Hornblende and biotite removal occurs early in the fractionation sequence along with plagioclase, while in the more felsic rocks removal of feldspar and quartz dominates. These calculations are consistent with petrographic observations, with hornblende occurring only in the quartz monzonite, and quartz becoming a major phenocryst phase approximately midway through the fractionation sequence. The requirement that biotite is a fractionating phase reflects the observation that this mineral does occur as minor discrete flakes in the mafic end-member, and only becomes exclusively interstitial in the more felsic rocks. Calculation of the total fractionation for the SiO2 range 66-76%, reveals a much greater degree of fractionation (~71%), dominated by removal of both plagioclase and orthoclase, quartz, biotite and hornblende, with minor amounts of ilmenite and ferrohypersthene.

The assumption that the mafic end-members of each suite approximate magma compositions rather than cumulates (particularly with reference to the quartz monzonites of the A-type suite) is justified by the absence of cumulate textures in any of these mafic granitoids, and the lack of positive Eu anomalies (see geochemistry section above) which would be present if these quartz monzonites were cumulates (considering that plagioclase is the major liquidus phase). It is also considered unlikely that the quartz monzonites are cumulates since their occurrence near the shallowly dipping northern margin of the intrusion (see earlier) implies that they were emplaced as part of the roof zone of the pluton. 163

Table 5A2 Major element modellingA A-type suite: (1) Fractionation from quartz monzonite (CC131) to adamellite (CC135) Reactant Analyses = 1 Product Analyses = 2 +3+4+5+6+7+8+9 123456789 Resultant compositions (normalized): Parent Daugher Horn- Plagio- Ortho- Ilmenite Quartz Biotite OPX Reactant Product Diff- Cumu- blende clase clase erence late

SiO2 66A75 71A50 45A68 60A39 63A69 0A50 100A00 35A29 49A94 67A84 67A84 0A00 62A39

TiO2 0A62 0A34 1A21 0A00 0A63 53A83 0A00 4A57 0A00 0A63 0A63 0A00 1A00

Al2O3 15A63 14A67 5A19 24A90 18A70 0A00 0A00 12A30 0A00 15A89 15A89 0A00 17A33 3FeO 3A67 2A12 22A85 0A00 0A00 43A44 0A00 25A99 34A43 3A73 3A73 0A00 5A81 MnO 0A08 0A05 0A51 0A00 0A00 4A43 0A00 0A27 1A04 0A08 0A08 0A00 0A12 MgO 1A05 0A48 8A68 0A00 0A00 0A00 0A00 7A25 13A69 11A07A 07 0A00 1A83 CaO 2A70 1A63 10A20 6A24 0A12 0A00 0A00 0A00 1A51 2A74 2A74 0A00 4A18

Na2O4A26 4A01 1A47 7A84 3A27 0A00 0A00 0A00 0A00 4A33 4A33 0A00 4A71

K2O3A63 4A47 0A58 0A23 12A36 0A00 0A00 9A24 0A00 3A69 3A69 0A00 2A63 % used 100A00 56A59 3A15 23A27 4A94 0A27 6A00 4A63 1A15 Residual sum of squares = 0A0000 Distance apart = 0A0007 Degree of fractionation (crystals removed) = 43A41%

(2) Fractionation from adamellite (CC135) to leucoadamellite (CC57) Reactant Analyses = 1 Product Analyses = 2 +3+4+5+6+7+8+9 1245678 Resultant compositions (normalized): Parent Daughter Plagio- Ortho- Ilmenite Quartz Biotite Reactant Product Diff- cumu- clase clase erence late

SiO2 71A50 75A69 60A39 64A97 0A50 100A00 35A32 72A03 72A04 -0A01 67A5

TiO2 0A34 0A10 0A00 0A00 53A83 0A00 3A39 0A34 0A34 0 0A6

Al2O3 14A67 13A00 24A90 18A40 0A00 0A00 12A54 14A78 14A8-0A02 16A6 3FeO 2A12 1A06 0A25 0A00 43A44 0A00 27A02 2A14 2A16 -0A02 3A31 MnO 0A05 0A04 0A00 0A00 4A43 0A00 0A53 0A05 0A06 -0A01 0A08 MgO 0A48 0A11 0A00 0A00 0A00 0A00 7A13 0A48 0A45 0A04 0A8 CaO 1A63 0A64 6A24 0A00 0A00 0A00 0A00 1A64 1A63 0A01 2A68

Na2O4A01 4A10 7A84 2A22 0A00 0A00 0A00 4A04 4A03 0A01 3A91

K2O4A47 4A45 0A23 13A85 0A00 0A00 9A32 4A54A504A52 % used 100 51A48 20A84 11A75 0A20 10A55 5A18 Residual sum of squares = 0A0028 Distance apart = 0A0533 Degree of fractionation (crystals removed) = 48A52%

Total fractionation from quartz monzonite to leucoadamellite: Parent Daughter Horn-Plagio- Ortho- Ilmenite Quartz Biotite OPX Proportions blende clase clase Used (%): 100 29A13 3A15 35A06 11A59 0A38 11A97 7A56 1A15

Total degree of fractionation for A-type suite (crystals removed) = 70A87%

Major element modelling of I-type suite: Reactant Analyses = 1 Product Analyses = 2+3+4+5+6 123456 Resultant compositions (normalized): Parent Daughter Biotite Horn- Ilmenite Plagio- Reactant Product Diff- Cumu- blende clase erence late

SiO2 67A56 70A84 35A85 45A70 0A33 58A05 68A76 68A74 0A02 52A03

TiO2 0A56 0A39 4A74 1A30 48A67 0A00 0A57 0A56 0A01 1A47

Al2O3 15A10 14A51 12A47 6A32 0A20 26A64 15A37 15A34 0A03 18A94 3FeO 3A25 2A34 21A71 17A77 43A80 0A00 3A31 3A29 0A02 8A41 MnO 0A07 0A05 0A33 0A61 5A85 0A00 0A07 0A09 -0A02 0A29 MgO 1A43 0A84 10A75 11A16 0A12 0A00 1A46 1A43 0A03 4A65 CaO 3A31 2A46 0A07 11A27 0A00 8A44 3A37 3A39 -0A02 8A39

Na2O3A85 3A79 0A35 1A56 0A19 6A87 3A92 3A94 -0A03 4A54

K2O3A13 3A53 9A68 0A59 0A00 0A00 3A19 3A23 -0A04 1A29 % used 100 84A79 4A63 4A48 0A16 8A89 Residual sum of squares = 0A0054 Distance apart = 0A0733 Degree of fractionation (crystals removed) = 15A21% 164

Trace-element modelling

The assumption that the major element modelling calculations (above) closely represent classical Rayleigh fractional crystallization (as opposed to equilibrium crystallization or other mixing/unmixing differentiation processes) may be tested using trace element models for fractional crystallization. Variation in selected trace elements (representing incompatible elements and elements compatible in the major crystallizing phases) was modelled using the equation for Rayleigh fractional crystallization:

i Where: C0 = concentration of element I in the original melt (parent) i C1 = concentration of element I in the remaining melt (daughter) f = fraction of remaining melt Di = bulk distribution coefficient for element I

Trace element concentrations were calculated for remaining melt fractions using the concentration in the mafic end-members of each suite together with the fraction of remaining melt and mineral proportions calculated by major element modelling. These results are compared to actual compositions for the felsic end-members of each suite in Table 5A3. The mineral proportions used in these calculations are as fractions of the total removed mineralogy (as required by the above equation), not as percentages of the parent melt as calculated in the major element modelling.

Whereas the calculated composition of residual magma of the I-type suite shows excellent agreement with the actual composition, several discrepancies occur between the calculated and actual compositions of the residual magma in the A-type suite. Calculations underestimate Rb and Zn contents, which may be due to an overestimation of the proportion of biotite removal calculated by major element modelling. Thus the abundance of Rb and Zn suggests that biotite played a much smaller role in the fractionation than major element modelling would suggest. Other discrepancies (e.g. Pb, Sr and Ba) may reflect inappropriate partition coefficients for these elements and the likely variation in these coefficients over such a large range of magma compositions (see discussion below). 165

Table 5A3. Trace element Modelling Partition coefficients used: Mineral OPX Hornblende Biotite Plag Orthoclase Ilmenite Quartz Bulk D Bulk D Ref** 111112 A-typeI-type Ba 0A003 0A044 9A70A308 6A12 - - 2A190 1A258 Rb 0A003 0A014 2A24 0A048 0A34 - - 0A319 0A278 Sr 0A009 0A022 0A12 4A43A87 - - 2A824 2A595 Pb - - 0A767 0A972 2A473 - - 0A967 0A653 Th - 0A11 0A997 - 0A02 - - 0A115 0A142 Y 0A960A03 0A06 0A006 2A09 - 0A326 1A827 V 61010- -35-1A797 4A411 Zn 0A9720----2A459 4A258 Residual Weight fraction used: liquid fraction: A-type 0A0162 0A0444 0A1067 0A4948 0A1635 0A0054 0A1690 0A2913 I-type - 0A2945 0A1098 0A5852 - 0A0105 0A8479 Weight fractions are as fractions of the total removed mineralogy

A-type suite results: Calculated Actual Concentration concentration of concentration of in parent * residual magma residual magma * Ba 1034 238 169 Rb 80 185 301 Sr 280 30 52 Pb 18 19 43 Th 92629 Y 29 68 92 V 47 18 13 Zn 65 11 43

I-type suite results: Calculated Actual Concentration concentration of concentration of in parent * residual magma residual magma (CC101) Ba 499 478 470 Rb 98 110 117 Sr 327 251 258 Pb 14 15 19 Th 11 13 13 Y 31 27 29 V 61 35 37 Zn 43 25 39

* Average values for the three most mafic or felsic samples used as appropriate ** References: (1) Arth 1976 (2) Nash & Crecraft 1985 (values for Zn taken from Henderson 1982)

Despite these inconsistencies, the close agreement of calculated and actual values for Th (which has the lowest bulk distribution coefficient Di in both suites) supports both a Rayleigh (disequilibrium) fractionation process and the degree of fractionation calculated by major element modelling. The proportion of remaining melt (f) may be approximated by using Th as an incompatible element and solving the above equation for f: 166

Estimation of f using actual values for parent and daughter magmas and the above equation, are 0A823 (c.f. 0A8479 for major element modelling) for the I-type suite and 0A267 (c.f. 0A2913 for major element modelling) for the A-type suite, which supports the degrees of fractionation calculated by major element modelling in these suites.

Rare-earth-element modelling

REE modelling (Table 5A4) was based on the mineral proportions calculated by major element modelling. Model compositions for the felsic end-members of both suites were calculated using partition coefficients for dacites and rhyolites and were then compared with actual compositions.

Calculations for the I-type suite (Table 5A4, Fig. 5A8c) agree well with the observed intrasuite variation. Removal of the two major fractionating minerals, hornblende and plagioclase accounts for the observed patterns, with hornblende removal causing depletion in the HREE

(an increase in the La/LuN ratio from 3A97 to 5A26) and plagioclase removal producing the small negative Eu anomaly. Although the size of the predicted Eu anomaly (Eu/Eu* . 0A45) exceeds that observed (Eu/Eu* . 0A60), this calculation is very sensitive to value used for the partition coefficient which is used for Eu in plagioclase (see below).

Fractionation in the A-type suite has produced a ‘seagull’ REE pattern, with a very large negative Eu anomaly (Eu/Eu* . 0A06) and a greatly reduced La/LuN ratio (.1A06) (Table 5A4, Fig 5A8b). This ‘seagull’ REE pattern cannot be adequately explained using only the mineral proportions calculated by major element modelling (model 1 - Table 5A4, Fig. 5A8b) as it produces a distinct rise in the LREE to ~120 × chondrite. The marked fall in LREE contents contrasts with the model 1 predictions (see Fig. 5A8b), and occurs approximately midway through the fractionation sequence (~71% SiO2). This coincides with the appearance of allanite as an accessory phase. The accompanying rise in HREE contents matches the disappearance of hornblende and zircon. This fall in LREE can only be 167

Table 5A4. Rare-earth-element Modelling

Partition coefficients used: Mineral OPX Hblnd Biotite Plag Orthoclas Ilmenite Allanite Quartz Bulk D - A-type Bulk D RefA**1112324 Model1Model2I-type La 0A11 0A85 0A33 0A38 0A072 3A6 820 - 0A294 1A934 0A510 Ce 0A15 0A899 0A32 0A267 0A046 3A46 635 - 0A235 1A505 0A457 Nd 0A22 2A89 0A29 0A203 0A038 3A22 463 - 0A287 1A213 1A002 Sm 0A27 3A99 0A26 0A165 0A025 2A83 205 - 0A310 0A720 1A300 Eu 0A17 3A44 0A24 5A417 2A60A55 81 - 3A289 3A451 4A237 Tb 0A460A28 0A08 0A033 2A47 71 - 0A361 0A503 1A845 Ho 0A55 6 0A32 0A06 0A006 2A09 50 - 0A351 0A451 1A837 Yb 0A86 4A89 0A44 0A09 0A015 1A64 8A9-0A334 0A352 1A541 Lu 0A94A53 0A46 0A092 0A031 1A14 7A7-0A322 0A337 1A439

Weight fraction used: Residual liquid A-type: fraction: Model 1 0A0162 0A0444 0A1067 0A4948 0A1635 0A0054 0A0000 0A1690 0A2913 Model 2 0A0162 0A0444 0A1067 0A4948 0A1635 0A0054 0A0020 0A1690 0A2913 I-type: 0A2945 0A1098 0A5852 0A0105 0A8479 Weight fractions are as fractions of the total removed mineralogy

A-type Suite results: Calculated Calculated Actual Concentration concentration concentration concentration in parent (CC132) Model 1 Model 2 in residual magma (CC57) La 36A70 87A67 11A60 14A40 Ce 77A60 199A41 41A63 41A50 Nd 36A90 88A92 28A38 24A80 Sm 6A34 14A84 8A95 7A82 Eu 1A67 0A10 0A08 0A18 Tb 0A77 1A69 1A42 1A78 Ho 0A98 2A18 1A93 2A84 Yb 2A63 5A98 5A85 8A87 Lu 0A41 0A95 0A93 1A40 Eu/Eu* 0A8350 0A0214 0A0271 0A0622

La/LuN 9A2233 9A5423 1A2865 1A0598

I-type Suite results: Calculated Actual Concentration concentration concentration in parent (CC7) in residual liquid in residual magma (CC101) La 18A50 20A06 19A40 Ce 41A80 45A72 40A00 Nd 19A70 19A69 20A50 Sm 4A48 4A26 4A27 Eu 1A11 0A65 0A86 Tb 0A81 0A70 0A67 Ho 1A24 1A08 0A97 Yb 3A27 2A99 2A56 Lu 0A48 0A45 0A38 Eu/Eu* 0A7160 0A4515 0A6035

La/LuN 3A9713 4A6292 5A2604

References (**): (1) Arth 1976 (2) Nash & Crecraft 1985 (3) Mahood & Hildreth 1983 (4) Brooks et al. 1981 10 30 20 40 1000 50 Plagioclase 10 500 20 10 Biotite Ba 200 Orthoclase

100 20

50 0.01 0.1 1.0 Eu/Eu*

Figure 5. 9. Plot illustrating the combined effect of plagioclase and orthoclase fractionation on Ba and Eu contents within the A-type suite (dashed line). The scales on each trend represent the percentage of the phase removed from the melt (calculated using partition coefficients from Arth (1976), Mahood & Hildreth (1983), & Nash & Crecraft (1985)). The star symbol represents the composition of the leucoadamellite as predicted by trace element and REE modelling. Symbols as for figure 5. 2. 168 169 reproduced by modelling if allanite is included in the calculation (model 2 - Table 5A4, Fig. 5A8b). Although an arbitrary amount (0A2%) of allanite has been used in the calculation for model 2, this modelling demonstrates that allanite must be used to explain the observed REE patterns of the fractionated A-types. The large negative Eu anomaly which characterizes the leucoadamellites resulted from both plagioclase and orthoclase fractionation (Fig. 5A9). The star symbol on this diagram is the predicted composition of the leucoadamellite and the dashed line is a best fit representation of the observed Ba-Eu*/Eu variations, demonstrating the increasing involvement of orthoclase in fractionation as the suite becomes more felsic. The predicted Eu anomaly exceeds that observed, as in the I-type suite. The value for the

Eu partition coefficient used in the calculations (DP lag = 5A42, Nash & Crecraft 1985) is that published for high-silica rhyolites. The use of other published partition coefficients such as

Eu for dacites (DP lag = 2A11, Arth 1976) results in a large shortfall in the predicted anomaly. REE modelling for such a compositionally extended suite introduces the problem of variation of partition coefficients within the suite due to wide range of silica contents. As magmas become increasingly silicic, and strongly polymerized, melt compositions and structure become more ‘feldspar-like’ and hence become increasingly exclusory to all trace elements, resulting in much larger partition coefficients (Mahood & Hildreth 1983). This effect is also observed for the many other trace elements within the suite: for example, although the degree of plagioclase removal decreases in the more felsic rocks of the A-type suite, Ca/Sr ratios (Fig. 5A8a) have the same rapid increase as K/Ba ratios. This effect may be partly due to partitioning of Sr in K-feldspar, but also reflects the likely strong increase in partition coefficients for compatible elements (e.g. Ba, Sr) previously documented in many high-SiO2 suites (e.g. Mahood & Hildreth 1983).

General implications of feldspar fractionation in A-type granites.

Although the quartz monzonites of the Chaelundi Complex exhibit many of the characteristic geochemical signatures of A-type granites, several properties regarded as distinctively A-type are only really evident as a result of extensive fractional crystallization. Ga, Rb, Nb, Y, HREE, Pb, Th, U, F and Fe3+/ΣFe all increase with fractionation due to the incompatible nature of these elements in A-type magmas (Fig. 5A6, 5A7 & 5A8b). Concentrations of these elements in the parent magma are the same as, if not lower than levels in the I-type suite. This suggests that fractionation may well be the dominant process in the generation of high 170 levels of these incompatible elements in some A-type magmas.

The feldspar controlled fractionation within the suite also markedly alters inter-element ratios (Fig. 5A10). Plagioclase fractionation has produced a characteristic pattern for Ca/Sr and

Rb/Sr, involving a slight increase in the ratios from 66-76% SiO2 followed by a more rapid increase only at the high-SiO2 extreme. On the other hand, orthoclase produces a more pronounced increase at around 74% SiO2, as reflected in changing K/Ba, Rb/Ba and Ga/Al. The decrease in K/Rb mirrors the increase in Rb/Ba.

A comparison with the Gabo and Mumbulla suites of the Lachlan Fold Belt, southeastern Australia (Collins et al. 1982) reveals interesting similarities and differences (Fig. 5A10).

Rb/Sr and K/Rb ratios are similar to the Chaelundi Complex granites at given SiO2 contents, indicating that the Gabo and Mumbulla suites were derived by feldspar fractionation of a more mafic parent. The lower Rb/Ba and K/Ba ratios of Mumbulla imply lesser degrees of fractionation than is required for the high-SiO2, Chaelundi A-types. However, contrasting Ca/Sr variation of the Gabo and Mumbulla suites, relative to the Chaelundi fractionation trends, is more suggestive of varying degrees of partial melting. Furthermore, the Ga/Al ratios of Gabo and Mumbulla are uniformly high, at the Chaelundi extreme of the feldspar fractionation trend, which is not indicated by the other ratios. Thus, for the Lachlan Fold Belt suites, higher Ga/Al appears to have resulted from lower degrees of partial melting compared with the Chaelundi A-type suite, which is consistent with their elevated REE contents (Fig. 5A8b).

Similar differences also occur between leucoadamellites within the New England Batholith, many of which have similar SiO2 contents to the leucoadamellites of the Chaelundi Complex, but which commonly have higher K/Rb and lower Rb/Sr ratios (Landenberger, unpublished data). This implies that different degrees of partial melting and later fractionation apply to each pluton, and therefore each should be treated individually in petrogenetic considerations. This effect, along with the extreme degree of fractionation possible in A-type melts, may partly explain the large variety of compositions of A-type granites in arc-related environments (Eby 1990). Classification of the New England leucoadamellites as A-type is ambiguous if they are highly fractionated, due to the inherent difficulties of applying a genetic classification to high silica granites. However, wherever more mafic rocks occur within these plutons they 171

(a) (b) 300 500 Ca/Sr K/Ba 300 200

100 100

50 50 30 30 20

100 2.0 1.8

(e) (f) 100 Rb/Ba Rb/Sr 1.0 10

0.1 1

(g) (h) 5000 0.08 Fe/V Mn/Fe 2000 0.06 1000 0.04 500

0.02 200

66 68 70 72 74 76 66 68 70 72 74 76 SiO2 SiO2 Figure 5. 10. Selected plots of inter-element ratios vs silica. Symbols as for figure 5. 2 with the addition of average for C-type magmas at 67% SiO2 - red star (Kilpatrick & Ellis 1992). Shaded fields represent the ranges of composition observed for the Gabo Suite (blue) and Mumbulla Suite (green) of the Lachlan Fold Belt (Collins et al. 1982). 172 always share the distinctive A-type characteristics (e.g. high LREE and HFSE) of the quartz monzonites of the Chaelundi Complex (Landenberger, unpublished data). This observation suggests that most leucoadamellites in New England are A-type in character.

5A5A2 Parental magmas

Major, trace and REE modelling indicate that the most mafic rocks analysed in the I- and A- type suites may be parental magmas, but the possibility that they represent fractionated magmas needs to be assessed. Petrography indicates plagioclase was a liquidus phase in mafic rocks of both suites, so any prior fractionation would have substantially increased Ca/Sr and Rb/Sr ratios. However, the Ca/Sr and Rb/Sr ratios for both suites are similar and very low (Rb/Sr < 0A4, Ca/Sr < 70, Fig. 5A10) indicating that earlier removal of plagioclase was very limited. Orthoclase also occurs as a near-liquidus phase early in the crystallization history of the A-type suite, and its removal by fractionation would have considerably increased Rb/Ba and K/Ba ratios. These ratios are also extremely low (Rb/Ba < 0A1, K/Ba < 35, Fig. 5A10) in the quartz monzonites, precluding prior removal of orthoclase, and also suggesting minimal fractionation. Furthermore, fractionation of feldspars would also have produced negative Eu anomalies, but these are not significant in either suite (Eu/Eu* = 0A83 for mafic A-type granite, Eu/Eu* = 0A72 for mafic I-type granite, Fig. 5A8a). The relatively flat, ‘primitive’ REE patterns (Fig. 5A8a) together with the evidence above, indicate that the mafic end-members of both suites have not undergone significant differentiation, and can be considered as parental magmas.

5A5A3 Geochemical contrasts between the Chaelundi Complex I- and A-type parental magmas.

The general geochemical similarities of the two granitoid suites of the Chaelundi Complex, and their similar initial Sr isotopic composition, suggest derivation by partial melting of similar source rocks. Furthermore, similar incompatible element contents (e.g. Rb, Pb, Th;

Fig. 5A7) and in particular, the similar Rb/Sr and Ca/Sr ratios at 66-67% SiO2 (Fig. 5A10 a,e) suggest that similar degrees of partial melting occurred during production of both suites. Nonetheless, the A-type parental magma has elevated contents of HFSE, LILE and LREE, and lower contents of most compatible elements compared with the I-type magma. In 173 1000

WPG

100 Fractionated A-types

Nb VAG 10 + syn-COLG ORG

10 100 1000 Y Figure 5. 11. Y vs Nb tectonic discrimination diagram for both suites (Pearce et al. 1984). Fields: WPG = within plate granites, VAG = volcanic arc granites, syn-COLG = syn-collision granites, ORG = oceanic ridge granites. Symbols as for figure 5. 2.

IAB 10.0 CG

5.0 VAG Ce/Nb

2.0 OIB MORB

1.0 Fractionated A-types

0.1 1 10 Y/Nb Figure 5. 12. Ce/Nb vs Y/Nb discrimination plot for the two major subgroups of A-type granites (Eby 1990). Fields: IAB = island arc basalts, OIB = oceanic island basalts, MORB = mid-ocean ridge basalts, CG = syn-collision granites, VAG = volcanic arc granites. Symbols as for figure 5. 2. 174 particular, the relatively low K/Ba and Rb/Ba ratios, and high K/Rb ratios of the A-type parent magma need to be explained by any petrogenetic model.

The bulk differences in major element compositions of the suites are best summarized by comparing the CIPW normative compositions of the two parent magmas. By subtracting the CIPW normative values for the two parent magmas from each other, these differences can be expressed as:

I-type + 3A61% Or + 5A02% Ab = A-type + 2A86% An + 0A30% Hy

The major difference between the two parent magmas is the Or and Ab normative contents, which can be explained as a higher orthoclase component in the A-type parent magma, as indicated by the higher modal orthoclase in the A-type suite. Also, the lower K/Ba, Rb/Ba and higher K/Rb of the A-type parent magma, with respect to the I-type parent magma, indicates that the normative Or difference is due to modal differences in orthoclase rather than biotite. This conclusion implies that the source-rock is alkali feldspar rich, rather than biotite-rich, and is unlikely to have been a granodiorite or tonalite (Creaser et al. 1991).

5A5A4 Nature of the source rocks

The relatively flat REE patterns (La/LuN ~4 and ~8 for the I- and A-type suites, respectively), low absolute REE contents (LREE less than 100 times chondrite), and primitive initial Sr

87 86 ratios (Sr /Sr i = 0A7042-0A7044, S.E. Shaw - unpublished data) imply that relatively juvenile and unevolved source rocks are required for the production of both suites. The necessity of alkali feldspar as a residual mineral during melt generation further constrains the range of possible compositions of source rocks. A high-K, mafic to intermediate source-rock composition is required. The necessity of amphibole and biotite for fluid-absent partial melting in the case of the I-type suite also requires that the source is hydrous (ie. amphibole and/or biotite bearing). Suitable rock types are high-K amphibolites or diorites.

The occurrence of widespread mafic and intermediate volcanics associated with early Permian rifting in the southern New England Fold Belt imply underplating of the crust at this time. Although some of these volcanics are alkaline (eg. Asthana & Leitch 1985) and have 175 suitable compositions to act as source rocks, most are tholeiitic or sub-alkaline (Leitch 1988)

87 86 and isotopic studies show extreme depletion (Sr /Sr i = 0A7035, Flood et al. 1988). Hence, they are not considered to be suitable source components. Alternative source rocks include underplated arc magmas and arc detritus (a major component of the accretionary prism) associated with the Devonian to late Carboniferous arc system of the southern New England Fold Belt.

5A6 Petrogenetic Models

5A6A1 Previous Models

The comparable geochemical features of the mafic A- and I-type Chaelundi Complex granites are used to evaluate the various models postulated for the origin of A-type granites. Models involving fractional crystallization from a basaltic parent magma to produce A-type magmas (Loiselle & Wones 1979, Turner et al. 1992) require that extreme depletion of compatible elements must occur along with extreme enrichment in incompatible elements. Fractionation of a basalt to intermediate compositions would involve removal of plagioclase, producing a negative Eu anomaly and strong depletion of Sr, which has not occurred in the quartz monzonites of the Chaelundi Complex. Rather, the relatively flat REE pattern and undepleted compatible element contents suggest an unfractionated ‘parent’ magma with

~66% SiO2. Hence, the geochemical characteristics of the mafic end-member of the Chaelundi A-type suite rule out derivation by fractionation from a basaltic liquid.

The residual source model proposed by Collins et al. (1982) invokes the melting of a source rock dehydrated and geochemically depleted by the earlier generation of I-type granites. Given the similar geochemistry of the I-type and A-type parent magmas in the Chaelundi Complex, a residual source model involving remelting of a melt-depleted source appears unfeasible (q.v. Creaser et al. 1991). Also, high fluorine and chlorine activities have been postulated in the production of A-type magmas (Collins et al. 1982, Whalen et al. 1987, Eby 1990), but these abundances in the I- and A-type parental magmas of the Chaelundi Complex (Fig. 5A7) are similar. The elevated fluorine contents of in the felsic granites of this A-type suite (and most leucoadamellites of the New England Batholith) are generated solely by fractionation. Thus, elevated aHF and aHCl are probably minor factors in the production of this 176

A-type granite suite.

Partial melting of an I-type tonalite as the source to produce A-type magmas has also been proposed (e.g. Cullers et al. 1981, Creaser et al. 1991, Skjerlie & Johnston 1993), but the Chaelundi Complex A-type parent magma is more mafic than that for the I-type suite. Indeed, the A-type is more mafic than most I-types of the New England Batholith. Melting of a previous I-type granite is also untenable because the source terrane was an active accretionary prism prior to the late Carboniferous, at which stage it melted to produce S-type granites (Flood & Shaw 1977). Clearly, no I-type granites existed at any crustal level before the late Permian.

Partial melting of tonalitic I-type granites, whilst geochemically possible, is not a likely scenario in those terranes where I-types and A-types are generated at similar times and emplaced at similar structural levels: a situation which applies to the southern New England Fold Belt (see below) and the Lachlan Fold Belt A-types of eastern Australia. Any I-type granites produced by partial melting would leave the source (e.g. Clemens & Mawer 1992, Petford et al. 1993) and be emplaced at upper crustal levels, as evidenced by the presence of high-level I-type granites and volcanic equivalents in the southern New England Fold Belt and Lachlan Fold Belt. Thus, granites remaining in the source region, to then melt several million years later, can be dismissed as an A-type source. The possibility of melting older I-type granites is also eliminated by the isotopically primitive character of the Chaelundi Complex A-type suite, ruling out the remote possibility of melting an exotic crustal fragment in the Carboniferous accretionary prism of the southern New England Fold Belt. Thus, for the Chaelundi Complex example, generation of the A-types from I-types is virtually impossible, a situation which is likely in many subduction-related environments.

5A6A2 Charnockitization of the lower crust and generation of A-type magmas

An alternative model for the origin of A-type granites involves partial melting of a lower crustal source that was dehydrated, but not geochemically depleted, during the thermal event that generated the I-type magmas. In the Chaelundi Complex, where the two parental magmas have been derived from similar source rocks and through similar degrees of partial melting, the differences between each type must be explained by the different conditions 177 prevailing during partial melting. The ‘dry’ nature of A-type melts has long been recognized (Barker et al. 1975, Loiselle & Wones 1979, Collins et al. 1982) and this observation is supported by preserved orthopyroxene and late crystallizing (interstitial) biotite and hornblende in the Chaelundi A-type. Elevated contents of HFSE, LILE and LREE in A-type magmas has been attributed to low a , high a and high a by many authors (Collins H2O HCl HF et al. 1982, Whalen et al. 1987, Eby 1990), but the Chaelundi A-type parent magma does not show enrichment of F or Cl compared to its I-type counterpart. Accordingly, it is low a H2O that produces a relative increase in aHCl and aHF in A-type magmas, a feature that is magnified with fractionation (Fig. 5A7). Therefore, low water activity in the source rocks is a prime determining factor in producing the geochemical characteristics of the Chaelundi A- type suite. Low a also requires higher temperatures during partial melting (Clemens & H2O Vielzeuf 1987, Holtz & Johannes 1991), which is consistent with evidence from A-type magmas (Clemens et al. 1986).

Experimental work involving production of granitic magmas under water-saturated and undersaturated conditions (for SiO2 oversaturated source compositions) indicates that with decreasing a , melt compositions become enriched in normative orthoclase at the expense H2O of plagioclase (Conrad et al. 1988, Holtz & Johannes 1991), as the An-Ab-Or cotectic trough shifts towards the Or apex. This may also explain why the particular source rocks involved in the generation of the A-type magma did not undergo partial melting during production of the I-types, as source rocks higher in normative Or-content are less likely to undergo partial melting because they are removed from the ternary minimum or cotectic trough in the Q-An-

Ab-Or-H2O system. Collins et al. (1989) used field and geochemical evidence to indicate that partial melting of quartzo-feldspathic rocks only occurs if the Q-Ab-Or ratio closely corresponds to the ternary minimum (or eutectic). With decreasing a , the higher-K source H2O rock composition will eventually coincide with a cotectic trough closer to the Or apex, and will then undergo partial melting.

I-type granite production, and dehydration of the A-type source rocks.

The production of I-type granites by fluid-absent partial melting occurs under granulite facies conditions (e.g. Brown & Fyfe 1970, Stevens & Clemens 1993). The metaluminous I-type 178 character of the Triassic New England Batholith granites, suggests that a mantle-derived component was involved, but whether or not it was the heat source needs to be ascertained. Chappell (1978) has argued that mantle-component of these granites was produced by prior magmatic underplating, but the heat source for remelting to produce I-type granites was unspecified. The Permian tectonic history of the southern New England Fold Belt is complicated by major extensional and compressional events, and a general lack of granite production (e.g. Collins et al. 1993), so independent evidence is required to elucidate the cause of partial melting.

Rb-Sr mica ages from older (~300 Ma) S-type granites and mylonites of the southern New England Fold Belt (Landenberger et al. 1995) show the terrane was uplifted to upper crustal levels at least 10 Ma before production of the I-type granites. The terrane remained at those levels during and after voluminous granite emplacement, indicated by the preservation of extensive I-type volcanic equivalents and subvolcanic complexes (Shaw & Flood 1981). Accordingly, the lower crustal granulites produced during granite generation cannot have formed by crustal overthickening and must have undergone isobaric cooling (Ellis 1987). Given the lack of substantial early Triassic extension, the heat source for partial melting and generation of the granulitic lower crust must have been intraplating and underplating by mantle-derived basaltic magmas, probably associated with subduction, which generated the Permian and Triassic granites of the northern New England Fold Belt (Gust et al. 1993).

Heating of the lower crust from external sources is essential for the production of later A-type granites. It is the continued supply of mantle-derived heat that is capable of elevating lower crustal temperatures to above 850E-900EC (q.v. Clemens et al. 1986), once the buffering effect of fluid-absent dehydration melting reactions is removed. In this way, anhydrous granulite facies lower crustal assemblages can be melted.

Experimental evidence indicates that melting begins at grain corners and is dispersed along grain edges at low melt fractions (Watson 1982, Laporte 1994). For aqueous fluids, the dihedral (or wetting) angle, tends to be ~60E, which will generate an interconnected fluid network, capable of dramatically increasing the permeability of the crust (Watson & Brenan

1987). Thus, for only small degrees of partial melting, the resultant H2O-rich magmas (e.g. Stevens & Clemens 1993), are capable of rapid infiltration through the source region. Under 179 fluid absent conditions, melting to produce I-type magmas proceeds by the reactions:

Bt + Qtz + Pl = Opx ± Cpx + Kfs + Melt (1) (Clemens & Wall 1981) and Hbl + Qtz = Opx + Cpx + Pl + Melt (2) (Rushmer 1991) for rocks of metaluminous composition. Under these conditions, the water-rich melt phase that is produced, is capable of immediate segregation from its source, particularly if the crust is subjected to deviatoric stress (Sawyer 1994), which is the typical situation in subduction- related environments. Therefore, even with low degrees of partial melting, either fluid present or fluid absent, the hydrous granitoid magmas are capable of efficient migration, at least on a metre scale, as indicated by the existence of migmatitic terranes (e.g. Phillips 1980).

Larger-scale granitoid melts are also efficiently extracted from the lower crust. Field evidence and numerical modelling indicate that they may be rapidly extracted along narrow, self-generating, dyke networks (Clemens & Mawer 1992, Petford et al. 1993), or in shear zones (e.g. D’Lemos et al. 1992), where the surface:volume ratio of the ascending melt to its host source rock is high. Active tectonism will enhance the rate of extraction and migration, by compression of magma and generation of new, structurally concordant conduits (Collins & Sawyer 1995). Under these conditions, the source region may be drained as soon as the interconnected melt framework forms, at melt fractions of <5% (Sawyer 1994). Rapid, efficient extraction of granitic magmas at such low melt fractions, favours disequilibrium melting (Sawyer 1994), which suggests that the chemical composition of the lower crust above and within parts of the source region will remain virtually unchanged after I-type granite melt removal (Fig. 5A13). The major effect is selective loss of the water-saturated melt phase or H2O-rich fluid fraction, and retention of the insoluble CO2-rich fluid fraction (Watson & Brenan 1987), effectively dehydrating the lower crust.

The effect of dehydration in and above the source region is recrystallization to produce anhydrous, but not refractory, granulite facies mineral assemblages. If biotite was originally a oe Crust Lower

Biotite and amphibole breakdown reactions at T>850°C. Small degrees of melt production in less fertile source rocks. Evolved fluids escape and are transported away from source by I-type magmas evolved from nearby fertile a sources. Source is effectively

b b a b

Large degrees of partial melting in fertile source rocks, involving biotite and amphibole breakdown reactions at T>850°C. Extraction of I-type granites to upper crust leaves a charnockitic residue. Mantle input (basalt) Figure 5. 13. Schematic representation of partial melting processes during I-type magma production, with charnockitization of unmelted rocks facilitated by effective removal of volatiles from the lower crust by magma transport. Underplating and injection of basaltic sills into the lower crust initiates dehydration melting of fertile lower crustal rocks

(amphibole and mica breakdown) and I-type granites are formed (b). Dehydration reactions also occur in less fertile rocks and in the layer above the melt zone without significant 180 melt generation (a). Overall temperatures are buffered by dehydration and fusion reactions. Reaction zones ‘a’ and ‘b’ are expanded in detail at left. 181 the dominant mafic phase, the new assemblage now has alkali feldspar as a major component and the crust is charnockitized (reaction 1). The effect will be enhanced by the entrapment of the unreactive, CO2-rich, volatile dregs during the dehydration event (Watson & Brenan

1987, Stevens & Clemens 1993), producing the typical CO2-rich granulites of charnockitized crust (e.g. Hansen et al. 1984, Friend 1985). If the dehydration zone is hornblende-rich, the dominant feldspar produced is plagioclase (reaction 2).

Partial melting of the dehydrated source rocks.

Melting of the dehydrated and ‘charnockitized’ lower crust at elevated (>900EC) temperatures will produce A-type granites. During production of the voluminous I-type magmas from fertile source rocks by reactions 1 and 2, less fertile rocks will undergo similar reactions accompanied by loss of fluid (primarily H2O), but with little or no melt generation:

Bt + Qtz + Pl = Opx ± Cpx + Kfs + H2O ± Cl2 ± F2 (3) and

OH-Hbl + Qtz = Opx + Cpx + Pl ± F-Hbl + H2O (4)

Once the temperature buffering effect of these reactions is removed, temperatures will continue to rise to >900EC and melting of the dehydrated source will then proceed by the congruent fluid-absent melting reaction:

Qtz + Kfs + Pl ± F-Hbl = Melt (5)

Though many authors regard melting of anhydrous crustal rocks as unlikely (e.g. Clemens & Vielzeuf 1987, Whitney 1988), it is possible, provided that underplating by mantle-derived magmas continues after dehydration. This can occur in a subduction-related environment, as is suggested for the Chaelundi Complex, or in a rift-setting. Our model is similar to that of Kilpatrick & Ellis (1992) for the origin of ‘C-type’ magmas (igneous charnockites), in which a fertile, but ‘anhydrous’ granulite is melted at high temperatures. 182

A- and C-type magmas C-type magmas share many of the characteristics of A-type granites, and indeed some rock types have been assigned to both groups. For example, the pigeonite-bearing Yardea Dacite from South Australia has been classified as A-type (Creaser & White 1991) and C-type (Kilpatrick & Ellis 1992), and some Precambrian charnockitic granite suites have also been classified as A-types (e.g. Whitney 1992). However, the igneous charnockites are generally more mafic, with parental magmas containing ~62% SiO2, compared with the 66% SiO2 content of the Chaelundi A-type parent, and contain hornblende rather than biotite as the late crystallizing (interstitial) phase in fractionated members of the charnockite magma series

(Kilpatrick & Ellis 1992). At similar SiO2 contents (67%), C-type granites have higher TiO2, total FeO, MgO, CaO, P2O5 and K2O, but Al2O3 and Na2O are consistently lower (q.v. Kilpatrick & Ellis 1992, Table 3). K/Ba, Rb/Ba and Ca/Sr are also higher, and K/Rb lower (Fig. 5A10). Nonetheless, the trace-element variation patterns are remarkably similar and suggest a common origin (Fig. 5A14).

Kilpatrick & Ellis (1992) have suggested that C-type magmas are produced by melting of mafic to intermediate high-K sources in which hornblende was previously dehydrated (reaction 4), and no longer a stable phase. The low CaO contents (and high Ca/Sr ratios, Fig. 5A10a) of C-type magmas is evidence that supports residual plagioclase, but not hornblende, in the source region. In contrast, the lower Ca/Sr ratios of the Chaelundi parental A- and I- type magmas are identical, indicating that both hornblende and plagioclase were stable in the A-type source prior to partial melting. On the other hand, the lower K/Ba and higher K/Rb of the Chaelundi A-type relative to the I-type (Fig. 5A10b,c) indicate that biotite was not a residual phase during A-type magma production. Note, however, the higher K/Ba, but lower K/Rb of C-type magmas relative to the Chaelundi Complex and Lachlan Fold Belt A-type granites, suggesting that alkali feldspar is probably an important residual phase during C-type magma production. Thus, both magma types may be produced by partial melting of a dehydrated, K-rich mafic to intermediate source, with the differences in melt composition reflecting variations in the anhydrous mineral assemblages that remained in the lower crust after the previous I-type granite forming event. against (a) primordial mantle (Sun & McDonough 1989) and (b) the mafic Chaelundi I-type (CC7). Chaelundi Complex with C-type magmas (shaded) at ~67% SiO (Kilpatrick & Ellis Chaelundi 1992), Complex normalized with C-type magmas (shaded) at ~67% SiO Figure Rock / I-type Rock / Primordial Mantle 100 14. Spidergram plots comparing 5 the A-types (solid lines) and average I-type (dashed line) of the 0.2 1.0 5.0 . 10 Pb Pb bTh Rb bTh Rb aU Ba aU Ba ai Calni A-type Chaelundi Mafic magmas C-type ai Calni A-type Chaelundi Mafic magmas C-type ai Calni I-type Chaelundi Mafic La K La K Nb Nb Ce Ce Sr Sr 2 Nd Nd P P rY Zr rY Zr (a) Ti Ti (b) Na Na 183 184

5A7 Conclusions

The geochemical characteristics of the mafic end-member of the Chaelundi A-type suite, in particular the REE patterns, imply that it was an unfractionated parent magma. The A-type suite were produced under the following conditions: (1) low a and consequent higher H2O temperatures of melting produced a parent magma with higher alkalis, HFSE, and LREE contents, but lower Ca, Mg and compatible trace element contents, compared to the associated I-type; (2) the resultant hot, dry, low viscosity parent magma underwent extreme fractionation, which magnified the differences of the A-type leucoadamellite - high SiO2, Rb, Pb, F, Ga, Fe3+, and produced the characteristic flat REE pattern with a large negative Eu- anomaly.

Disequilibrium partial melting to form the slightly older I-type magmas generated localized areas of depleted residue, but a large proportion of the source region dehydrated only. Therefore, prior I-type melt production charnockitized the lower crust, producing a source rock that contained alkali feldspar as a stable mineral phase. Continued mafic under- and intra-plating of the lower crust during the terminal stages of subduction elevated geotherms towards the ‘dry’ ternary feldspar solidus, generating K-rich granitoid melts (Conrad et al. 1988). Consequently, the resulting, lower crustal derived A-type magmas have a similar geochemical character to the I-types, but the subtle differences reflect the drier, hotter conditions during melting. The mafic character of the Chaelundi Complex A-type parent magma, precludes an origin by melting of pre-existing I-type tonalitic/granodioritic granitoids (c.f. Creaser et al. 1991, Pitcher 1993).

The New England A-type granites formed in a tectonic environment no different to that of the I-types, and merely reflect dehydration (charnockitization) and subsequent melting of the lower crust. Nonetheless, this process is capable of generating petrographically and geochemically distinctive magmas, which readily fractionate to compositions unlike those of normal I-type granites. 185

CHAPTER 6. PETROGENESIS OF BASALTIC ENCLAVES IN THE A-TYPE WOODLANDS QUARTZ MONZONITE

6A1 Introduction

The term ‘enclave’ was introduced by Lacroix (1890) to describe fragments of distinctive rock enclosed in otherwise homogenous igneous rocks. Today, the term enclave is used to cover all types inclusions in igneous rocks, from millimetre (single crystal) to metre scale, and includes both cognate (endogenous inclusion or ‘autolith’) and accidental (exogenous inclusions or ‘xenoliths’) types. While this subdivision makes a primary distinction between the two major types of inclusion, several terms have been proposed for the nomenclature of the differing types of cognate inclusions (Didier & Barbarin 1991). These include mafic microgranular enclave (MME), felsic microgranular enclave (FME), schlieren, cumulate enclaves and surmicaceous enclaves.

Over the past twenty years numerous studies have been published dealing with the petrogenesis of enclaves, most involving the MME type, probably because these are by far the most abundant and enigmatic of the cognate enclaves. Several contrasting theories have been proposed for their origin, and these essentially fall into three groups. Firstly, the restite theory proposes that these inclusions, together with much of the mafic mineralogy of granites, represent refractory residuum (restite) remaining from the partial melting event which gave rise to the granitoid magma (e.g. Chappell et al. 1987). In contrast, the magma mingling theory, favoured by many petrologists, proposes that these inclusions represent globules of immiscible mafic magma which has mingled, but not mixed (homogenized), with their granitoid hosts (e.g. Barbarin & Didier 1991, Didier 1987, Vernon 1983, 1991). In addition, several other theories have been proposed, mainly suggesting a cumulate origin for MME, either as coarse-grained cumulates (e.g. Dodge & Kistler 1990), or as disrupted fine- grained mafic quench margins (e.g. the pressure-quench model, Flood & Shaw 1991).

The term microgranitoid enclave (ME) was introduced by Vernon (1983), and will be used herein in preference to mafic microgranular enclave (MME), as it emphasizes the significance of the igneous texture which is displayed by these enclaves, and avoids the possible inference 186 that they have a metamorphic ‘granular’ texture, and hence a metamorphic origin.

Although microgranitoid enclaves are common in ‘calc-alkaline’ granitoids (Barbarin & Didier 1991), particularly those of I-type affinity (Chappell & White 1991), they are generally rare in leucogranites (Vernon 1983), particularly those of A-type affinity (Bonin 1991). This study deals with an unusual occurrence of enclaves in an A-type granites of the New England Batholith of northeastern New South Wales. These enclaves also have particular significance, since they are unusually mafic (basaltic) and contain intact vestiges of their primary basaltic mineralogy.

6A2 Geological Setting

During the latest Permian & Triassic, the compressional events of the late Permian (see Chapter 3) were followed by rifting and intrusion of the voluminous I- and A-type granites of the New England Batholith (Shaw & Flood 1981, Collins et al. 1993) and extrusion of their volcanic equivalents. The intrusives forming this belt have been subdivided on the basis of textural and geochemical differences into three major suites (the Moonbi, Uralla and Clarence River Plutonic Suites) and an group of leucoadamellites which have not been assigned to suites. Together, these four groups of granitoids comprise ~60% of the New England Batholith (Shaw & Flood 1981).

The leucoadamellites of the New England Batholith mostly occur as large plutons in the northern and eastern portions of the Batholith (figure 6A1). Most are coarse-grained silicic rocks containing biotite, sometimes accompanied by minor Fe-rich hornblende, or alternatively muscovite (Shaw & Flood 1981), together with the accessory minerals zircon, allanite and apatite. Most are highly silicic (>74% SiO2, Shaw & Flood 1981) and are mildly corundum normative. The petrography and geochemistry of this group of granites is broadly similar to the A-type granites of the Lachlan Fold Belt (Whalen et al. 1987), and as such, are often referred to as A-type (Landenberger & Collins 1996). Although the leucoadamellites have a similar age span to the accompanying I-type granites, field relationships demonstrate that they always post-date the enclosing I-type plutons, despite isotopic ages for spatially associated I- and A-type plutons being indistinguishable (Landenberger & Collins 1996, Shaw & Flood 1993). 187

Cover rocks

Map area Leucoadamellite and related rocks Other granites and volcanics of the New England Batholith New England Fold Belt 0 50 km N

Clarence- Great Moreton Australian Basin Basin

30° 00' S

Peel

- Figure 6. 2

Southern New England Fold Belt Tablelands Complex

Manning

Tamworth Pacific Belt Sydney Fault 152° 00' E Ocean Basin System

Figure 6. 1 Geological map showing the distribution of leucoadamellites and related rocks in the southern New England Fold Belt. 188

Despite the leucocratic nature of most plutons in the group, more mafic rocks do occur. These are dominantly adamellites and quartz monzonites which represent parental melts which have fractionated to leucoadamellite compositions (Landenberger & Collins 1996, chapter 5). The quartz monzonites form small discrete intrusive bodies (such as at Woodlands), or occur as small mafic margins of the larger leucoadamellite plutons (e.g. in the Chaelundi Complex, Landenberger & Collins 1996). Although these quartz monzonites are more mafic (varying down to 65% SiO2), they share the felspar-rich nature of their fractionated derivatives. These A-type quartz monzonites are distinguished from other monzonitic rocks of the batholith (i.e. monzonites associated with the I-type Moonbi Plutonic Suite) by lower MgO, CaO, Sr and higher HFSE and LREE contents. The most distinguishing feature of these two groups of monzonites is Mg# (100Mg/(Mg+3Fe)), with Mg# for the A-type quartz monzonites typically very low (Mg#.20 for Woodlands Quartz Monzonite) while the Moonbi Suite quartz monzonites have Mg# >60 (Chappell 1978).

6A3 Microgranitoid enclaves in A-type granites of the New England Batholith

Microgranitoid enclaves are exceedingly rare in the coarse grained leucogranites of the New England Batholith. These leucogranites are generally free of inclusions on all scales, although the more mafic members of this group (quartz monzonites) often contain single corroded crystals of calcic plagioclase rimmed by sodic plagioclase, and magnesian augite and hypersthene, which have reaction coronas of actinolite and cummingtonite respectively (e.g. Landenberger & Collins 1996). Where these more mafic A-types form smaller intrusions and are consequently finer grained (e.g. the quartz monzonite at Woodlands) or form the finer grained margins of larger plutons (e.g. Smokey Cape Adamellite), microgranitoid enclaves up to 1 metre in diameter are not uncommon, and are particularly abundant in the Woodlands Quartz Monzonite.

6A4 Field relationships of the Woodlands Quartz Monzonite and its enclaves

6A4A1 Age and intrusive relationships

The Woodlands Quartz Monzonite (informal name) is a relatively small intrusion, ~7 km long and varying in width from 200 m to 1 km (figure 6A2), which was emplaced at shallow 189

Glen Innes

30°00' S River Sara

Glen Innes

River

Guyra Fault

Gle Oban 152°00' E n N

Bluff F

ibinda

lt Guyra Won 0 10

Woodlands Quartz Monzonite km Triassic Oban River Leucoadamellite Roads Wards Mistake Adamellite » Major Mornington Diorite drainage Hillgrove Suite granitoids Late Carboniferous Carboniferous & Permian metasediments

Figure 6. 2. Geological map of the Kookabookra - Ward's Mistake area showing the location of the Woodlands Quartz Monzonite. 190 crustal depths in the early Triassic. Paired biotite - whole-rock Rb-Sr analyses give an emplacement age of 249±2 Ma and an initial 87Sr/86Sr of 0A7061±0A0001 (see appendix F). The pluton predominantly intrudes the deformed Kookabookra Adamellite of the late Carboniferous Hillgrove Plutonic Suite, and appears to coincide with the northerly extension of the Glen Bluff Fault (a mylonite zone in this deformed granite).

Intrusive relationships at the northern extremity of the Woodlands pluton are unclear, as it has either intruded, or is intruded by the Oban River Leucoadamellite (see figure 6A2). The Oban River pluton has biotite Rb-Sr ages of 248 Ma (Shaw & Flood 1993) and 246 Ma (Hensel 1982), which are indistinguishable from that of the Woodlands intrusion. The I-type Wards Mistake Adamellite also has biotite Rb-Sr ages indistinguishable from the Oban River and Woodlands plutons (246 Ma, Shaw & Flood 1993; 249 Ma, Hensel 1982), but predates the Oban River pluton on the basis of field relationships (see figure 6A2).

6A4A2 Outcrop detail

The most accessible outcrop of the Woodlands Quartz Monzonite (WQM) occurs near its northern end, along a 200m stretch of the Sara River (Fig. 6A2). At this locality, sharp, near- vertical contacts with the enclosing Kookabookra Adamellite occur, and two phases of intrusion are evident. The first phase is a more mafic quartz-poor monzonite which is only preserved as a narrow (~3 metres) margin at the western extremity of the intrusion. The second phase is a quartz monzonite which dominates the mass, and forms sharp vertical contacts with the first phase. This phase gradually becomes more felsic, grading into adamellite in the centre of the body. The internal contact between the two phases parallels the primary margin of the intrusion (Plate 6A1a), and shows that more felsic phase is younger. The first phase of intrusion is relatively enclave-free, whereas the second contains abundant enclaves varying from ~5 cm up to ~1 m in diameter.

Three main types of enclave occur in the quartz monzonite. The first type are highly silicic, well-rounded inclusions similar to the felsic microgranular enclaves (FME) considered to represent disrupted early fine-grained margins of granitic hosts (Didier & Barbarin 1991). These felsic enclaves (plate 6A1b) have similar meso- and microscopic textures to quenched, 191

Plate 6A1. Outcrop features of enclaves in the Woodlands Quartz Monzonite.

(a) Intrusive contact between the early mafic monzonite phase and the (later) more felsic quartz monzonite, showing a mafic microgranitoid enclave in the early phase with a preserved rind of its host, protruding into the latter phase. Hammer handle is 3 cm thick.

(b) Fine-grained, felsic microgranular enclave, showing a schlieren tail (bottom right of enclave). Hammer handle is 3 cm thick.

(d) Contact between a fine-grained phenocryst-free enclave, and a coarser-grained phenocrystic enclave. Note that the fine-grained enclave is convex into the phenocrystic enclave. Lens cap is 50 mm in diameter. 03

(a) (b) 14

13 192 highly felsic carapace rocks which outcrop to the north of the Chaelundi Complex, and form part of this A-type suite.

The two remaining types of enclave dominate the population and are the microgranitoid enclaves of Vernon (1983). The two types of mafic microgranitoid enclaves are end- members of an essentially continuous textural (and compositional) spectrum. At one end of this spectrum are the ‘typical’ fine-grained microgranitoid enclaves which are very similar to enclaves occurring in New England Batholith I-type suites. They have a uniform fine- grained texture apart from occasional inclusions of orthoclase (plate 6A2a) and quartz ‘ocelli’, which are common in enclaves and are often inferred to be xenocrystic - i.e. they are derived from the host magma (e.g. Vernon 1991). Orthoclase inclusions always have plagioclase overgrowths which can be identified in hand specimen, a feature also observed in enclaves from the Bundarra Suite (S-type) and the Chaelundi and Moonbi I-type suites. These fine- grained enclaves form approximately half of the population of mafic enclaves present and are referred to herein as phenocryst-poor enclaves or PPE.

The second type of mafic enclave occurring in the Woodlands pluton is a more unusual type. These enclaves are strongly porphyritic with pyroxene and plagioclase phenocrysts averaging ~4 mm in size, which are set in a fine grained matrix (plate 6A1d). This enclave type seldom contains the xenocrysts common in the PPE variety, and occasionally show quenched margins (plate 6A2b), and will be referred to herein as phenocryst-rich enclaves or PRE.

The shape of microgranitoid enclaves in the WQM is variable, some are well rounded (plate 6A1a), while others are lobate, showing both convex and concave margins (plate 6.2a) with the host (a feature common in enclaves of the Smokey Cape Adamellite). Enclaves commonly appear to be at various stages of disaggregation and invasion by the host quartz monzonite (plates 6A2c,d). Where the finer-grained phenocryst-poor enclaves are in contact with the phenocryst-rich type, the former are always convex into the latter (plate 6A1d). 193

Plate 6A2. Outcrop detail of mafic microgranitoid enclaves.

(a) Fine-grained mafic enclave with orthoclase xenocrysts and crenulate margins showing ‘embayment’ by host granitoid. Lens cap is 50 mm in diameter.

(b) Mafic microgranitoid enclave with a very fine-grained quench margin and phenocryst-rich core. Lens cap is 50 mm in diameter.

(c) Enclave showing partial disaggregation along the rim, with invasion of the host granitoid. Lens cap is 50 mm in diameter.

(d) Phenocryst-rich mafic enclave showing severe disruption and invasion by the host quartz monzonite (see arrow). Hammer handle is 3 cm thick. 03

(a) (b) 14

13 194

6@5 Petrography and Mineral Chemistry

6A5A1 The host quartz monzonite

The Woodlands Quartz Monzonite varies from a quartz-poor monzonite phase to mafic adamellite. Mineralogy is dominated by normally zoned plagioclase (An20-30) and perthitic orthoclase (Or50-75), which form phenocrysts up to 7 mm in length. Interstitial to these are hornblende (α=straw yellow, β=brownish green, γ=bluish green, Mg#~22), biotite (α=straw yellow, β=γ=chocolate brown, Mg#~22) and quartz which are generally restricted to less than 1 mm (for compositions see appendix A4). Accessories include zircon, magnetite and allanite. This feldspar-dominated assemblage is similar to the mafic members of the Chaelundi A-type suite, excepting that it bears magnetite rather than ilmenite as the sole oxide phase. Another similarity to the Chaelundi A-type suite is the occasional isolated crystals or small mafic clots (xenocrysts) of magnesian clinopyroxene, orthopyroxene and calcic plagioclase, which bear the same composition as the equivalent phenocryst phases in the enclaves.

6A5A2 Enclave matrix

The fine grained mafic enclaves described above (PPE type), have microstructures and mineralogy which are common to microgranitoid enclaves of I-type granites (e.g. Vernon 1990). That is, they bear essentially the same mineralogy as their host, but are always more mafic and finer grained. The dominant groundmass phases are hornblende (α=straw yellow, β=brownish green, γ=olive green), and plagioclase (see modes appendix E) with minor biotite (α=straw yellow, β=γ=chocolate brown), acicular apatite and magnetite (plate 6A3e). However, unlike the groundmass of enclaves from other New England Batholith suites, and microgranitoid enclaves generally (Vernon 1983), the groundmass of the Woodlands enclaves contains no quartz or orthoclase, with these phases only occurring as xenocrysts. There is also a general trend of decreasing matrix grainsize from phenocryst-rich enclaves (0@25-0@5 mm) to phenocryst-poor enclaves (0@125-0@25 mm).

Although hornblende, biotite and plagioclase are common phases to the enclave matrix and host, the mineral chemistry of these phases in the enclaves is consistently different to their 195

Plate 6A3(a-d). Woodlands Quartz Monzonite - Enclave microstructure.

(a) Orthopyroxene phenocryst (phenocryst-rich enclave), showing replacement by fibrous cummingtonite and biotite (extreme margin). Oxide inclusions are titanomagnetite showing replacement by biotite. Base of plate = 3A5 mm.

(b) Cluster of clinopyroxene and plagioclase phenocrysts (phenocryst-rich enclave) - the large CPX phenocryst displays the typical limited replacement by fibrous actinolite and a magmatic overgrowth of hornblende. Base of plate = 7 mm.

(c) Large sieve-textured plagioclase phenocryst with a uniform An90 core and overgrowth and infilling by oligoclase. Base of plate = 7 mm.

(d) Partly rounded plagioclase phenocryst showing a relatively uniform calcic core with

minor oscillatory zoning (An90-70) and overgrowth of oligoclase. Base of plate = 3A5 mm. (a) (b)

Plate 6A3(e-h). Woodlands Quartz Monzonite - Enclave microstructure.

(e) Fine-grained matrix typical of all mafic enclaves, showing intergrowth of plagioclase and hornblende along with minor magnetite and acicular apatite. Base of plate = 1 mm.

(f) Orthoclase xenocryst in a phenocryst-poor enclave showing typical overgrowth of oligoclase. Base of plate = 7 mm.

(g) Xenolith of host granite (oligoclase + orthoclase + quartz) in a phenocryst-poor enclave. Base of plate = 7 mm.

(h) Rim of a phenocryst-poor enclave showing minor calcic plagioclase phenocrysts. The large alkali-feldspar xenocryst transgressing the enclave margin shows an overgrowth of oligoclase on that part contacting the enclave matrix. Base of plate = 7 mm. (e) (f)

(g) (h) 197 equivalents in the host granitoid. Both biotite and hornblende in the enclave matrix are more magnesian than that of the host rock (figure 6A3a and 6A4). Hornblende in the quartz monzonite has a consistent Mg#~22, while the hornblende in enclave matrix varies from Mg# 26 - 52, with the fine-grained matrix of the phenocryst-free enclaves forming the more magnesian extreme of this range (figure 6A3a). A similar trend is evident for biotite, with that in the host restricted to Mg#~19, while the enclave biotites vary from Mg#~28 for PRE to Mg#~38 for PPE (figure 6A3a). Plagioclase compositions overlap for host and enclave matrix

(figures 6A3b, 6A6) and range from An20 in the phenocryst-bearing enclaves which coincides with a restricted range in the host, to An52 in phenocryst-free enclaves. Ilmenite is a rare groundmass phase in the Woodlands enclaves, with magnetite usually the sole oxide phase, although it also occurs in trace quantities.

6A5A3 Enclave phenocrysts

The phenocryst-rich enclaves of the Woodlands Quartz Monzonite have a primary igneous (phenocryst) mineral assemblage similar to that of mildly evolved island-arc (‘calc-alkaline’) basalts (Ewart 1982). The phases commonly present are calcic plagioclase, augite, hypersthene and titanomagnetite (in decreasing order of abundance). All of these primary phases show evidence of magmatic corrosion (rounding and embayment), display gross disequilibrium textures with the matrix and are at various stages of alteration. Some enclaves are extremely phenocryst-rich and bear textures similar to cumulates, however, this feature is somewhat disguised by the embayment and alteration present.

Plagioclase phenocrysts occur in two main forms. The first type (type 1 plagioclase) are similar to the ‘spongy cellular plagioclase’ (plate 6A3c) described by Hibbard (1991). They are essentially skeletal crystals formed by high growth rates, which have undergone subsequent partial dissolution, followed by infilling by more sodic plagioclase and other phases which occupy the enclave groundmass. These phenocrysts are relatively uniform in composition, and are highly calcic, varying from An80 to An95. They are commonly corroded (well-rounded) and are overgrown (both on the rims and interstices) by sodic plagioclase of the same composition as that in the enclave matrix (~An20-30). Crystals showing the same texture as these phenocrysts are occasionally observed in the host granitoid. The second type 198 (a) augite Mg# hypersthene 60 hornblende

biotite 40

enclave groundmass 20 host phases

40 50 60 70

SiO2 (b) An# 80 plagioclase phenocrysts 60 enclave groundmass 40

groundmass 20 plagioclase host plagioclase (& xenocrysts in enclaves) 2 50 60 70 SiO 2 (c) K/(K+Na) Orthoclase

0.6 enclave groundmass

0.4

0.2 . 6 3phenocryst for granitoid. Field host and groundmass phenocrysts, enclave in (100*Mg/(Mg+Fe)) Mg# phase ferromagnesian of Comparison (a) Figure phases (CPX & OPX) - red & cyan fill respectively, field for matrix phases (biotite & hornblende) - brown & green fill respectively. The grey field covers bulk covers field grey The respectively. fill green & brown - hornblende) & (biotite phases matrix for field respectively, fill cyan OPX) & & red (CPX - phases enclave groundmass compostions. Diamond symbols are the host quartz monzonite (purple) and bulk enclaves (black), inverted triangle is the Oban River Leucoadamellite. The yellow field covers A-type graniotids of the SNEFB. the of graniotids A-type covers field yellow The Leucoadamellite. Comparison of plagioclase compositions (An# vs SiO ); plagioclase phenocrysts and groundmass plagioclase are dark blue and light blue fields respectively. respectively. fields blue SiOlight vs and (An# blue compositions dark plagioclase are of plagioclase Comparison groundmass (b) and phenocrysts plagioclase ); Filled areas as for (a) - average host plagioclase composition is also plotted for comparison. for plotted also is composition plagioclase host average - (a) for as areas Filled Alkali feldspar compositions of both host granitoid and enclave xenocrysts plotted as molecular 100*K/(K+Na). Filled areas as for (a) with the addition of the of addition the with (a) for as areas Filled 100*K/(K+Na). molecular as plotted xenocrysts enclave and granitoid host both of compositions feldspar Alkali (c) 50 60 70 (pink). field feldspar alkali SiO2 199 Di Hd CPX phenocrysts

matrix hornblende Actinolite after CPX host hornblende

Cummingtonite after OPX OPX host phenocrysts matrix biotite biotite

En Fs

Figure 6. 4 Comparative compositions of ferromagnesian phases in enclave phenocrysts, matrix and host granitoid plotted in the pyroxene quadrilateral.

1 Tremolite Tr Hb Actinolite Actinolite secondary Tsch after augite Act Magnesio-Hbl Tschermakite Hbl Hbl

2+ Primary hornblende

Mg/(Mg+Fe ) Fe- Fe- Ferro- Ferro- Act Ferro-Hbl Tsch Actinolite Tschermakite Hbl Hbl

0 8.0 7.5 7.0 6.5 6.0 5.5 TSi

Figure 6. 5 Classification of primary matrix hornblende and secondary actinolite after clinopyroxene phenocrysts - Woodlands enclaves (Hawthorne 1981). 200 An

Enclave phenocrysts orthite an

ite

bytown

labradorite

Enclave matrix, xenocrysts and host phenocrysts andesine

goclas oli

te orthoclase

albi Enclave xenocrysts and Ab host phenocrysts Or Figure 6. 6. Feldspar compositions from the Woodlands Quartz Monzonite and enclaves plotted in the An-Ab-Or ternary system.

Wo Wo Di Hb 500° 600° 700° 800° 900° 1000° 1100°

1200°

1200° 1100° 1000° 900° 800° 700° 600° 500° En Fs

Figure 6. 7. Pyroxene thermometry of enclave phenocrysts from the Woodlands Quartz Monzonite. Analyses are plotted on the 1 kb graphical thermometer of Lindsley (1983). 201 of plagioclase phenocryst (type 2 plagioclase) are not ‘cellular’, are more euhedral (due to limited magmatic corrosion), and display weak oscillatory zoning (plate 6A3d). These plagioclases are somewhat less calcic, ranging from An85-70. Like the cellular type, they are also overgrown by sodic plagioclase (~An20-30). Both types of plagioclase phenocrysts vary in size up to 7 mm.

Augite is the second most common phenocryst phase, and occurs in all phenocryst-bearing enclaves and attains crystal sizes up to 7 mm (plate 6A3b). The augite present is similar to that found in island arc basalts (Mg#~70, Ewart 1982). These augite phenocrysts are commonly overgrown by hornblende of the same composition to that occurring in the enclave matrix (Mg# 50-30), and occasionally has a skeletal appearance with inclusions of hornblende + plagioclase, similar to those occurring in the plagioclase phenocrysts. Rounding by magmatic corrosion is less evident than for plagioclase, but alteration (uralitization), commonly to fibrous aggregates of actinolite (Mg# 70-40), is common (figure 6A3a, 6A4). This alteration varies in degree from partial, where it is concentrated around the rims of the phenocrysts (immediately beneath overgrowths of magmatic hornblende), to complete pseudomorphous replacement. The texture formed by complete replacement results in pseudomorphous decussate aggregate of fibrous actinolite rimmed by magmatic hornblende, a feature which is common in many enclaves from I-type granites, and within the host granitoid at Woodlands. Examples include enclaves in the Chaelundi I-type suite (Landenberger 1988), the Moruya Suite I-type suite of the Lachlan Fold Belt (Keay 1992), and in the Moonbi Suite of the New England Batholith (Collins & Landenberger, unpublished data).

Unaltered hypersthene is somewhat more restricted than augite, occurring only in the most mafic enclaves, where phenocrysts vary in length up to 3 mm in length (plate 6A3a). Where inclusion relationships can be found, augite appears as overgrowths on hypersthene, and therefore the latter has probably begun crystallization before augite. The phenocrysts are commonly mildly rounded and show magmatic overgrowths of biotite which is of the same composition as that occurring in the enclave matrix (Mg# 25-35). Like the augite, hypersthene also shows partial to complete alteration (uralitization) which pervades from the rim of the phenocrysts inwards. However, in this case, cummingtonite (Mg# 58-75) is the replacement mineral, and occurs as fibrous aggregates, pseudomorphing hypersthene (plate 202

6A3a), forming a similar texture to the actinolite which replaces augite. Hypersthene generally shows somewhat more extensive alteration than does augite, and while many enclaves have no unaltered phenocrysts, they commonly have decussate aggregates of fibrous cummingtonite. The outer margins of these aggregates of cummingtonite are commonly composed of biotite (plate 6A3a). Whether this biotite is a primary magmatic overgrowth (as in the case of less altered phenocrysts) or whether it is of secondary origin, is uncertain. The degree of K-saturation in the matrix ‘magma’ early in the alteration history of these phenocrysts may have favoured the formation of biotite rather than cummingtonite as the replacement mineral.

Titanomagnetite is a common accessory phase in most enclaves reaching a maximum size of ~1 mm. Like other phenocrysts, it commonly shows partial or complete replacement. Most phenocrysts show exsolution into paired lamellae (<500 μm in width) of magnetite and ilmenite, with the ilmenite portion commonly replaced by biotite leaving a ‘skeleton’ of magnetite.

Accessory phenocrysts include magnesian hastingsitic hornblende (α=pale yellow, β=greenish brown, γ=reddish brown, classification of Hawthorn 1981) which has been identified in only one enclave (WQMX10).

6A5A4 Enclave xenocrysts

In addition to the basaltic phenocrysts described above, crystals of granitic origin occur in the more felsic, phenocryst-poor, finer grained enclaves (PPE). These crystals vary in size up to 10 mm and include orthoclase, quartz and sodic plagioclase in decreasing order of abundance. Orthoclase crystals occurring in enclaves are the same composition as those in the host granite (Or55-75), but are distinguished by displaying partial resorption (they are strongly rounded) and have overgrowths of sodic plagioclase (rapakivi texture - plate 6A3f). Where these crystals occasionally transgress enclave margins, this rapakivi overgrowth is restricted to that portion of the crystal which is enclosed in the enclave, while the portion in contact with the granitoid host is rimmed by a further growth of orthoclase (which is usually more potassic - plate 6A3h). This is a feature commonly observed in other granitoid suites of the NEB in which have a mineralogy containing orthoclase phenocrysts, e.g. the Bundarra 203 and Moonbi Plutonic Suites (see chapter 4). The crystals are interpreted as originating from the host granitoid (i.e. they are xenocrysts) due to their similarity in composition to orthoclase in the host, the absence of orthoclase in the matrix of enclaves (including a complete absence as either crystals or matrix in PRE), and the disequilibrium textures they exhibit.

Quartz occurs as rarer crystals (up to 10 mm) in the most felsic enclaves. Like the orthoclase they display disequilibrium textures, including strong embayment and reaction rims of hornblende. These are the quartz ocelli described by Vernon (1983) and are considered to represent xenocrystic quartz derived from the host granitoid, which has undergone reaction with a more mafic magma to produce a rim of small hornblende crystals generated by localized undercooling (Hibbard 1991).

Large, sodic plagioclase crystals (An20-30) are also the same composition as plagioclase in the host granitoid, but are extremely rare. They exhibit no distinct disequilibrium textures as their composition is also similar to plagioclase in the enclave matrix. However, their composition is obviously at odds with the abundant calcic plagioclase in phenocryst-rich enclaves, and therefore a different explanation is required for their origin. Like the associated orthoclase and quartz xenocrysts, it is likely that these crystals have originated from the host granitoid.

Additional support for the interpretation that these crystals of orthoclase, quartz and sodic plagioclase are xenocrysts, comes from the observation that these three phases also occur as composite microxenoliths within enclaves, forming granitic aggregates of up to ten grains (plate 6A3g) which are frequently accompanied by minor hornblende. These microxenoliths have margins which display the same disequilibrium textures as the isolated xenocrysts, but preserve an internal texture which is identical to that of the host granitoid.

6A5A5 Pyroxene thermometry

The coexistence of pyroxene pairs in some of the phenocryst-rich enclaves lends itself to the estimation of the temperature of phenocryst crystallization. The graphical two - pyroxene thermometer of Lindsley (1983) was applied to mineral probe data for this purpose. Due to 204 the shallow nature of Triassic intrusives of the New England Batholith (Shaw & Flood 1981) the graphical thermometer for 1 kb was applied. Data for enclave pyroxene compositions are presented in figure 6@7. Only two enclaves contain coexisting pyroxene pairs and these are linked by tielines on the plot. The plot demonstrates that the pyroxene pairs are not at mutual equilibrium, with hypersthene recording temperatures of 950-1000EC, while augite records lower temperatures of ~800EC. Whether this discrepancy reflects a partial equilibration of augite to lower temperatures while hypersthene reacted more sluggishly, or whether it records the true temperature of crystallization is unclear. A lower crystallization temperature for augite is consistent with the observation that hypersthene generally appears to have begun crystallization prior to augite (as discussed earlier). Regardless of this discrepancy the data show (for single pyroxene thermometry) that these phenocrysts have crystallized at high temperatures, which is consistent with crystallization from basaltic magma.

6A6 Geochemistry

6A6A1 The host quartz monzonite

The Woodlands Quartz Monzonite has a bulk chemistry very similar to the Chaelundi A-type suite (Landenberger & Collins, see chapter 5). One distinguishing factor of the samples collected is a more restricted and lower silica range (Woodlands 65669% SiO2, Chaelundi

A-type 66676% SiO2). At similar silica contents (~69% SiO2) TiO2, MgO, Sr are lower, and Ba, Zr, Fe3+/ΣFe (Fig. 6A8 & 6A9) and LREE (Fig. 6A11) are higher in the Woodlands pluton. In essence, the suite has more pronounced A-type characteristics than the Chaelundi Complex A-type suite. This A-type chemistry is supported by the typical A-type petrographic characteristics of early crystallizing feldspars and late interstitial biotite and hornblende.

Although no sampled members of the suite are higher than 70% SiO2, the presence of high silica felsic microgranular enclaves (FME - Didier & Barbarin 1991) suggests that more felsic phases of the suite do exist. In addition, the nearby Oban River Leucoadamellite (see figure 6A2), which has similar petrographic and geochemical characteristics to other felsic A-types of the New England Batholith, may be related by a fractionation series to the more mafic Woodlands Quartz Monzonite (a single analysis of this pluton is represented by an inverted 205

TiO 8 1.2 2 CaO

6 0.8 4

0.4 2

7.0

18 Al O 2 3 4.5 17

16 4.0 15

14 3.5 Na2 O 13 3.0

10 5 EFeO 8 4 6 3 4 KO2 2 2

3 MgO 0.3 Fe3+ /EFe

2 0.2

1 0.1

55 60 65 70 75 55 60 65 70 75 SiO SiO2 2 Figure 6. 8 Major element plots for Woodlands Quartz Monzonite and enclaves. Key: Purple diamonds = Woodlands Quartz Monzonite; black diamonds = enclaves; half-filled diamonds = enclave groundmass (EDS analysis); inverted yellow triangle = Oban River Leucoadamellite. The yellow shaded area is that for A-type suites of the southern New England Fold Belt (Chaelundi, Woodlands, Smokey Cape), the green shaded field is for enclaves from the Smokey Cape Adamellite. 206

Ba 600

1000 Sr 400

500 200

Zr 200 Rb 400

100 200

2000 30 Pb F

1500 20 1000

10 500

80 Y 200 Zn 60

40 100 20

55 60 65 70 75 55 60 65 70 75 SiO 2 SiO2 Figure 6. 9 Trace element plots for Woodlands Quartz Monzonite and enclaves. Key as for figure 6. 8. 207

1000

500 Ca/Sr K/Ba

200 100

100

400 Ga*10000/Al K/Rb 3 300

200 2 100

100 1000 Mg/Cr 800 10 Rb/Sr 600

400 1 200

0.08 Mn/Fe 100 0.06

0.04 50 Mg/V 0.02

55 60 65 70 75 55 60 65 70 75 SiO 2 SiO2

6. 10 Inter-element ratio plots for Woodlands Quartz Monzonite and enclaves. Key as for figure 6. 8. 208 Yb enclaves phenocryst - rich } Ho Lu WQM - host quartz monzonite WQMX2 WQMX10 WQMX8 - phenocryst-poor enclave Chaelundi A-type quartz monzonite Tb Eu Sm Nd Ce La . 6 11 REE chondrite normalized (Nakamura 1977) plot for Woodlands Quartz 1 10 Figure Monzonite and enclaves.

100

1000 Rock/Chondrite 209 triangle in figures 6A8 to 6A10). Although the field relationship between these two plutons has not been established, their indistinguishable Rb-Sr biotite ages (see section 6A4A1) also supports a possible genetic link.

6A6A2 Enclaves

The most notable feature of enclaves from the Woodlands Quartz Monzonite is their highly basaltic character. Enclaves from other granitoid suites of the New England Batholith, such as the Hillgrove and Bundarra suites (see Chapter 4), Moonbi Suite (Landenberger & Collins, unpublished data) and the I-type suite of the Chaelundi Complex (Landenberger 1988), all have compositional ranges which overlap with that of the host granitoid, and generally do not extend below ~55% SiO2. In contrast, enclaves from the WQM are restricted to the range

~52659% SiO2, with a large compositional gap occurring between enclaves and host (figures 6A8-6A10, appendix E).

Trends exhibited by WQM enclaves on Harker plots (Figs. 6A8, 6A9) fall into three groups. For a limited number of elements, enclaves display linear trends which form extensions of the differentiation trend exhibited by the host granitoid, either decreasing in abundance with increasing silica (CaO, 3FeO), or increasing with increasing silica (K2O). However, for most elements, linear trends for the enclaves are different to the trends for host granitoid, in many cases exhibiting increasing abundance with increasing silica as abundances in the host decrease (TiO2, Na2O, Ba, Zr, Zn). The third case is one where trends exhibited by the enclaves are similar to the host granitoid, but do not form a linear extension of the host differentiation trend, and exhibit either more steeply negatively sloping trends (Al2O3, MgO, Sr), or positively sloping trends which are offset with that of the host (Rb, Pb, Y, F). This pattern is in contrast to the geochemical behaviour for many enclave suites, in which enclave compositions plot on or near extrapolation of the differentiation trends of the host granitoids (Vernon 1983). Although enclaves from some granitoid suites of the New England Batholith partly conform to the chemical behaviour described by Vernon (1983), (e.g. the Chaelundi I-type suite, Landenberger 1988) this pattern is rarely observed for all elements, and such chemical behaviour for trace elements may well be process-specific, as described by Tindle (1991). These three separate styles of trends exhibited by the Woodlands enclaves are similar to those described in the summary of enclave geochemical behaviour discussed by Poli & 210

Tommasini (1991).

The REE contents of the Woodlands enclaves also contrasts strongly with the behaviour of other enclave suites. REE contents in other New England granitoid suites (e.g. the Chaelundi I-type suite, Landenberger 1988) commonly exceed contents in the host granitoid, a feature which is also true for many incompatible trace elements (e.g. Rb). This ‘partitioning’ of incompatible elements into enclaves relative to host is a commonly observed feature of microgranitoid enclaves (e.g. Eberz & Nicholls 1990). However, REE contents do not exceed contents in the host quartz monzonite except for Eu in one sample (Fig 6A11). REE contents for the Woodlands enclaves do not exceed 100 times chondrite except for the more felsic phenocryst-free enclaves (WQMX8), there is a general trend to higher contents with increasing silica. The host granitoid exhibits a REE trend consistently steeper (La/LuN =

11A83) than do any of the enclaves (La/LuN = 7A0067A58). Both phenocryst-bearing enclaves also exhibit slight positive Eu anomalies (Eu/Eu* = 1A0461A10), while the phenocryst-free enclave analysed has a moderate negative Eu anomaly (Eu/Eu* = 0A67) similar to that in the host granitoid (Eu/Eu* = 0A69).

6@7 Enclave Petrogenesis

6A7A1 Restite, Cumulate or Mingled Magma?

As discussed earlier, models for the origin of microgranitoid enclaves may be grouped under three broad categories. Firstly, the restite theory (Chappell et al. 1987) proposes that enclaves represent refractory source rock entrained in the granitoid magma after partial melting. Secondly, theories of a cumulate origin propose that enclaves originate either as coarse grained cumulates (Dodge & Kistler 1990), or as disrupted fine-grained mafic quench margins (the pressure-quench model, Flood & Shaw 1991). In contrast to the first two groups, the magma mingling theory proposes that these inclusions represent globules of immiscible mafic magma which have mingled, but not mixed (homogenized), with their granitoid host (e.g. Vernon 1991). Petrographic, geochemical and isotopic evidence from the Woodlands example, will be used to place tight constraints on which model may be applied to this particular example. 211

Petrographic evidence The coarse-grained basaltic phenocrysts within the Woodlands enclaves, which display features such as igneous zoning, are unequivocally magmatic. This observation alone, renders an origin as restite untenable for the Woodlands enclaves. Even if the implied source for the host granitoid were meta-igneous, such a primary igneous petrographic features would be unlikely to survive high grade metamorphism and partial melting. Any refractory meta- igneous material entrained in the magma would more likely resemble high grade metamorphic rocks (metabasites), and features such as primary igneous zoning would not be preserved. The common occurrence of igneous microstructures in microgranitoid enclaves is one of the crucial arguments against a restite origin (Vernon 1983, 1991).

The preservation of these basaltic phenocrysts in the Woodlands enclaves also renders a cumulate origin unlikely. The pressure-quench cumulate model (Flood & Shaw 1991) was proposed to explain the fine grained nature of microgranitoid enclaves and their mineralogical similarity to host rocks. Both the size of phenocrysts in the Woodlands enclaves, and their strongly contrasting chemistry and mineralogy, make application of the pressure-quench cumulate model unlikely. If enclaves are to have originated as an accumulation of early formed crystals from the host granitoid, then they should contain a concentration of the liquidus or near-liquidus phases of the host, which in the Woodlands example are sodic plagioclase and ferro-hornblende. However, as already discussed, the primary phenocrysts of the Woodlands enclaves are magnesian pyroxenes and calcic plagioclase, which are clearly not in equilibrium with the granitoid host. Hence, this gross disequilibrium in phenocryst mineralogy between enclaves and host in the Woodlands example, also renders a coarse-grained cumulate origin (Dodge & Kistler 1990) unlikely for the Woodlands example. The occurrence of quenched margins rimming many enclaves within the Woodlands pluton also infers that these inclusions were magmatic, providing additional evidence against both a restite and cumulate origin.

Geochemical evidence Major and trace element contents of the Woodlands enclaves provide additional evidence against application the restite theory. One of the principle arguments proposed as evidence for restite unmixing is the linear geochemical variations apparent in many granitoid suites. In brief, linear geochemical trends are proposed to be the result of progressive removal of 212 entrained restite from the minimum melt component of granitoid magmas (see chapter 4 for a more detailed discussion). Examples of entrained restite include features such as plagioclase cores, mafic clots, and microgranitoid enclaves. Since removal of restite is proposed as the cause of the linear geochemical variation, then by definition, the compositions of enclaves should plot on extensions of the differentiation trends in the host. This geochemical observation is evident in some enclave - host pairs (e.g. Vernon 1983), and is used as one of the primary arguments supporting a restite origin for enclaves (Chappell et al. 1987). However, such geochemical behaviour is certainly not evident the case in the Woodlands example, with enclaves forming linear trends that are independent from those of the host granitoid on most Harker plots.

A similar argument may be made against a cumulate origin for enclaves in the Woodlands example. If fractional crystallization is the inferred differentiation process in the host granitoid, and enclaves are postulated to represent early cumulate phases, then enclaves should also plot along extensions of the fractional crystallization trends. However, if processes other than fractional crystallization (e.g. magma mixing, fractional partial melting) are the inferred differentiation mechanism for the host, then an argument against a cumulate origin for enclaves, on geochemical grounds alone, cannot be made.

Isotopic evidence During the last twenty years, isotopic investigations have frequently been applied to provide constraints on the genetic relationships between enclaves and their host granitoids (Didier & Barbarin 1991b). However, the usefulness of isotopic systems is limited to enclave-host pairs in which the enclaves retain at least some of their primary signatures during their residence in the enclosing granitoid magma (Pin 1991). Enclaves-host pairs in which enclaves and host share the same isotopic signature, may be used to invoke any of the three major models origin of microgranitoid enclaves discussed above, and thus the isotopic evidence in these cases is equivocal (e.g. Pin et al. 1990) . Restite enclaves should record the same isotopic signature as their host, since they represent the source rock from which the host granitoid gained its signature. Likewise, cumulate enclaves which originated as early crystallization products of their host granitoid, should also share the same isotopic signature as they host, since the two are cogenetic. Even in mingled magma systems, enclaves may reach isotopic equilibration with their hosts through diffusional exchange (Pin 1991), or may have had similar isotopic 213 signatures to begin with, hence limiting the usefulness of isotopic studies. Conversely, an isotopic difference between enclave and host is diagnostic of mingled magmas that were not cogenetic (Pin 1991).

The petrographic and geochemical evidence suggesting limited equilibration between enclaves and host in the Woodlands example suggests that isotopic analysis might prove useful in this case. Four samples were selected for Nd-Sr isotopic analysis, including the host granitoid, two enclaves (one phenocryst-bearing, one phenocryst-free), together with an augite separate from the phenocryst-bearing enclave. Results are summarized in table 6A1

87 86 and are presented in a conventional gNd - Sr/ Srinitial diagram in figure 6A12. In addition to the results tabulated below, a field for the Uralla Plutonic Suite, and analyses of the Wards Mistake Adamellite and the Oban River Leucoadamellite (data recalculated from Hensel et al. 1985), and a field for the Moonbi Plutonic Suite (Collins et al. unpublished data) are plotted for comparison (Fig. 6A12).

87 86 Sample gNd Sr/ Srinitial WQM (host) 0A26 0A70610

WQMX5 (phenocryst free enclave) 1A21 0A70523

WQMX2 (phenocryst bearing enclave) 0A57 0A70576

WQMX2P (augite separate from WQMX2) 1A69 0A70382

Table 6A1. Summary of isotopic results from the Woodlands Quartz monzonite and enclaves (all ratios are calculated at 249 Ma, which is the Rb-Sr biotite age for the Woodlands pluton).

The isotopic composition of the host quartz monzonite (WQM) plots within the field for the Uralla Suite (Fig. 6A12), suggesting derivation from source rocks of similar isotopic composition as that suite. The isotopic composition of the phenocryst-bearing enclave (WQMX2) also plots within this field. However, both the phenocryst-free enclave (WQMX8) and the augite separate from WQMX2 show significant isotopic differences to the host. The augite separate in particular, has an isotopic composition which plots near the mantle array, and is far removed from the composition of the host quartz monzonite. 214 . 0.710 al. 1985). et Moonbi Suite Uralla Suite monzonite quartz host phenocryst-rich enclave Leucoadamellite Oban River i Sr/ Sr 87 86 fine-grained enclave Ward's Mistake Ward's Adamellite 87 86

ntle array ma Nd vs initial Sr/ Sr plot of isotopic analyses of the Woodlands Quartz Monzonite and enclaves. Key as for figure 6 8. e

. enclave augite separate 6 12. All ratios calculated at 249 Ma

0 2 4 6 -4 -2 Figure Fields for the Uralla and Moonbi plutonic suites are included for reference (recalculated data from Hensel 0.702 0.703 0.704 0.705 0.706 0.707 0.708 0.709 Nd e 215

The basaltic character of the phenocrysts from the WQM enclaves is consistent with the observation that the isotopic composition of the augite separate plots near the mantle array, and provides additional evidence that the original magma in which the phenocrysts crystallized was mantle-derived. If the isotopic composition of the augite separate represents the original isotopic composition of the basaltic magma, then the array of isotopic composition defined by the two enclaves analysed represents a process of partial isotopic equilibration with the host. The greater degree of isotopic equilibration of the phenocryst- bearing enclave with the host is consistent with the petrographic and mineral chemistry of these enclaves. The chemistry of matrix minerals from phenocryst-bearing enclaves is much closer the host quartz monzonite mineral assemblage than for the matrix minerals from phenocryst-free enclaves (see section 6A5A2). This observation, together with the coarser- grained character of the matrix-free enclaves, suggests that the matrix of these enclaves underwent more complete equilibration with the host, during slower crystallization.

In summary, these isotopic data provide unequivocal evidence that the Woodlands enclaves are not cogenetic with their host granitoid. An origin as commingled basaltic magma is the only alternative of the three possible origins of the enclaves discussed above, an origin which is also the most consistent with the petrographic and geochemical evidence.

6A7A2 Composition of the enclave parent magma

Identification of magmas which are parental to mafic microgranitoid enclaves, remains one of the most outstanding problems in the study of enclaves (Barbarin & Didier 1991). The unusual petrography of the Woodlands enclaves, with their preserved basaltic phenocrysts, provides a unique insight into the composition of parent magma. This petrography, in conjunction with geochemical and isotopic evidence will be briefly reviewed in order to place tighter constraints on the precise nature of this magma.

Petrographic evidence The phenocryst mineral assemblage of calcic plagioclase, augite, hypersthene and titanomagnetite (± magnesian hastingsitic hornblende), provides the best evidence as to the type of magma parental to the enclaves. The presence of orthopyroxene alone rules out 216 alkaline (OIB) magmas as a possibility (Wilson 1989). Additionally, the presence of highly calcic plagioclase (~An90) further constrains the parental magma type, as calcic plagioclase is a feature peculiar to high-alumina arc basalts (Crawford et al. 1987). Therefore, the phenocryst assemblage present in the Woodlands enclaves is typical of island arc (‘calc- alkaline’) basalts (Ewart 1982). Furthermore, the absence of olivine is also suggestive of a mildly evolved arc basalt, and explains the calc-alkaline geochemical trends exhibited by the enclaves (Fig. 6A13a).

Geochemical evidence Although the primary major element compositions of mingled magmas are likely to be somewhat altered form their primary compositions (e.g. Eberz & Nicholls, see section 6A7A3), the ratios of relatively immobile trace elements may be used to infer primary magma compositions, through the use of tectonic discrimination diagrams. In addition to the AFM plot, which indicates a calc-alkaline trend for the enclaves (Fig. 6A13a), the discrimination diagrams of Meschede (1986), Mullen (1983) and Pearce & Cann (1973), all indicate a calc- alkaline or arc-basalt (HAB) geochemical signature for the more mafic of the Woodlands enclaves. However, differentiation trends within the Woodlands enclave suite cause the more felsic enclaves to plot outside the fields for arc basalts, with most of these felsic samples plotting outside all fields (Fig. 6A13). The best examples of this behaviour are the trace element plots of Ti-Zr-Y (Fig. 6A13c, Pearce & Cann 1973) and Nb-Zr-Y (Fig. 6A13d, Meschede 1986), which both show the effects of Zr-enrichment in the more felsic enclaves. This Zr-enrichment trend can also be observed on the Harker plot for Zr (Fig. 6A9), and is probably the result of xenocrystic contamination (see section 6A7A3).

High-alumina basalts (HAB, or ‘calc-alkaline basalts’) are a strongly varied group, ranging from low-K tholeiites to shoshonites (Ewart 1982), with differences between the various subdivisions primarily based on K2O content. In order to restrict the possible subdivisions to which the Woodlands enclave primary magmas belong, samples were plotted on the SiO2 vs K2O plot of Gill (1981). Although post emplacement processes (e.g. diffusion & xenocrystic contamination - see section 6A7A3) have probably enhanced the K2O contents of the Woodlands enclaves, Figure 6A13e shows that the primary magmas were probably of medium-K to high-K HAB composition. The vector indicating orthoclase xenocryst contamination shows that this process is responsible for part of the high-K character of the 217

EFeO (b) CAB = Calc-Alkaline Basalts (a) AFM Plot IAT = Island Arc Tholeiites MORB = Mid-Ocean Ridge Basalts OIA = Ocean Island Andesites OIT = Ocean Island Tholeiites

Tholeiitic TiO2/10

OIT

Calc-Alkaline

MORB

Na2O+K2O MgO IAT

(c) A,B = LKT Low Potassium Tholeiites B = OFB Ocean OIA Floor Basalts B,C = CAB Calc-Alkaline Basalts D = WPB Within plate basalts CAB Ti/100 H H MnO 10 P2O5 10

(d) AI-AII = WPA (within plate Alkaline Basalts) AII-C = WPT (within plate Tholeiites) B = P MORB (Mid-Ocean Ridge Basalts) D = N MORB (Mid-Ocean Ridge Basalts) C-D = VAB (Volcanic Arc Basalts) NbH2

D A

C AI

AII Zr YH3 5.0 B Orthoclase Dacite KO2 (WQM) C 4.0 Shoshonites High-K Andesite calc-alkaline D Basaltic Andesite 3.0 Zr/4 Y Basalt

2.0 Calc-alkaline

1.0 Low-K tholeiites 0.0 50 55 60 65 70 SiO2

(e) SiO2 - K 2 O classification

. ´´ Figure 6 13. Discrimination plots for Woodlands enclaves. (a) AFM plot, (b) MnO 10-P2 O 5 10-TiO2 /10 (Mullen 1983), (c) Zr-Y/3-Ti (Pearce & Cann 1973), (d) Zr/4-Y-Nb/2 (Meschede 1986), (e) SiO2 - K 2 O classification (Gill 1981). Symbols as for figure 6. 8. 218 more felsic, orthoclase xenocrystic enclaves. Nonetheless, it is also unlikely that even the more mafic xenocryst-free enclaves have retained the original K2O contents of the parental basalt, since matrix biotite crystallization favouring the diffusion of K2O into the enclave magma during the early stages of mingling, has probably also contributed K2O (Eberz & Nicholls 1990 - see discussion section 6A7A3). In summary, although the more mafic Woodlands enclaves plot on the boundary between medium-K and high-K HAB, the influence of post emplacement K2O enrichment processes implies that the original magmas contained lower K2O contents, and hence these magmas were probably of medium-K HAB composition.

Isotopic evidence As discussed earlier, even though the isotopic compositions of enclaves from the WQM have isotopic compositions which fall close to that of the host granitoid, rather than a basaltic composition, the augite separated from the phenocryst-bearing enclaves represents an isotopic composition closer to that of the original enclave magma. In particular, it can be

87 86 seen on the gNd - Sr/ Srinitial plot that the isotopic composition of the augite separate plots within the mantle array, indicating that the enclave parental magma was most likely of basaltic composition. In addition, the relatively low gNd value (+1A69) indicated for the parental basalt indicates that this parent magma was unlikely to be of MOR affinity, since most MOR basalts have gNd values of >+6 (Wilson 1989). In summary, although the isotopic evidence alone does not tightly constrain the original composition of the enclave magmas, it strongly supports the petrographic and geochemical evidence for an arc basalt (HAB) of medium-K composition.

6A7A3 Post-emplacement chemical evolution of enclave magmas

Mechanisms of enclave magma differentiation While large bodies of mafic magma injected into a more felsic host may take a relatively long time to chemically equilibrate, enclave-sized mafic magma globules reach thermal equilibration with their host occurs more rapidly (Fernandez & Barbarin 1991). Since the liquidus temperatures are lower for granitoid magmas than for mafic magmas, thermal equilibration results in undercooling of the mafic inclusions, resulting in rapid crystallization. Once thermal equilibration has been attained, the relative rheological properties of enclave and host undergo an inversion, and viscosities become similar, due to the advanced state of 219 crystallization of the more mafic magma (Fernandez & Barbarin 1991). Once thermal and rheological contrasts have been overcome, chemical equilibration may proceed, provided both enclave and host are still partly molten, thus allowing open-system behaviour.

Since thermal diffusion between juxtaposed magmas is typically three to five orders of magnitude larger than chemical diffusion (Sparks & Marshall 1986), then chemical equilibration of enclave and host is relatively sluggish. Like thermal equilibration, the size of the mafic inclusion is the overriding factor determining the rate of chemical equilibration with the host, since larger bodies of mafic magma have a larger volume : surface area ratio. Even so, most enclaves reach chemical equilibration in the plutonic environment, since enclave and host may remain in contact with each other in a magmatic state for long periods, in the order of 1 Ma (Orsini et al. 1991). However, the preservation of basaltic phenocrysts, and the large compositional range of enclaves in the Woodlands example, suggests that equilibration has only been partial, a point discussed in more detail below. The rate of chemical transfer between host and enclave also depends on the chemical contrast and the state of crystallization of both components. Chemical exchanges are fastest in the fully molten state, and become progressively slower with crystallization of the enclave, and is extremely slow once the enclave reaches its solidus (Fernandez & Barbarin 1991).

Chemical evolution of primary enclave magmas is generally thought to proceed by three main mechanisms, which are combinations of chemical (diffusional) and physical separation or accumulation of crystallized components as outlined by Eberz & Nicholls (1990) and Orsini et al. (1991). These three processes are:

1. Closed system fractionation/accumulation within the enclave magma 2. Liquid state differentiation by chemical diffusion between enclave and host 3. Aqueous metasomatic exchange (near- or sub-solidus)

A further mechanism of chemical evolution of enclave magmas not discussed by Eberz & Nicholls (1990) or Orsini et al. (1991), is the mechanical exchange of crystalline material between enclave and host (Barbarin & Didier 1992, Vernon 1990). This process is considered important in the Woodlands example, where xenocrysts and xenoliths of host material are common, particularly within the more felsic enclaves, and hence this process is 220 added here as a fourth mechanism of chemical evolution:

4. Mechanical exchange of crystalline material between enclave end host.

These four mechanisms will now be discussed in detail, and their relevance and contribution to the evolution of the Woodlands enclaves will be assessed.

(i) Internal fractionation & accumulation: Although some authors consider that interaction processes (chemical and physical exchanges between mafic enclaves and host) are the primary causes of variation between and within enclaves (e.g. Barbarin & Didier 1992, Blake & Koyaguchi 1991), many others infer that fractional crystallization within the mafic magmas parental to enclaves is also an important process (Eberz & Nicholls 1990, Orsini et al. 1991, Poli & Tommasini 1991).

Rapid crystallization processes, as deduced from the fine-grained character of most microgranitoid enclaves, together with the small size of enclaves, infers that fractionation, at least on the scale of individual enclaves, is unlikely to be an efficient process. In addition the fluidity required for efficient fractionation is unlikely on the enclave scale, since enclaves reach thermal and rheological equilibration with their host rapidly (Fernandez & Barbarin 1991), and are thus relatively viscous. Removal and/or accumulation of crystals on the enclave scale also introduces a spatial problem - where are the crystals removed to? Although intra-enclave fractionation has been deduced by detailed analysis of the cores and rims of individual enclaves, this process is unlikely to be efficient, and large intra-enclave silica contents such as those observed by Eberz & Nicholls (1990), are accompanied by distinct petrographic variations, with chilled margins enriched in the ferromagnesian mineral assemblage (Eberz & Nicholls 1990).

Nonetheless, fractionation is likely to be an efficient process in larger bodies of mafic magma injected into a more felsic host, since larger bodies will not undergo immediate undercooling, and will retain a low viscosity (Pitcher 1991). This process will then be recognized by large scale geochemical variation between enclaves of a particular enclave suite derived from an initially larger mafic body which has undergone fractionation after emplacement in the granitoid host, as proposed by Poli & Tommasini (1991). Such fractionation processes have 221 also been inferred to occur prior to emplacement in the host, when mantle-derived magmas were emplaced in the lower crust (Orsini et al. 1991).

The presence of the basaltic phenocryst assemblage in the Woodlands enclaves, together with their large variation in bulk composition and mineralogy, provides a unique opportunity to assess the role of larger scale fractionation/accumulation processes prior to enclave breakup. Mass-balance calculations were conducted using the Genmix petrological mixing program (Le Maitre 1981), to discern if the bulk compositional differences between the mafic (~53%

SiO2) and felsic (~59% SiO2) end-members of the Woodlands enclaves could be generated by removal or accumulation of the early crystallizing basaltic phenocrysts. Modal data for the enclaves alone (see appendix C) suggest that this is a possibility, with the phenocryst-rich enclaves having total phenocryst contents of up to 60%, while the more felsic phenocryst- poor enclaves contain <20% of these phenocrysts. The calculations were applied to the compositions of the most felsic and mafic enclaves, together with mineral analyses of the basaltic mineral assemblage augite + hypersthene + plagioclase (An90) + titanomagnetite. Although calculations initially suggested this scenario provided a good working model, with the degree of fractionation/accumulation approximately matching the differences in total phenocryst modes (~40%) between PRE and PPE, calculations failed to match the elevated

K2O contents of the PPE (WQMX5). The other major differences in modal mineralogy were then taken into account, and representative analyses for the xenocrysts of orthoclase and quartz present within the more felsic enclaves were added to the equation. Although these xenocrysts represent a separate process of enclave differentiation (i.e. bulk contamination), the residual sum of squares (see table 6A2) improved significantly with the addition of these mineral phases, indicating that the calculations approximated the observed compositions more precisely.

The modelling presented in table 6A2 demonstrates that the felsic end-member of the enclaves may be generated from the mafic end-member by the removal of ~5% each of augite and hypersthene, ~29% plagioclase and ~3% titanomagnetite, together with the addition of ~2% orthoclase and ~1% quartz as xenocrysts. In addition to the calculated degree of fractionation/accumulation (~43%) matching the approximate difference in total phenocryst abundance, the proportions of phenocryst phases used in calculations also approximates the 222 proportional modes each phase in the more mafic enclaves (see appendix C). A simple subtraction of the modal abundance of phenocryst phases (from appendix C), reveals modes that differ by of 2A8% augite, 7A7% hypersthene, 28% plagioclase and 0A5% titanomagnetite.

WQMX Augite Hypers- Plagio- Fe-Ti Ortho- Quartz WQMX Element 2 thene clase oxide clase 5 (PRE) (PPE)

SiO2 52A94 50A75 53A28 49A32 0A49 63A74 100A00 59A23

TiO2 1A23 0A88 0A35 0A02 11A10 0A23 - 1A36

Al2O3 18A27 2A47 1A27 31A54 0A14 18A48 - 15A47 3FeO 8A33 9A31 16A99 - 82A71 0A13 - 6A93 MnO 0A22 0A29 0A51 - A68 - - 0A26 MgO 3A66 15A37 26A10 - - - - 2A40 CaO 8A27 20A67 1A49 14A90 - 0A21 - 4A61

Na2O 3A58 - - 3A02 - 3A18 - 4A63

K2O 1A51 - - 0A07 - 12A32 - 2A92 Residual sum of squares = 0A00005 Degree of fractionation/accumulation 43A1% (sum of removed/accumulated phases) % Used 100 5A27 5A43 29A19 3A21 2A07 1A12 59A76 Equation: WQMX2 - augite - hypersthene - plagioclase - FeTiOxide+Orthoclase + Quartz =WQMX5

Table 6A2. Modelling results - Woodlands enclaves

Although the numerical modelling above demonstrates that the differences in bulk composition between mafic and felsic enclaves in the Woodlands example, can be largely modelled as a fractionation/accumulation process, several problems remain. Firstly, while it has been demonstrated that the modal differences (for phenocryst phases) between ‘mafic’ and ‘felsic’ enclaves closely approximates the mineral proportions used in mixing calculations, there are nonetheless small differences between these values (particularly the minor phases), suggesting that other processes may be operating. Additionally, if the differences in composition between mafic and felsic enclave are solely due to differences in the modal proportions of phenocryst phases, then this implies that the bulk matrix composition is the same for all enclaves. However, the matrix composition (as determined by broad-beam electron microprobe analysis - appendix D, figure 6A8) does vary between enclaves by as much as 5% SiO2. This variation in the matrix composition between enclaves, 223 and the gross disequilibrium between phenocrysts and matrix within enclaves, must be due to other differentiation mechanisms such as diffusional exchange between enclave and host, as discussed in the next section.

The second question arising from the above modelling, is whether the bulk differences modelled are due to accumulation or fractionation processes. Although the modal abundances of phenocrysts in most phenocryst-rich enclaves are high (~60%), and phenocrysts are often in partial contact, no distinct cumulate texture is obvious. The non- recognition of cumulate textures in the case of these enclaves, may be partly due to textural modifications such as the magmatic corrosion displayed by all phenocrysts, and the extensive uralitization developed around the margins of pyroxenes. The recognition of a cumulate texture may also be hindered by quenching, forming the finer grained doleritic texture (typical of many enclaves) developed in the matrix, preventing the development of typical intercumulus texture. Nonetheless, no preferred orientation of ‘accumulated’ plagioclase phenocrysts has been observed, suggesting that accumulation has not been extensive. However, the development of a slight positive europium anomaly (Eu/Eu* = 1A0461A10) in the phenocryst-rich enclaves (see Fig. 6A11) suggests that some accumulation of plagioclase has occurred.

(ii) Diffusional chemical exchange between enclave matrix ‘magma’ and host granitoid magma (liquid state differentiation)

If two magmas are juxtaposed in a partially crystallized state, they form a diffusion couple which may exchange chemical components (Eberz & Nicholls 1990). This diffusional exchange has been discussed in detail for many examples of granitoid host - enclave pairs (e.g. Barbarin & Didier 1992, Eberz & Nicholls 1990, Orsini et al. 1991, Poli & Tommasini 1990). This diffusion is bidirectional, with transfer from felsic host towards mafic inclusion dominant over diffusion in the opposite direction (Orsini et al. 1991). Diffusional exchange in juxtaposed magmas is dominated by transfer from felsic towards mafic magma, with SiO2,

K2O, Na2O, Rb, Li, Zr, Nb, Y and REE commonly increasing in the mafic magma (e.g. Orsini et al. 1991). Transfer from mafic towards felsic magma is more limited, involving

TiO2, Al2O3, FeO, MgO, CaO and Sr (Barbarin & Didier 1992, Orsini et al. 1991). 224

Two major processes are thought to be responsible for the chemical transfer between enclaves and their hosts (Orsini et al. 1991). The chemical exchanges that occur between juxtaposed magmas are largely ‘down-gradient’ (i.e. in the direction of lower concentration), and hence the diffusion is largely gradient driven (e.g. Barbarin & Didier 1992). However, volatiles are also normally in higher concentration in felsic magmas, and hence also will diffuse into a juxtaposed mafic magma. The addition of volatiles to the system, which also largely migrate from felsic to mafic component, favours the transfer of elements such as

SiO2, K2O, Na2O, Rb, Zr, P and Cs (e.g. Orsini et al. 1991). In addition, rapid crystallization within an undercooled mafic magma, may regulate the activities of elements which are partitioned into the crystallizing assemblage, hence accelerating gradient driven diffusion.

For example, biotite crystallization, which strongly partitions K2O, will maintain a strong compositional gradient between the remaining mafic melt and the felsic host.

Experimental interactions (Johnston & Wyllie 1988, Lesher 1986, Van Der Laan & Wyllie 1993) between mafic and felsic melts demonstrate that the alkaline elements diffuse most rapidly, with diffusion rates of D = 6×10-7 cm2/sec and D = 3×10-7 cm2/sec (Van Der Na2O K2O Laan & Wyllie 1993). In contrast, diffusion of divalent transitional cations (MgO, CaO,

-8 2 -8 2 FeO) is more sluggish (DMgO=9×10 cm /sec, DCaO=4-6×10 cm /sec) and SiO2 and Al2O3 display the slowest diffusion (D and D =(3-0A6)×10-8 cm2/sec, Van Der Laan & Wyllie SiO2 Al2O3 1993). These experimental data are in accord with compositions for many natural examples, with alkali elements commonly in similar concentrations in both enclave and host. Detailed studies of zoned enclaves (e.g. Eberz & Nicholls 1990) also suggest that alkali diffusion has far exceeded that of other elements.

The importance of diffusive mass transfer process between enclave magma and host is difficult to quantitatively assess. However, the contrasting phenocryst and matrix mineralogy of the Woodlands enclaves provides unequivocal evidence that significant diffusional chemical exchange has occurred between enclave and host. Additionally, the relative significance of diffusive transfer processes for individual elements may be appraised by comparing whole-rock, phenocryst and matrix compositions for individual enclaves. Differences in isotopic composition between enclaves and phenocrysts is also a testament to diffusional exchange between enclaves and host. 225

Petrographic evidence The dominance of anhydrous phases as the phenocryst assemblage of PRE (except for rare magnesian hastingsitic hornblende) and the absence of anhydrous ferromagnesian phases in the matrix, testify to an abrupt change in a , which was presumably effected by volatile H2O transfer into the mafic magma, once mingling with the host granitoid had been accomplished. Further, the occurrence of biotite in the matrix assemblage (and the absence of orthopyroxene) also requires that a has also been increased by diffusion of K O from the K2O 2 host into the mafic magma.

Additionally, the marked change in elemental ratios between phenocryst and matrix minerals also testifies to changes in the relative activities of other elements. Whereas Mg# (molecular Mg/[Mg+Fe] ×100) averages ~76 for the pyroxenes of the phenocryst assemblage, Mg# for the matrix minerals is highly variable, and ranges from ~54626 (see Fig. 6A3), and approaches that of the host granitoid (Mg# ~20). This stark contrast attests to a relative change in aMgO and aFeO occurring once mingling had occurred. The abrupt change in plagioclase compositions, with bytownite often rimmed by oligoclase (see Fig. 6A3), in turn implies a dramatic change in relative a and a . Although these relative activity changes cannot Na2O CaO be quantified, the (absolute) major element abundances of enclave, bulk matrix and phenocryst will be reassessed, to determine which elements have been most mobile in effecting these changes.

Geochemical evidence In addition to the contrast in mineralogical compositions between phenocryst and matrix assemblage, the net changes in enclave magma chemistry effected after mingling has occurred, may be assessed by comparing whole-rock compositions of enclaves with the bulk composition of the enclave matrix (as determined by broad-beam EDS analysis). Averaged whole-rock major element analyses for phenocryst-rich enclaves and phenocryst-poor enclaves are presented in Table 6A3, together with averaged EDS (electron microprobe) analyses for the matrix of enclaves. In addition, individual matrix EDS major element analyses were plotted on figure 6A8 with the whole-rock analyses.

The differences between whole-rock compositions and matrix composition for the PPE are relatively minor, reflecting the high proportion of matrix (>90%) in these enclaves. 226

However, comparative compositions generally show the same changes (to a smaller extent) exhibited by the PRE, with a decrease in MgO and an increase in K2O the most significant differences.

Phenocryst-rich enclaves Phenocryst-poor enclaves Element Whole-rock Bulk Whole-rock Bulk

SiO2 53A23 57A62 59A11 59A07

TiO2 1A20 0A63 1A37 1A05

Al2O3 18A21 15A215A45 15A36 3FeO 8A17 8A14 6A94 5A49 MnO 0A23 0A21 0A27 0A22 MgO 3A49 2A24 2A47 1A05 CaO 8A03 5A33 4A68 3A47

Na2O3A60 5A31 4A67 4A21

K2O1A65 3A00 2A90 5A27 Mg#43333925

XCa A71 A53 A53 A48 ASI 0A82 0A70 0A80 0A81

Table 6A3. Comparison of average bulk (whole-rock, XRF) enclave compositions and matching average matrix compositions (broad beam EDS analysis), for two phenocryst-rich enclaves (WQMX2, WQMX7) and two phenocryst-poor enclaves

(WQMX5, WQMX8). Mg# is molecular 100Mg/(Mg+Fe), XCa is molecular Ca/(Ca+Na) and ASI is the Alumina Saturation Index of Zen (1986).

Because of the intact primary phenocryst mineralogy of the PRE, whole-rock analyses for these enclaves represent values closer to the initial enclave magma composition, and hence the contrast between whole-rock and matrix compositions is greater. Unlike the PPE, matrix

SiO2 contents are significantly higher than for whole-rock values, which may in part reflect phenocryst accumulation, but also inward diffusion of SiO2 from the host. Likewise, matrix compositions for the PRE show significant increases in Na2O and K2O contents, and decreases in TiO2, Al2O3, MgO and CaO, while FeO and MnO contents are essentially unchanged.

As a result of the drop in MgO contents and unchanged 3FeO contents, Mg# for the matrix is markedly lower than for whole-rock values, a feature also reflected in Mg# for the ferromagnesian phases of the matrix. Although not accurately reflected in the EDS analyses of matrix, the diffusional loss of MgO is more pronounced in the PRE relative to the PPE, 227 as reflected in the composition of matrix hornblende and biotite compositions. Both of these matrix minerals show their lowest Mg# in the PRE, and constitute the low Mg# end of the compositional spectrum depicted in figure 6A3a.

Similarly, outward diffusion of CaO and inward diffusion of Na2O is also more pronounced in the PRE relative to PPE, with this change resulting in the dramatic difference in XCa values for the matrix versus whole-rock for the PRE. Similar to the matrix ferromagnesian phases of the PRE, matrix plagioclase compositions are also more sodic in the PRE than in PPE, and constitute the Ab-rich extreme of the matrix plagioclase filed in figure 6A3b. The significant drop in ASI values for PRE matrix composition also indicates that inward Na2O diffusion has also outstripped outward CaO diffusion, although this has also been effected by loss of Al2O3.

Although inward diffusion of K2O has probably also played a role in this exchange, the change in K2O contents is complicated by xenocrystal contamination involving host orthoclase, as discussed below.

These diffusional exchanges between enclave and host granitoid are summarized in figure 6A14. Although fractional crystallization trends are difficult to pick from diffusional exchange trends on most Harker plots, these interelement plots accentuate the subtle differences between these processes. The Mg# versus XCa plot (Fig. 6A14a) shows that change in Mg# is more accentuated by diffusional exchange processes, resulting in steeper trends on this plot. However, the drop in ASI values effected by inward diffusion of Na2O (Fig. 6A14b) provides a better contrast with fractional crystallization trends, as the latter usually cause an increase in ASI value (at least for metaluminous melts) due to the uptake of CaO by CPX or hornblende removal. In each plot, the bulk phenocryst composition is also plotted, calculated by combining individual phenocryst compositions with modes.

In summary, contrasts between the composition of phenocrysts, whole-rock and matrix of the Woodlands enclaves indicate that diffusional exchange has been a significant process in the differentiation of this suite of enclaves. Two main conclusions can be drawn from the observed data. Firstly, the more rapid exchange of Na2O and K2O relative to MgO, SiO2,

Al2O3, FeO and CaO is consistent with other empirical (e.g. Eberz & Nicholls 1990) and experimental data (e.g. Van Der Laan & Wyllie 1993). The greater degree of diffusion exchange of the latter group of elements in PRE relative to PPE in turn indicates that 228 60 (a) Bulk phenocryst 50 composition

40 Mg# FC

30 DE DE 20

Host 10

XCa 0 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 1.1 (b) Host 1.0

Bulk FC phenocryst 0.9 composition ASI DE 0.8

0.7

XMg DE 0.6 0.1 0.2 0.3 0.4 0.5 0.6 Figure 6. 14. Plots showing fractionational crystallization trends (FC) versus diffusional equilibration trends (DE) for Woodlands enclaves. Symbols are for figure 6.8 except half-filled diamonds are compositions of enclave margins as determined by EDS.

(a) XCa (molecular Ca/(Ca+Na)) vs. Mg# (100Mg/(Mg+Fe)). (b) XMg (Mg/(Mg+Fe)) vs. ASI (alumina saturation index). 229 diffusional exchange has been greatest in the PRE - a point also supported by the matrix mineral compositions in PRE which are much closer to their equivalents in the host granitoid. Given that diffusion rates are unlikely to change significantly, this more extensive diffusional exchange suggests that the PRE have been subject to these exchange processes for a longer time - a point supported by the coarser grained character of the PRE matrix.

The second major point arising from these data, is that unlike many other enclave-host pairs in which complete chemical equilibration has been accomplished, as indicated by identical mineral compositions in host and enclave (e.g. Barbarin & Didier 1992), the Woodlands enclaves have only attained partial equilibration through diffusional exchange. This conclusion is also supported by the textural disequilibrium displayed by these enclaves, and the isotopic data discussed below.

Isotopic evidence The difference in initial isotopic ratios between PRE, PPE and the augite separate from PRE (which provides a best-estimate of the isotopic composition of the original basaltic magma), suggests that diffusional exchange and equilibration of Sr isotopes has been significant for the Woodlands enclaves, while there is little change in Nd isotopes of these three samples (see Fig. 6A12). This may in part reflect the lack of a significant difference in gNd between the host and original magma composition. Nonetheless, these data support experimental data (e.g. Lesher 1986) and observations in other enclave-host pairs (e.g. Orsini et al. 1991) that Sr isotopic equilibrium is attained more rapidly than Nd isotopic equilibrium, due to the faster diffusion of Sr. The greater degree of isotopic equilibration of the PRE with the host (Fig. 6A12) is also consistent with the petrographic evidence and chemical composition of these enclaves, suggesting that the matrix of these enclaves underwent more complete equilibration with the host, during slower crystallization, as discussed above.

(iii) Metasomatism - alteration of primary magmatic (basaltic) assemblages? In addition to the volatile transfer from the host granitoid into enclaves inferred to occur during liquid state differentiation (as discussed above), a number of studies from volcanic and plutonic rocks have demonstrated that a metasomatic exchange, involving a near- or subsolidus exsolved fluid phase may also bring about major chemistry changes in enclaves (e.g. Debon 1980). However, evidence for such processes is lacking in many cases, as discussed 230 by Eberz & Nicholls (1990), and experimental constraints render such fluid exchanges unnecessary in explaining the observed data, as discussed by Orsini et al. (1991).

As discussed earlier, the switch in crystallization from anhydrous to hydrous ferromagnesian phases in the Woodlands enclaves, is strong evidence for an increase in a brought about H2O by an inward diffusion of water from the host granitoid. However, this same evidence requires that this diffusion took place prior to crystallization of the matrix phases (i.e. prior to solidification), and hence is unlikely to have involved any exsolved hydrous (pegmatitic) fluids. Nevertheless, the timing of alteration (uralitization) of the anhydrous ferromagnesian phases, is not well constrained, and may have occurred after complete solidification of the enclave magma. Petrographic evidence alone is equivocal. The marginal replacement of augite phenocrysts by fibrous actinolite occurs directly beneath a magmatic overgrowth of hornblende, and the timing of its development is unconstrained. Likewise the replacement of hypersthene by cummingtonite is accompanied by an overgrowth of biotite, and is also unconstrained. However, this process of alteration is likely to have taken place at a relatively high temperature, as evidenced by the lack of alteration of plagioclase phenocrysts, and the fresh, unaltered nature of all matrix phases within the enclaves. Additionally, the lack of evidence of pneumatolytic alteration in the host granitoid, suggests that the evolution of a separate fluid phase is unlikely to have occurred.

In contrast to the enclaves at Woodlands, the enclaves of the Smokey Cape Adamellite which bear the same primary basaltic assemblage, show a more pervasive alteration, with few relict intact primary phases. In addition, feldspars in both the enclaves and host display extensive sericitization and alteration to clays, and hydrous ferromagnesian phases (particularly biotite) show extensive chloritization, suggesting that both enclaves and host have undergone a pervasive pneumatolytic alteration in the subsolidus state.

(iv) Contamination by solid phases - xenocrysts and xenoliths (mechanical exchange)

Contamination of enclave magmas by solid (crystalline) phases derived from the host granitoid is considered as an important process of enclave evolution by many researchers (e.g. Barbarin & Didier 1992, Vernon 1990). Although such a processes may be difficult to recognize or quantify geochemically, the primary evidence for this process is petrographic. 231

Petrographic features such as orthoclase showing magmatic corrosion and overgrowth by plagioclase, quartz ocelli with rims of mafic minerals, and coarse-grained plagioclase more sodic than that of the enclave matrix, are all interpreted as xenocrysts derived from the host (Vernon 1990).

Although plagioclase xenocrysts are difficult to recognize in the Woodlands example, since enclave matrix plagioclase is similar in composition to the host, orthoclase and quartz xenocrysts are easily recognized due to their larger grain size than the enclave matrix, in which neither phase is present. Orthoclase and quartz xenocrysts are restricted to the PPE, where they are quite common, with rounded orthoclase is overgrown by oligoclase, and quartz displays reaction rims of hornblende (as discussed in detail earlier). Also, the two are occasionally found as composite xenoliths, accompanied by plagioclase, which also indicates derivation from the host.

Although these xenocrysts and xenoliths are common in the PPE, their overall abundance is low, not exceeding 5% (see appendix C). As such, their sum contribution to enclave differentiation is small, and swamped by the effects of fractional crystallization and diffusional exchange. However, those enclaves containing orthoclase xenocrysts do have significantly higher K2O contents relative to other enclaves (Fig. 6A8). Although this K2O increase may be due to diffusional exchange, it is accompanied by an increase in Ba (Fig 6A9), an element which diffuses more slowly (Van Der Laan & Wyllie 1993). Hence, this coupled increase in K2O and Ba is more likely a product of orthoclase contamination. Additionally, these orthoclase xenocrysts are chemically indistinguishable from their counterparts in the host (see earlier).

6A8 Conclusions

6A8A1 A model of enclave formation for the Woodlands example

Combined petrographic, geochemical and isotopic data provide unequivocal evidence that the enclaves present in the Woodlands Quartz Monzonite originated as from a coeval, but not cogenetic, basaltic magma, that has mingled with the host A-type quartz monzonite. Preserved geochemical and mineralogical characteristics within the phenocryst-rich enclaves 232 suggest that the parent magma was of high-alumina basalt affinity (medium- to high-K calc- alkaline). Deviations from the geochemical and petrographic characteristics common to basalts (e.g. low MgO and CaO, high Na2O and K2O, crystallization of hydrous ferromagnesian phases in the matrix), are the result of several magmatic differentiation processes which have affected these enclaves in the course of their evolution. In order to explain all the variation and characteristics present within the Woodlands enclave suite, enclave formation is viewed as a product of three stages (a modification of the model presented by Poli & Tommasini 1991), which are summarized below, and depicted in figure 6A15.

Some of the most important components of the model presented by Poli & Tommasini (1991) are the geochemical inferences. As in their model, it has been shown here, that the major cause of geochemical differentiation within the enclave suite is fractional crystallization. The Poli and Tommasini model also inferred that the major cause of geochemical variation within the host granitoid, was primarily of product of mixing of the host magma with injected mafic material. However, like the nearby Chaelundi complex A-type suite, variation within the Woodlands Quartz Monzonite is inferred here to be caused by an independent fractional crystallization trend. The efficiency and extent of fractionation processes in A-type granites is likely to swamp any variation in composition caused by magma mixing (Bonin 1991).

Stage 1 - Injection of the basaltic magma into the A-type host (Fig. 6A15a) After emplacement in the upper crust, the crystallizing Woodlands Quartz Monzonite was intruded by one or more injections of basaltic magma. These injections may have initially taken the form of fountain-like intrusions (Poli & Tommasini 1991) as depicted in figure 6A15a, or may have been dykes (Pitcher 1991). Regardless of the initial form, once injection was completed, the mafic magma remained as a discrete body within the host, and attained a more globular shape due to viscosity contrasts with the host magma.

The injected basalt may have been completely molten, or may have contained early phenocrysts at the time of emplacement into the host. Whether the basaltic phenocrysts within the Woodlands enclaves formed prior to, or after injection into the felsic host is unclear. However, the skeletal nature of most phenocrysts present in the PRE suggests that their crystallization occurred due to rapid undercooling after injection into the host granitoid, 233

Stage 1

Stage 2

Stage 3

Figure 6. 15. Schematic representation of the three stage model for formation of the Woodlands enclaves. See text for explanation. 234 which together with the near absence of phenocrysts in the PPE, implies that the basalt was essentially free of crystals at the time of injection. Nonetheless, the presence of zoned plagioclase phenocrysts which do not exhibit skeletal texture (type 2 plagioclase phenocrysts), and are the common type found in the PPE suggests that these phenocrysts may have formed prior to emplacement in the host. The weak oscillatory zoning present within these plagioclases phenocrysts, may have developed in response to lowering of pressure during ascent through the crust prior to emplacement.

Stage 2 - Differentiation of the parental basalt (Fig. 6A15b) At this initial stage, freezing of the basic magma began along the contact, and was accompanied by heating of the felsic host (Sparks & Marshall 1986). The differentiation mechanisms discussed above (section 6A7A3) also began during this stage.

At the margin of the mafic intrusion, initial exchange processes between mafic magma and host begin, with xenocryst incorporation (dominated by orthoclase and quartz) occurring during initial crystallization, increasing bulk SiO2 by as much as 5% (Platevoet & Bonin 1991). Chemical exchange between the intrusion margin and granitoid host also occurred at this time, with inward diffusion of water (crystallization of hydrous phases), Na2O and

K2O, and outward diffusion of MgO and CaO. However, diffusional exchange with the host at this stage was limited, due to rapid crystallization, resulting in the finer grained groundmass of the resulting PPE. This initial stage of quenching, together with the influx of water, also prevented crystallization of anhydrous ferromagnesian phases.

Away from the rapidly quenching margin, the core of the basaltic intrusion was shielded from the exchange processes operating at the margin. The core of the intrusion was also somewhat thermally shielded, and hence quenching was not as rapid. However, the magma was sufficiently undercooled to precipitate the crystallization of skeletal phenocrysts of calcic plagioclase, augite and hypersthene. Shielding from the exchange processes that operated at the intrusion margin, not only excluded xenocryst incorporation, but the lack of diffusional exchange also permitted the crystallization of this primary anhydrous mineral assemblage, from an unaltered magma of basaltic composition. This slower crystallization also permitted effective fractional crystallization and accumulation to occur, and enhanced compositional contrasts with the rim of the intrusion. 235

Stage 3 - Intrusion breakup, diffusional exchange and final crystallization (Fig. 6A15c) By the completion of stage 2, thermal and rheological equilibration was advanced, and the viscosity contrasts between intrusion and host were substantially lowered, allowing breakup of the initial intrusion into fragments of enclave size, assisted by magmatic flow within the host.

By this time, enclaves formed from the rim of the initial intrusion were largely crystallized, and hence underwent no further chemical or physical exchanges with the host. Physical disaggregation into enclaves during this stage, may have caused some shape adjustments (e.g. rounding). The high sphericity of the PPE, as opposed to the lobate ‘gloopy’ appearance of the PRE also suggests that these finer grained enclaves were near-solid at the time of breakup of the initial mafic intrusion. Juxtaposed PPE and PRE always occur with the former convex into the latter (Plate 6A2d), also supporting this conclusion. The resultant PPE are depicted as the lighter coloured enclaves in figure 6A14c.

In contrast, magma globules that originated from the core of the initial intrusion, although rich in phenocrysts, were still largely magmatic. Upon contact with the host granitoid, these magma globules underwent a second, more rapid stage of undercooling. However, the thermal contrast was now somewhat lower than for the initial intrusion, due to the advanced state of crystallization of the globules, and heating of the host by the mafic magma, and hence quenching is not as rapid as for the initial intrusion margin (as evidenced by the coarser grained matrix of the PRE relative to PPE). Breakup of the initial intrusion also removed the barriers to chemical exchange with the host. Rapid diffusional exchange with the host than occurred, with inward diffusion of water caused the switch to crystallization of hydrous ferromagnesian phases. Concomitant diffusional loss of CaO and MgO, and inward diffusion of Na2O and K2O, resulted in the other chemical contrasts between phenocrysts and matrix (i.e. low An plagioclase, and low Mg# hornblende and biotite). Although xenocryst incorporation is theoretically possible at this stage, no evidence of xenocrysts occur in the resultant PRE. Mechanical exchange at this stage may have been hindered by the advanced state of crystallization of both magma globule and host, by preventing free movement of crystallized phases. The resultant PRE are depicted as the darker coloured enclaves in figure 6A14c. Finally, at the end of this stage, crystallization of all enclaves and host was complete. 236

6A8A2 General implications for the origin of microgranitoid enclaves

The unique preservation of the basaltic phenocryst assemblage, and the basaltic geochemical character (<55% SiO2) of the PRE, suggests that the three stages of crystallization described above, all took place relatively rapidly. Such rapidity is supported by the medium grained, mildly porphyritic character of the host quartz monzonite. This relatively rapid crystallization of the host magma has ‘frozen in’ and preserved these enclaves before thorough chemical and rheological equilibration with the host granitic magma could be accomplished. Such cases have been described as arrested hybridism by Vernon (1990).

Preservation of enclaves per se, requires that the processes of hybridism have been ‘arrested’, since the ultimate products of the mingling, mixing and hybridism process, are blended, homogenous, hybrid magmas. The preservation of enclaves in A-type granites is rare (Bonin 1991), and so such thorough mixing and homogenization may be particularly relevant in the case of A-types. The reason for the scarcity of enclaves in A-type granites may relate to the different rheological properties of A-type magmas. A-type magmas are generally regarded as high temperature, water - poor magmas (e.g. Landenberger & Collins 1996), with viscosities more akin to those of basaltic magmas. The lack of contrast in rheological properties between host and any injected mafic magma, would favour complete mixing rather than the preservation of enclaves. As such, arresting the processes of mixing and homogenization is critical to preserving enclaves in these magmas. Evidence for such homogenization of injected mafic magmas in A-type granites is present in the nearby Chaelundi Complex, where rare small mafic clots and individual corroded crystals of calcic plagioclase, together with magnesian hypersthene and augite, occur in the mafic end-member of the A-type suite (see chapter 5, Landenberger & Collins 1996).

Although the enclaves preserved in the Woodlands Quartz Monzonite are an unusual example of arrested hybridism in A-type granites, their similarities to enclaves present in other granitoid types (particularly I-types) in the New England Batholith and elsewhere, suggest that a the model discussed above may be applied to many examples of microgranitoid enclaves. The more primitive character of the Woodlands enclaves is largely a result of arresting the hybridization process, which in contrast, has moved closer to completion in many I-type suites. Enclaves in I-type granites of the New England Batholith commonly 237 display textures which are similar to those of the Woodlands example, but they show evidence of more extensive exchange with their host granitoids. Both calcic plagioclase phenocrysts and orthoclase xenocrysts overgrown by sodic plagioclase are common in many New England I-types (Vernon 1990), and although preservation of anhydrous ferromagnesian phases is uncommon, pyroxene pseudomorphs (‘mafic clots’) composed of actinolite and/or cummingtonite are common in the Chaelundi Complex I-type (Landenberger 1988), and in I-types of the Moonbi Plutonic Suite (Landenberger & Collins, unpublished). The occurrence of these ‘mafic clots’ of secondary amphiboles in the enclaves of many I-type suites (e.g. the Moruya Suite, Keay 1993) indicates that pyroxenes were probably a common early crystallization product in the history of many enclave magmas. This petrographic evidence, and its similarity with the Woodlands example, for which hybridism has been demonstrated, suggests that magma mixing and hybridism may be a general process which should be considered for all granitoid suites. 238

CHAPTER 7. A TECTONIC SYNTHESIS.

The tectonic setting of the southern New England Fold Belt prior to the late Carboniferous is widely accepted as that of a convergent margin plate boundary, with a magmatic arc in the west, flanked by a fore-arc basin, and a subduction complex in the east (e.g. Murray et al. 1987). The character of volcanism within the belt indicates a change from intra-oceanic island arc magmatism during the Devonian, to continental margin arc magmatism by the middle Carboniferous (McPhie 1987, Roberts & Engel 1987). Magmatism along the ‘Kuttung’ arc resulted from westward directed subduction beneath the eastern margin of the Australian continent (Fig. 7A1a). Development of the accretionary prism also occurred at this time, producing early blueschist facies fabrics (D1-D2) preserved in the Tia Complex.

The late Carboniferous heralded a fundamental change in tectonic and magmatic styles within the southern New England Fold Belt (SNEFB). Westward migration of arc magmatism occurred during the middle Carboniferous (Collins 1996b). Rapid easterly migration ensued in the late Carboniferous, establishing a new arc within the original accretion complex (Tablelands Complex) of the SNEFB. This migration may have been associated with outboard growth of the accretion complex or by arc rifting (Collins 1996b). However, the rapidity of this migration suggests that subduction may have terminated due to collision of oceanic plateau or seamount with the trench (Fig. 7A1b), hence initiating a new, more easterly subduction zone. Such a collision is consistent with the intense deformation and uplift within the accretionary prism during D3.

Establishment of a new arc in the old accretion complex during the latest Carboniferous (Fig. 7A1c) was accompanied by the onset of high-T/low-P metamorphism and uplift in parts of the accretion complex. This thermal perturbation culminated with intrusion of granitoids of the Hillgrove Supersuite and gabbros and diorites of the Bakers Creek Suite (of island arc tholeiite affinity). Intrusion of these arc magmas occurred during further compressional deformation (D5), causing uplift of the entire Wollomombi Zone along its eastern margin. to late Carboniferous. See text for explanation. Figure c 3020M D- ) -D (c) 310-290Ma(D ) (b) 320-310Ma(D a 3030M D- ) -D (a) 360-320Ma(D 1(a-c). Schematic profiles of 7 the southern New England Fold Belt from the early Carboniferous . amtc arc Magmatic amtc arc Magmatic amtc arc Magmatic xicin of Extinction Extinct 5 4 3 2 1 awrh Belt Tamworth awrh Belt Tamworth Belt Tamworth OEAC BASIN FORE-ARC BASIN FORE-ARC

(Peel - Manning (Peel - Manning FORE-ARC RIDGE Fault System) Fault System) (Peel - Manning Fault System) Tablelands Complex crtoay prism accretionary hreig of Shortening Tablelands Tablelands ACCRETIONARY Complex Complex D) (D 3 nrso o Hlgoe Supersuite Hillgrove of intrusion PRISM crtoay rs and prism accretionary ute sotnn of shortening Further ca plateau Ocean r seamount or

TRENCH udcin zone subduction Trench new of Initiation 239 240

After the late Carboniferous, compressional tectonics gave way to rifting in the early Permian, with the development of early Permian basins such as the Nambucca and Manning basins within the Tablelands Complex, and the Sydney Basin further to the west (Fig. 7A1d). During this period, there was little manifestation of arc magmatism within the SNEFB. However, intrusion of the Bundarra Plutonic Suite, and extrusion of bimodal volcanics in the rift basins, probably represented inter-arc or back-arc rifting, associated with establishment of an intra-oceanic island arc further to the east (as represented by the Gympie province, Sivell & Waterhouse 1988). Evidence for an outboard Permian arc may also be found in the northern Sydney Basin, where abundant tuffs of the middle to late Permian Newcastle Coal Measures indicate an easterly source.

Compressional tectonics dominated the late Permian, with climactic deformation of the Hunter-Bowen Orogeny (Fig 7A1e). This event initially generated large-scale folding of earlier fabrics within the accretion complex (F6). On a regional scale, this folding was related to development of the Texas - Coffs Harbour Orocline and dispersal of discrete structural blocks within the Tablelands Complex. F6 folds in the Wollomombi Zone and S1 cleavage in the adjacent early Permian Nambucca Block, are both truncated by D7 ductile shear zones which represent the culmination of this deformation. D7 involved westward tilting of the entire Wollomombi Zone with up to 8 km of uplift, along mylonite zones such as the Wongwibinda-Yarrowitch Fault System. This deformation coincides with docking of the Gympie province in southern Queensland (Murray et al. 1987). As such, it is postulated that the deformation recorded in the SNEFB at this time may have ultimately been associated with collision of outboard island arc formed during the early Permian (Fig. 7A1e).

Late Permian deformation was followed by re-establishment of arc-related volcanism in the early Triassic (Fig. 7A1f), which involved minor crustal extension and intrusion of I- and A- type granites of the New England Batholith. I- and A-type granites of the Chaelundi Complex were generated at this time, in a subduction-related tectonic setting. Basaltic enclaves preserved in the nearby A-type Woodlands Quartz Monzonite, provide evidence that basaltic magmas of island arc affinity were still providing the heat source necessary for partial melting in the lower crust during the Triassic. The change from S-type to I- and A- type magmatism at this time is considered to result from dehydration and depletion of fusible components in the accretion complex during ongoing subduction. 241

(d) 290-275 Ma (D6 )

Sydney Basin Tamworth Belt Tablelands (initiation) Complex Establishment of outboard island arc (= Gympie Province) Backarc extension Extension of old accretionary prism and intrusion of Bundarra Suite (Peel - Manning Fault System) Trench

(e) 275-255 Ma (D7 )

Sydney Basin Tamworth Belt Tablelands Complex

Climactic deformation and uplift of old accretionary prism Collision of outboard island arc with Tablelands Ccomplex Hunter- Mooki Thrust (Peel - Manning Fault System) Trench

(f) 255-225 Ma (post-tectonic extension)

Sydney Basin Tamworth Belt Tablelands Complex Re-establishment of arc or backarc (?) magmatism in Tablelands Complex (Triassic I- and A-types) (Peel - Manning Fault System) Trench

Figure 7. 1(d-f). Schematic profiles of the southern New England Fold Belt from the early Permian to middle Triassic. See text for explanation. 242

REFERENCES

AITCHISON J.C. 1988. Radiolaria from the southern part of the New England Orogen, eastern Australia. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 49-60.

AITCHISON J.C. & IRELAND T.R. 1995. Age of ophiolitic rocks across the late Palaeozoic New England Orogen, New South Wales: implications for tectonic models. Australian Journal of Earth Sciences 42, 11-23.

AITCHISON J.C., IRELAND T.R., BLAKE M.C.JR. & FLOOD P.G. 1992. 530Ma zircon age for ophiolite from the new England Orogen, oldest known rocks from eastern Australia. Geology 20, 125-128.

ALLAN A.D. & LEITCH E.C. 1990. The tectonic significance of unconformable contacts at the base of Early Permian sequences, southern New England Fold Belt. Australian Journal of Earth Sciences 37, 43-49.

ALLÈGRE C.J. & MINSTER J.F. 1978. Quantitative models of trace element behaviour in magmatic processes. Earth & Planetary Science Letters 38, 1-25.

ANDREWS E.C. & MINGAYE J.C.H. 1907. The geology of the New England Plateau, with special reference to the granites of northern New England. Part IV - Petrology. Records of the Geological Survey of N.S.W. 8, 196-238.

ARCULUS R.J. 1987. The significance of source versus process in the tectonic controls of magma genesis. Journal of Volcanology and Geothermal Research 32, 1-12.

ARCULUS R.J. & RUFF L.J. 1990. Genesis of continental crust: evidence from island arcs, granulites, and exospheric processes. In Vielzeuf D. & Vidal Ph. (Eds.), Granulites and crustal evolution. NATO ASI Series. Kluwer Academic Publishers, The Netherlands. 1990. pp. 7-23.

ARTH J.G. 1976. Behaviour of trace elements during magmatic processes - a summary of theoretical models and their applications. Journal of the U.S. Geological Survey 4, 41-47.

ASTHANA D. & LEITCH E.C. 1985. Petroi Metabasalt: alkaline within-plate mafic rocks from the Nambucca Slate Belt, northeastern New South Wales. Australian Journal of Earth Sciences 32, 261-277.

ATHERTON M.P. 1993. Granite magmatism. Journal of the Geological Society 150, 1009-1023.

BACON C.R. 1986. Magmatic inclusions in silicic and intermediate volcanic rocks. Journal of Geophysical Research 91, 6091-6112.

BAILEY D.K. 1978. Continental rifting and mantle degassing. In: Neumann, E.R. & Ramberg, I.B. (Editors) Petrology and geochemistry of continental rifts. Reidel, Dordrecht, pp. 1-13.

BALL C.W., DALLWITZ W.B. & NOAKES L.C. 1948. Geological reconnaissance of the proposed hydro-electric works in Kosciusko area between Waste point and Khancoban. Records BMR Geology & Geophysics 1948/7.

BARBARIN B. 1990. Plagioclase xenocrysts and mafic magmatic enclaves in some granitoids of the Sierra Nevada Batholith, California. Journal of Geophysical Research 95, 17747-17756.

BARBARIN B. & DIDIER J. 1991. Review of the main hypotheses proposed for the genesis and evolution of mafic microgranular enclaves. In: Didier J. & Barbarin B. (Editors) Enclaves & granite petrology. Elsevier, Amsterdam - Oxford - New York - Tokyo, pp. 367-373.

BARBARIN B. & DIDIER J. 1992. Genesis and evolution of mafic microgranular enclaves through various types of interaction between coexisting felsic and mafic magmas. Transactions of the Royal Society of Edinburgh: Earth Sciences 83, 145-153. 243

BARKER F., FARMER G.L., AYUSO R.A., PLAFKER G. & LULL J.S. 1992. The 50 Ma granodiorite of the eastern Gulf of Alaska: Melting in an accretionary prism in the forearc. Journal of Geophysical Research B97, 6757-6778.

BARKER F., WONES D.R., SHARP W.N. & DESBOROUGH G.A. 1975. The Pikes Peak Batholith, Colorado Front Range, and a model for the origin of the Gabbro-Anorthosite-Syenite-Potassic Granite suite. Precambrian Research 2, 97-160.

BATEMAN P.C. & CHAPPELL B.W. 1979. Crystallization, fractionation and solidification of the Tuolumne Intrusive Series, Yosemite National Park, California. Geological Society of America Bulletin 90, 465-482.

BEAKHOUSE G.P., MCNUTT R.H. & KROGH T.E. 1988. Comparative Rb-Sr and U-Pb zircon geochronology of late- to post-tectonic plutons in the Winnipeg River Belt, northwestern Ontario, Canada. Chemical Geology 72, 337-351.

BEARD J.S., ABITZ R.J. & LOFGREN G.E. 1993. Experimental melting of crustal xenoliths from Kilbourne Hole, New Mexico and implications for the contamination and genesis of magmas. Contributions to Mineralogy & Petrology 115, 88-102.

BEARD J.S. & BORGIA A. 1989. Temporal variation of mineralogy and petrology in cognate gabbroic enclaves at Arenal Volcano, Costa Rica. Contributions to Mineralogy & Petrology 103, 110-122.

BEARD J.S. & LOFGREN G.E. 1989. Effect of water on the composition of melts of greenstone and amphibolite. Science 244, 195-197.

BÉDARD J. 1990. Enclaves from the A-type granite of the Mégantic Complex, White Mountain Magma Series: Clues to granite magmagenesis. Journal of Geophysical Research 95, 17797-17819.

BINGEN B., DEMAIFFE D., HERTOGEN J., WEIS D. & MICHOT J. 1993. K-rich calc-alkaline gneisses of Grenvillian age in SW Norway: mingling of mantle-derived and crustal components. Journal of Geology 101, 763-778.

BINNS R.A. 1966. Granitic intrusions and regional metamorphic rocks of Permian age from the Wongwibinda district, northeastern New South Wales. Journal & Proceedings of the Royal Society of N.S.W. 99, 5-36.

BINNS R.A., CHAPPELL B.W., FLOOD R.H., GUNTHORPE R.J., HOBSON E., NEILSON M.J., RANSLEY J.E. & SLADE M.J. 1967. Geological map of New England 1:250,000 - New England Tableland, southern part, with explanatory text. University of New England. NSW.

BIRCH W.D. & GLEADOW J.W. 1974. The genesis of garnet and cordierite in acid volcanic rocks: evidence from the Cerberean Cauldron, central Victoria, Australia. Contributions to Mineralogy & Petrology 45, 1-13.

BLACK L.P. & MCCULLOCH M.T. 1990. Isotopic evidence for the dependence of recurrent felsic magmatism on new crust formation: An example from the Georgetown region of Northeastern Australia. Geochimica et Cosmochimica Acta 54, 183-196.

BLAKE S. & KOYAGUCHI T. 1991. Insights into the magma mixing model from volcanic rocks. In: Didier J. & Barbarin B. (Editors) Enclaves & granite petrology. Elsevier, Amsterdam - Oxford - New York - Tokyo, pp. 403-413.

BLUNDY J.D. & SPARKS R.S.J. 1992. Petrogenesis of mafic inclusions in granitoids of the Adamello Massif, Italy. Journal of Petrology 33, 1039-1104.

BONIN B. 1991. The enclaves of alkaline anorogenic granites: an overview. In: Didier J. & Barbarin B. (Editors) Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New York, Tokyo. pp. 179-189. 244

BROOKS C.K., HENDERSON P. & RONSBO J.G. 1981. Rare earth element partitioning between allanite and glass in the obsidian of Sandy Braes, northern Ireland. Mineralogical Magazine 44, 157-160.

BROWN G.C. & FYFE W.S. 1970. The production of granitic melts during ultrametamorphism. Contributions to Mineralogy & Petrology 28, 310-318.

BROWN G.C., THORPE R.S. & WEBB P.C. 1984. The geochemical characteristics of granitoids in contrasting arcs and comments on magma sources. Journal Geological Society London 141, 413-426.

BROWNLOW J.W. 1988. Plate tectonics and the Mid-Carboniferous to Middle Triassic evolution of eastern Australia - a re-evaluation. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 172-180.

BRYANT C.J. & ARCULUS R.J. 1993. Geochemistry of the Clarence River Suite granitoids. In: Flood, P.G. & Aitchison, J.C. (eds.) New England Orogen, Eastern Australia Symposium volume, Department of Geology and Geophysics, University of New England, Armidale, NSW, Australia pp. 349-352.

BURKE W.H., DENISON R.E., HETHERINGTON E.A., KOEPNICK R.B., NELSON H.F. & OTTO J.B. 1982. Variation of seawater 87Sr/86Sr throughout Phanerozoic time. Geology 10, 516-519.

BURLINSON G.K. 1977. The Geology of the Dundurrabin - Bostobrick district, New South Wales. B.Sc. Hons thesis. University of New England, Armidale, N.S.W. (unpubl.).

BURNHAM C.W. 1979. The importance of volatile constituents. In: Yoder, H.S. (Editor) The evolution of the igneous rocks: Fiftieth anniversary perspectives. Princeton, New Jersey: Princeton University Press, pp.439-482.

CAWOOD P.A. 1982b. Structural relations in the subduction complex of the Paleozoic New England Fold Belt, eastern Australia. Journal of Geology 90, 381-392.

CAWOOD P.A. 1982c. Tectonic reconstruction of the New England Fold Belt in the Early Permian: an example of development of an oblique-slip margin. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 25-34.

CAWOOD P.A. 1983. Modal composition and detrital clinopyroxene geochemistry of lithic sandstones from the New England Fold Belt (east Australia): a Paleozoic forearc terrane. Geological Society of America Bulletin 94, 1199-1214.

CAWOOD P.A. 1984a. A geochemical study of metabasalts from a subduction complex in eastern Australia. Chemical Geology 43, 29-47.

CAWOOD P.A. 1984b. The development of the SW Pacific margin of Gondwana: Correlations between the Rangitata and New England Orogens. Tectonics 3, 539-553.

CAWOOD P.A. & LEITCH E.C. 1985. Accretion and dispersal tectonics of the southern New England Fold Belt, eastern Australia. In D.G. Howell ed. Tectonostratigraphic terranes of the Circum - Pacific region. Circum - Pacific Council Energy & Mineral Resources. Earth science series 1, 481-492.

CAWTHORN R.G. & COLLERSON K.D. 1974. The recalculation of pyroxene end-member parameters and the estimation of ferrous and ferric iron content from electron microprobe analyses. American Mineralogist 59, 1203-1208.

CHAPPELL B.W. 1966. Petrogenesis of the granites at Moonbi, New South Wales. Ph.D. thesis (unpublished) Australian National University, Canberra, ACT Australia.

CHAPPELL B.W. 1978. Granitoids from the Moonbi district, New England Batholith, eastern Australia. Journal of the Geological Society of Australia 25, 267-283.

CHAPPELL B.W. 1994. Lachlan and New England: fold belts of contrasting magmatic and tectonic development. Journal and Proceedings of the Royal Society of New South Wales 127, 47-59. 245

CHAPPELL B.W. 1996. Magma mixing and the production of compositional variation within granitoid suites: Evidence from the granites of southeastern Australia. Journal of Petrology In press.

CHAPPELL B.W. & STEVENS W.E. 1988. Origin of infracrustal (I-type) magmas. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 71-86.

CHAPPELL B.W. & WHITE A.J.R. 1974. Two contrasting granite types. Pacific Geology 8, 173-174.

CHAPPELL B.W. & WHITE A.J.R. 1992a. I- and S-type granites in the Lachlan Fold Belt. Transactions of the Royal Society of Edinburgh: Earth Sciences 83, 1-26.

CHAPPELL B.W. & WHITE A.J.R. 1992b. Restite enclaves and the restite model. In: Didier J. & Barbarin B. (Editors) Enclaves & granite petrology. Elsevier, Amsterdam - Oxford - New York - Tokyo, pp. 375- 381.

CHAPPELL B.W., WHITE A.J.R. & HINE R. 1988. Granite provinces and basement terranes in the Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences 35, 505-524.

CHAPPELL B.W., WHITE A.J.R. & WYBORN D. 1987. The importance of residual source material (restite) in granite petrogenesis. Journal of Petrology 28, 1111-1138.

CHEN Y.D., PRICE R.C. & WHITE A.J.R. 1989. Inclusions in three S-type granites from Southeastern Australia. Journal of Petrology 30, 1181-1218.

CHEN Y.D., PRICE R.C., WHITE A.J.R. & CHAPPELL B.W. 1990. Mafic inclusions from the Glenbog and Bluegum granite suites, southeastern Australia. Journal of Geophysical Research 95, 17757-17785.

CHEN Y.D. & WILLIAMS I.S. 1990. Zircon inheritance in mafic inclusions from Bega Batholith granites, southeastern Australia: An ion microprobe study. Journal of Geophysical Research 95, 17787-17796.

CHRISTIANSEN E.H. & VENCHIARUTTI D.A. 1990. Magmatic inclusions in rhyolites of the Spor Mountains Formation, Western Utah: Limitations on compositional inferences from inclusions in granitic rocks. Journal of Geophysical Research 95, 17717-17728.

CLEMENS J.D. 1988. Volume and composition relationships between granites and their lower crustal source regions: An example from central Victoria, Australia. Australian Journal of Earth Sciences 35, 445-450.

CLEMENS J.D. 1989. The importance of residual source material (Restite) in granite petrogenesis: A comment. Journal of Petrology 30, 1313-1316.

CLEMENS J.D. 1990. The granulite - granite connection. In: D. Vielzeuf & Ph. Vidal (Eds.), Granulites, and Crustal Evolution. NATO ASI Series, Kluwer Academic Publishers, The Netherlands. 1990. pp. 25-36.

CLEMENS J.D., HOLLOWAY J.R. & WHITE A.J.R. 1986. Origin of an A-type granite: experimental constraints. American Mineralogist 71, 317-324.

CLEMENS J.D. & MAWER C.K. 1992. Granitic magma transport by fracture propagation. Tectonophysics 204, 339-360.

CLEMENS J.D. & VIELZEUF D. 1987. Constraints on melting and magma production in the crust. Earth & Planetary Science Letters 86, 287-306.

CLEMENS J.D. & WALL V.J. 1981. Origin and crystallization of some peraluminous (S-type) granitic magmas. Canadian Mineralogist 19, 111-131.

CLEMENS J.D. & WALL V.J. 1988. Controls on the mineralogy of S-type volcanic and plutonic rocks. Lithos 21, 53-66. 246

COLEMAN M.L. 1979. Isotopic analysis of trace sulphur from some S- and I-type granites: heredity or environment? In Atherton, M.P. & Tarney, J. eds. Origin of granite batholiths: geochemical evidence. pp. 129-133. Shiva Publishing.

COLLINS W.J. 1991. A reassessment of the ‘Hunter-Bowen Orogeny’: Tectonic implications for the Southern New England Fold Belt. Australian Journal of Earth Sciences 38, 409-423.

COLLINS W.J. 1996a. S- and I-type granitoids of the eastern Lachlan Fold Belt: 3-component mixing, not restite unmixing. Journal and Proceedings of the Royal Society of Edinburgh In press.

COLLINS W.J. 1996b. Tectonic setting of gold deposits in eastern Australia. In: Mesothermal gold deposits: A global overview. Pub. University of Western Australia (in press).

COLLINS W.J., BEAMS S.D., WHITE A.J.R. & CHAPPELL B.W. 1982. Nature and origin of A-type granites with particular reference to southeastern Australia. Contributions to Mineralogy & Petrology 80, 189-200.

COLLINS W.J., FLOOD R.H., VERNON R.H. & SHAW S.E. 1989. The Wuluma granite, Arunta Block, central Australia: An example of in situ, near-isochemical granite formation in a granulite-facies terrane. Lithos 23, 63-83.

COLLINS W.J., OFFLER R., FARRELL T.R. & LANDENBERGER B. 1993. A revised Late Palaeozoic - Early Mesozoic tectonic history for the southern New England Fold Belt. In: Flood, P.G. & Aitchison, J.C. (eds.) New England Orogen, Eastern Australia Symposium volume, Department of Geology and Geophysics, University of New England, Armidale, NSW, Australia pp. 69-84.

COLLINS W.J., & SAWYER E.W. 1995. Pervasive magma transfer through the lower - middle crust during non - coaxial compressional deformation: an alternative to dyking. Journal of Metamorphic Geology (submitted).

COLLINS W.J. & VERNON R.H. 1991. Mid crustal contamination of granitic magmas. In Second Hutton Symposium on Granites and Related Rocks (Abstacts), B.W. Chappell (Ed.) BMR record 1991/25:p27.

COLLINS W.J. & VERNON R.H. 1992. Palaeozoic arc growth, deformation and migration across the Lachlan Fold Belt, southeastern Australia. Tectonophysics 214, 381-400.

COLLINS W.J. & VERNON R.H. 1994. A rift-drift-delamination model of continental evolution: Palaeozoic tectonic development of eastern Australia. Tectonophysics 235, 249-275.

COMPSTON W., WILLIAMS I.S., KIRSCHVINK J.L., ZANG ZICHAO & MA GUOGAN. 1992. Zircon U-Pb ages for the Early Cambrian time-scale. Journal of the Geological Society, London. 149, 171-184.

CONEY P.J. 1988. The Tasman orogenic system and the Pacific margin of Gondwana. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 11-19.

CONEY P.J., EDWARDS A., HINE R., MORRISON F. & WINDRAM D. 1990. The regional tectonics of the Tasman orogenic system, eastern Australia. Journal of Structural Geology 12, 519-543.

CONRAD W.K., NICHOLLS I.A. & WALL V.J. 1988. Water-saturated and -undersaturated melting of metaluminous and peraluminous crustal compositions at 10kb: Evidence for the origin of silicic magmas in the Taupo Volcanic Zone, New Zealand, and other occurrences. Journal of Petrology 29, 765-803.

COOPER J.A., RICHARDS J.R. & WEBB A.W. 1963. Some potassium - argon ages in New England, New South Wales. Journal of the Geological Society of Australia 10, 313-316.

COSCA M.A., ESSENE E.J. & BOWMAN J.R. 1991. Complete chemical analyses of metamorphic hornblendes:

implications for normalizations, calculated H2O activities, and thermobarometry. Contributions to Mineralogy & Petrology 108, 472-484. 247

CRAWFORD A.J., FALLOON T.J. & EGGINS S. 1987. The origin of island arc high-alumina basalts. Contributions to Mineralogy & Petrology 97, 417-430.

CREASER R.A., PRICE R.C. & WORMOLD R.J. 1991. A-type granites revisited: Assessment of a residual-source model. Geology 19, 163-166.

CREASER R.A. & WHITE A.J.R. 1991. Yardea Dacite - Large-volume, high-temperature felsic volcanism from the Middle Proterozoic of South Australia. Geology 19, 48-51.

CROSS K.C., FERGUSSON C.L. & FLOOD P.G. 1987. Contrasting structural styles in the Palaeozoic subduction complex of the southern New England Orogen, eastern Australia. In Leitch, E.C. & Scheibner, E. eds. Terrane accretion and orogenic belts. American Geophysical Union Geodynamics Series 19, 83-92.

CUDDY R.G. 1978. Internal structures and tectonic setting of part of the New England Batholith and associated volcanic rocks, northern New South Wales. Unpub. PhD. Thesis University of New England, Armidale N.S.W.

CULLERS R.L., KOCH R.J. & BICKFORD M.E. 1981. Chemical evolution of magmas in the Proterozoic terrane of the St. Francois Mountains, southeastern Missouri. 2. Trace element data. Journal of Geophysical Research 86, 10388-10401.

D’LEMOS R.S., BROWN M. & STRACHAN R.A. 1992. Granite magma generation, ascent and emplacement within a transpressional orogen. Journal of the Geological Society, London 149, 487-490.

D’LEMOS R.S., BROWN M. & STRACHAN R.A. 1992. The relationship between granite and shear zones: magma generation, ascent and emplacement within a transpressional orogen. Journal of the Geological Society, London 149, 487-490.

DAVIDSON J.P., DE SILVA S.L., HOLDEN P. & HALLIDAY A.N. 1990. Small-scale disequilibrium in a magmatic inclusion and its more silicic host. Journal of Geophysical Research 95, 17661-17675.

DAY R.W., MURRAY C.G. & WHITTAKER W.G. 1978. The eastern part of the Tasman orogenic zone. Tectonophysics 48, 327-364.

DEBON F. 1980. Genesis of three concentrically zoned granitoid plutons of Cauterets - Panticosa (French and Spanish Western Pyrenees). Geologische Rundschau 69, 107-130.

DEER W.A., HOWIE R.A., ZUSSMAN J. 1966. An introduction to the rock forming minerals. Longman Publishers.

DEPAULO D.J. 1981. Trace element and isotopic effects of combined wallrock assimilation and fractional crystallization. Earth and Planetary Science Letters 53, 189-202.

DEWEY J.F. 1988. Extensional collapse of orogens. Tectonics 7, 1123-1139.

DEWEY J.F. & BIRD J.W. 1970. Plate tectonics and geosynclines. Tectonophysics 10, 625-638.

DIDIER J. & BARBARIN B. 1991. The different types of enclaves in granites - Nomenclature. In: Didier J. & Barbarin B. (Editors) Enclaves & granite petrology. Elsevier, Amsterdam - Oxford - New York - Tokyo, pp. 19-23.

DIDIER J. & BARBARIN B. 1991b. Enclaves & granite petrology. Elsevier, Amsterdam - Oxford - New York - Tokyo, 625pp.

DIRKS P.H.G.M., HAND M., COLLINS W.J. & OFFLER R. 1992. Structural - metamorphic evolution of the Tia Complex, New England fold belt; thermal overprint of an accretion - subduction complex in a compressional back-arc setting. Journal of Structural Geology 14, 669-688. 248

DIRKS P.H.G.M., LENNOX P.G. & SHAW S.E. 1992. Implications of two Rb-Sr biotite dates for the Tia Granodiorite, southern New England Fold Belt, NSW, Australia. Australian Journal of Earth Sciences 39, 111-114.

DIRKS P.H.G.M., OFFLER R. & COLLINS W.J. 1993. Timing of emplacement and deformation of the Tia Granodiorite, southern New England Fold Belt, NSW: Implications for the metamorphic history. Australian Journal of Earth Sciences 40, 103-108.

DODGE F.C.W. & KISTLER R.W. 1990. Some additional observations on inclusions in the granitic rocks of the Sierra Nevada. Journal of Geophysical Research 95, 17841-17848.

DORAIS M.J. 1993. Pyroxene in enclaves and syenites of the Red Hill complex, New Hampshire: an ion and electron microprobe study. Contributions to Mineralogy and Petrology 114, 130-138.

DOUCE A.E.P & JOHNSTON A.D. 1991. Phase equilibria and melt productivity in the pelitic system: Implications for the origin of peraluminous granitoids and aluminous granulites. Contributions to Mineralogy and Petrology 107, 202-218.

EASTELL J. & WILLIS J.P. 1990. A low dilution fusion technique for the analysis of geological samples. 1. Method and trace element analysis. X-ray Spectrometry 19, 3-14.

EBERZ G.W. & NICHOLLS I.A. 1990. Chemical modification of enclave magma by post-emplacement crystal fractionation, diffusion and metasomatism. Contributions to Mineralogy & Petrology 104, 47-55.

EBY G.N. 1990. The A-type granitoids: A review of their occurrence and chemical characteristics and speculations on their petrogenesis. Lithos 26, 115-134.

EBY G.N. 1992. Chemical subdivision of the A-type granitoids: Petrogenetic and tectonic implications. Geology 20, 641-644.

EGGINS S. & HENSEN B.J. 1987. Evolution of mantle derived augite - hypersthene granodiorites by liquid - crystal fractionation. Barrington Tops Batholith, eastern Australia. Lithos 20, 295-310.

ELLIS D.J. 1987. Origin and evolution of granulites in normal and thickened crusts. Geology 15, 167-170.

EVANS P.R. & ROBERTS J. 1980. Evolution of central eastern Australia during the Late Palaeozoic and Early Mesozoic. Journal of the Geological Society of Australia 26, 325-340.

EVERNDEN J.F. & RICHARDS J.R. 1962. Potassium - argon ages in eastern Australia. Journal of the Geological Society of Australia 9, 1-49.

EWART A. 1982. The mineralogy and petrology of Tertiary - Recent orogenic volcanic rocks: with special reference to the andesitic - basaltic compositional range. In: Thorpe R.S. (Ed.) Andesites Pub. John Wiley & Sons. pp. 25-87.

EWART A., SCHÖN R.W. & CHAPPELL B.W. 1992. The Cretaceous volcanic-plutonic province of the central Queensland (Australia) coast - a rift related ‘calc-alkaline’ province. Transactions of the Royal Society of Edinburgh: Earth Sciences 83, 327-345.

FAGAN R.K. 1979. Deformation, metamorphism and anatexis of an Early Palaeozoic flysh sequence in northeastern Victoria. Unpub. PhD thesis, University of New England, Armidale N.S.W.

FARRELL T.R. 1988. Structural geology and tectonic development of the Wongwibinda Metamorphic Complex. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 117-124.

FARRELL T.R. 1992. Deformation, metamorphism and migmatite genesis in the Wongwibinda Metamorphic Complex. PhD thesis (unpublished), Department of Geology, University of Newcastle. 249

FAURE G. & POWELL J.L. 1972. Strontium isotope geology. Pub. Springer-Verlag.

FERGUSSON C.L. 1982a. An ancient accretionary terrain in eastern New England - evidence from the Coffs Harbour block. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp.63-70.

FERGUSSON C.L. 1982b. Structure of the Late Palaeozoic Coffs Harbour Beds, northeastern N.S.W. Journal Geological Society Australia 29, 25-40.

FERGUSSON C.L. 1984a. Tectono-stratigraphy of a Palaeozoic subduction complex in the central Coffs Harbour Block of northeastern N.S.W. Australian Journal of Earth Sciences 31, 217-236.

FERGUSSON C.L. 1984b. The Gundahl Complex of the New England Fold Belt, eastern Australia: a tectonic melange formed in a Palaeozoic subduction complex. Journal of Structural Geology 6, 257-271.

FERGUSSON C.L. 1985. Trench floor sedimentary sequences in a Palaeozoic subduction complex, eastern Australia. Sedimentary Geology 42, 181-200.

FERGUSSON C.L. 1988. Tectonostratigraphic terranes in the central part of the Coffs Harbour Block, northeastern New South Wales. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 42-48.

FERNANDEZ A.N. & BARBARIN B. 1991. Relative rheology of coeval mafic and felsic magmas: Nature of resulting interaction processes and shape and mineral fabrics of mafic microgranular enclaves. In: Didier J. & Barbarin B., Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New york, Tokyo. pp. 263-275.

FLOOD P.G. 1988a. New England Orogen: Geosyncline, Mobile Belt and Terranes. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 1-6.

FLOOD P.G. & AITCHISON J.C. 1988. Tectonostratigraphic terranes of the southern part of the New England Orogen. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 7-10.

FLOOD P.G. & FERGUSSON C.L. 1982. Tectono-stratigraphic units and structure in the Texas - Coffs Harbour region. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp.71-78.

FLOOD R.H. 1971. A study of part of the New England Bathylith. Unpub. PhD thesis, University of New England, Armidale, N.S.W.

FLOOD R.H., CRAVEN S.J., ELMES D.C., PRESTON R.J. & SHAW S.E. 1988. The Warragundi Igneous Complex: volcanic centres for the Werrie Basalt N.S.W. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 166-171.

FLOOD R.H. & SHAW S.E. 1975. A cordierite bearing granite suite from the New England Batholith, N.S.W., Australia. Contributions to Mineralogy & Petrology 52, 157-164.

FLOOD R.H. & SHAW S.E. 1977. Two “S - type” granite suites with low initial 87Sr/86Sr ratios from the New England Batholith, Australia. Contributions to Mineralogy & Petrology 61, 163-173.

FLOOD R.H. & SHAW S.E. 1991. A pressure-quench cumulate origin for microgranitoid enclaves. In: Second Hutton symposium on granites and related rocks, Abstracts. Bureau of Mineral Resources, Canberra, Australia. pp. 37.

FLOOD R.H., SHAW S.E. & CHAPPELL B.W. 1980. Mineralogical and chemical matching of plutonic and associated volcanic units, New England Batholith, Australia. Chemical Geology 29, 163-170. 250

FLOOD R.H., SHAW S.E. & FARRELL T.R. 1991. Plutonic, volcanic and metamorphic rocks of the New England Batholith. Second Hutton Symposium on granites and related rocks - Excursion Guide. Bureau of Mineral Resources, Geology & Geophysics, Record 1991/23.

FLOOD R.H. & VERNON R.H. 1978. The Cooma granodiorite, Australia: An example of in situ crustal anatexis? Geology 6, 81-84.

FLOOD R.H. & VERNON R.H. 1988. Microstructural evidence of the orders of crystallization in granitoid rocks. Lithos 21, 237-245.

FRANCE-LANORD C. & LE FORT P. 1988. Crustal melting and granite genesis during the Himalayan collision orogenesis. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh, Earth Sciences 79, 183-195.

FREUNDT A. & SCHMINCKE H. 1992. Mixing of rhyolite, trachyte and basalt magma erupted from a vertically and laterally zoned reservoir, composite flow P1, Gran Canaria. Contributions to Mineralogy and Petrology 112, 1-19.

FRIEND C.R.L. 1985. Evidence for fluid pathways through Archaean crust and the generation of the Closepet Granite, Karnataka, south India. Precambrian Research 27, 239-250.

FUJIMAKI H. 1986. Partition coefficients of Hf, Zr, and REE between zircon, apatite and liquid. Contributions to Mineralogy and Petrology 94, 42-45.

FURMAN T. & SPERA F.J. 1985. Co-mingling of acid and basic magma with implications for the origin of mafic I-type xenoliths, Field and petrochemical relations of an unusual dike complex at Eagle Lake, Sequoia National Park, California, U.S.A. Journal of Volcanology & Geothermal Research 24, 151-178.

GAMBLE J.A. 1979. Some relationships between coexisting granitic and basaltic magmas and the genesis of hybrid rocks in the Tertiary central complex of Slieve Gullion, northeast Ireland. Journal of Volcanology and Geothermal Research 5, 297-316.

GAPAIS D. 1989. Shear structures within deformed granites, Mechanical and thermal indicators. Geology 17, 1144-1147.

GAPAIS D., BALE P., CHOUKROUNE P., COBBOLD P.R., MAHJOUB Y. & MARQUER. 1987. Bulk kinematics from shear zone patterns, some field examples. Journal of Structural Geology 9, 635-646.

GARCIA D., FONTEILLES M. & MOUTTE J. 1994. Sedimentary fractionations between Al, Ti and Zr and the genesis of strongly peraluminous granites. Journal of Geology 102, 411-422.

GILL J.B. 1981. Orogenic andesites and plate tectonics. Springer-Verlag, Berlin, 389 pp.

GILLIGAN L.B., BROWNLOW J.W., CAMERON R.G. & HENLEY H.F. 1992. Dorrigo - Coffs Harbour 1:250000 Metallogenic Map and explanatory notes. Geological Survey of New South Wales, Sydney.

GLAZNER A.F. 1991. Plutonism, oblique subduction, and continental growth: An example from the Mesozoic of California. Geology 19, 784-786.

GRAHAM I.J. & KORSCH R.J. 1985. Rb-Sr geochronology of coarse grained greywackes and argillites from the Coffs Harbour Block, eastern Australia. Chemical Geology 58, 45-54.

GRAY C.M. 1984. An isotopic mixing model for the origin of granitic rocks in southeastern Australia. Earth & Planetary Science Letters 70, 47-60.

GRAY C.M. 1990. A strontium isotopic traverse across the granitic rocks of southeastern Australia: Petrogenetic and tectonic implications. Australian Journal of Earth Science 37, 331-349. 251

GRAY C.M. 1995. Discussion of ‘Lachlan and New England: fold belts of contrasting magmatic and tectonic development’ by B.W. Chappell. Journal and Proceedings of the Royal Society of New South Wales. 128, 29-32.

GREEN R. & KRIDOHARTO P. 1975. New evidence on the form of the granitic intrusions in New England, N.S.W. Journal of the Geological Society of Australia 22, 51-59.

GREEN T.H. 1994. Experimental studies of trace-element partitioning applicable to igneous petrogenesis - Sedona 16 years later. Chemical Geology 117, 1-36.

GREEN T.H. & RINGWOOD A.E. 1968a. Genesis of the Calc-alkaline rock suite. Contributions to mineralogy and petrology 18, 105-162.

GREEN T.H. & RINGWOOD A.E. 1968b. Origin of garnet phenocrysts in Calc-Alkaline rocks. Contributions to Mineralogy & Petrology 18, 163-174.

GRIFFIN T.J., WHITE A.J.R. & CHAPPELL B.W. 1978. The Moruya Batholith and geochemical contrasts between the Moruya and Jindabyne suites. Journal of the Geological Society of Australia 25, 235-247.

GRIFFIN W.L. & O’REILLY S.Y. 1986. The lower crust in eastern Australia, xenolith evidence. In Dawson, J.B.; Carswell, D.A.; Hall, J. & Wedepohl, K.H. eds. 1986. The nature of the lower continental crust. Geological Society of London Special Publication No.24 pp. 363-374.

GROMET L.P. & SILVER L.T. 1983. Rare earth element distributions among minerals in a granodiorite and their petrogenetic implications. Geochimica et Cosmochimica Acta 47, 925-939.

GROMET L.P. & SILVER L.T. 1987. REE variations across the Peninsular Ranges Batholith, Implications for batholithic petrogenesis and crustal growth in magmatic arcs. Journal of Petrology 28, 75-125.

GUST D.A., STEPHENS C.J. & GRENFELL A.T. 1993. Granitoids of the northern NEO: their distribution in time and space and their tectonic implications. In: Flood P.G. & Aitchison J.C. (Eds.) New England Orogen, eastern Australia, Symposium volume, University of New England, Armidale, NSW, Australia. pp. 565-571.

GUST D.A. & PERFIT M.R. 1987. Phase relations of a high-Mg basalt from the Aleutian Island Arc, implications for primary island arc basalts and high-Al basalts. Contributions to Mineralogy & Petrology 97, 7-18.

HALL A. 1987. Igneous petrology. Longman publishers.

HAMILTON W.B. Plate tectonics and island arcs. Geological Society of America Bulletin 100, 1503-1527.

HAND M. 1988. Structural analysis of a deformed subduction accretion complex sequence, Nowendoc, N.S.W. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 105-116.

HANSEN E.C., NEWTON R.C. & JANARDHAN A.S. 1984. Fluid inclusions in rocks from the amphibolite facies gneiss to charnockite progression in southern Karnataka, India: direct evidence concerning the fluids of granulite metamorphism. Journal of Metamorphic Geology 2, 249-264.

HANSON G.N. 1978. The application of trace elements to the petrogenesis of igneous rocks of granitic composition. Earth & Planetary Science Letters 38, 26-43.

HARLEY S.L. 1989. The origins of granulites: a metamorphic perspective. Geological Magazine 126.3, 215- 247.

HARRINGTON H.J. & KORSCH R.J. 1985a. Tectonic model for the Devonian to Middle Permian of the New England Orogen. Australian Journal of Earth Sciences 32, 163-179. 252

HARRINGTON H.J. & KORSCH R.J. 1985b. Late Permian to Cainozoic tectonics of the New England Orogen. Australian Journal of Earth Sciences 32, 181-203.

HARRINGTON H.J. & KORSCH R.J. 1987. Oroclinal bending in the evolution of the New England - Yarrol Orogen and the Moreton Basin. In: Proceedings of Pacific Rim Congress 87 pp. 797-800.

HARRISON T.M. & MCDOUGALL I. 1980. Investigations of an intrusive contact, northwest Nelson, New Zealand - 1/ Thermal, chronological and isotopic constraints. Geochimica et Cosmochimica Acta 44, 1985-2003.

HART S.R. & ALLEGRE C.J. 1980. Trace element constraints on magma genesis. In Hargraves, R.B. ed. Physics of magmatic processes. Princeton University press, 1980 pp. 121-159.

HAWTHORNE F.C. 1981. Crystal chemistry of the amphiboles. Mineralogical Society of America: Reviews in Mineralogy 9A, 1-102.

HAWTHORNE F.C. 1983. The crystal chemistry of amphiboles. Canadian Mineralogist 21, 173-480.

HAYDON R.C. 1974. The geology of the Winterbourne area, south of Armidale, New South Wales, including the petrology of the Cheyenne Complex. B.Sc. Honours thesis (unpublished), University of New England, Armidale, NSW, Australia.

HELZ R.T. 1976. Phase relations of basalts in their melting ranges at P = 5kb. Part II. Melt Compositions. H O Journal of Petrology 17, 139-193. 2

HENDERSON P. 1982. Inorganic Geochemistry. Pergamon press, Oxford.

HENSEL H.D. 1982. The mineralogy, petrology and geochemistry of granitoids and associated intrusives from the southern part of the New England Batholith. Unpubl. PhD thesis, University of New England, Armidale, N.S.W.

HENSEL H.D., COMPSTON W., CHAPPELL B.W. & TAYLOR S.W. 1981. Primitive tholeiitic intrusives in the New England Batholith. Research School of Earth Science Australian National University Annual Report pp. 199-200.

HENSEL H.D., MCCULLOCH M.T., COMPSTON W. & CHAPPELL B.W. 1982. A neodymium & strontium isotopic investigation of granitoids and possible source rocks from New England, eastern Australia. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp.193-200.

HENSEL H.D., MCCULLOCH M.T. & CHAPPELL B.W. 1985. The New England Batholith : constraints on its derivation from Nd & Sr isotopic studies of granitoids and country rocks. Geochimica et Cosmochimica Acta 49, 369-384.

HIBBARD M.J. 1991. Textural anatomy of twelve magma-mixed granitoid systems. In: Didier J. & Barbarin B. (Editors) Enclaves & granite petrology. Elsevier, Amsterdam - Oxford - New York - Tokyo, pp. 431-444.

HILDRETH W. & MOORBATH S. 1988. Crustal contributions to arc magmatism in the Andes of central Chile. Contributions to Mineralogy & Petrology 98, 455-489.

HILL M., MORRIS J. & WHELAN J. 1981. Hybrid granodiorites intruding the accretionary prism, Kodiak, Shumagin, and Sanak Islands, Southwest Alaska. Journal of Geophysical Research 86, 10569-10590.

HOLLOWAY J.R. & BURNHAM C.W. 1972. Melting relations of basalt with equilibrium water pressure less than total pressure. Journal of Petrology 13, 1-29.

HOLTZ F. & JOHANNES W. 1991. Genesis of peraluminous granites I. Experimental investigation of melt

compositions at 3 and 5kb and various H2O activities. Journal of Petrology 32, 935-958. 253

HOLTZ F. & BARBEY P. 1991. Genesis of peraluminous granites II. Mineralogy and chemistry of the Tourem Complex (Northern Portugal). Sequential melting vs. restite separation. Journal of Petrology 32, 959-978.

HUPPERT H.E. & SPARKS R.S.J. 1988a. The generation of granitic magmas by intrusion of basalt into continental crust. Journal of Petrology 29, 599-624.

HUPPERT H.E. & SPARKS R.S.J. 1988b. The fluid dynamics of crustal melting by injection of basaltic sills. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh, Earth Sciences 79, 237-243

HUPPERT H.E. & SPARKS R.S.J. 1989. Chilled margins in igneous rocks. Earth & Planetary Science Letters 92, 397-405

HUTTON D.H.W. 1988. Granite emplacement mechanisms and tectonic control, inferences from deformation studies. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 245-255

IIZUMI S. & HONMA H. 1987. Magnetic susceptibility of granitoids in the southern part of the New England Batholith. Geological report of Shimane University, Japan 6, 149-160.

ISHIGA H., LEITCH E.C., WATANABE T., NAKA T. & IWASAKI M. 1988. Radiolarian and conodont biostratigraphy of siliceous rocks from the New England Fold Belt. Australian Journal of Earth Sciences 35, 73-80.

JAMES R.S. & HAMILTON D.L. 1969. Phase relations in the system NaAlSi3O8 - KAlSi3O8 - CaAl2Si2O8 - SiO2 at 1kilobar water vapor pressure. Contributions to Mineralogy & Petrology 21, 111-141.

JOHANNES W. & HOLTZ F. 1990. Formation and composition of H2O-undersaturated granitic melts. In: J.R. Ashworth & M. Brown (Eds.) High-temperature Metamorphism and Crustal Anatexis. Unwin Hyman, Boston, Sydney, Wellington. pp. 87-104.

JOHNSON M.C. & RUTHERFORD M.J. 1989. Experimental calibration of the aluminium-in-hornblende geobarometer with application to Long Valley caldera (California) volcanic rocks. Geology 17, 837-841.

JOHNSTON A.D. & WYLLIE P.J. 1988. Interaction of granitic and basic magmas: experimental observations at

10 kbar with H2O. Contributions to Mineralogy & Petrology 98, 352-362.

JONES J.G. & VEEVERS J.J. 1982. A Cainozoic history of Australia’s southeast highlands. Journal of the Geological Society of Australia 29, 1-12.

KENT A.J.R. 1993. Palaeozoic magmatic history of the Rockvale region, southern New England Orogen, NSW, Australia. In: Flood, P.G. & Aitchison, J.C. (eds.) New England Orogen, Eastern Australia Symposium volume, Department of Geology and Geophysics, University of New England, Armidale, NSW, Australia pp. 381-389.

KENT A.J.R. 1994. Geochronology and geochemistry of Palaeozoic intrusive rocks in the Rockvale region, southern New England Orogen, New South Wales. Australian Journal of Earth Sciences 41, 365-379.

KILPATRICK J.A. & ELLIS D.J. 1992. C-type magmas: igneous charnockites and their extrusive equivalents. Transactions of the Royal Society of Edinburgh 83, 155-164.

KIMBROUGHL D.L., CROSS K.C. & KORSCH R.J. 1993. U-Pb isotopic ages for zircons from the Pola-Fogal and Nundle granite suites, southern New England Orogen. In: Flood P.G. & Aitchison J.C. (Eds.) New England Orogen, eastern Australia University of New England, Armidale, NSW, Australia, pp.403- 412.

KLEEMAN J.D. 1975. Geological age measurements using fission tracks. Atomic Energy in Australia 18(2), 3-8. 254

KLEEMAN J.D. 1988. Constraints from granitic intrusives and felsic extrusives on the tectonics of the New England Orogen in the Late Palaeozoic - Early Triassic. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 129-133.

KORSCH R.J. 1977. A framework for the Palaeozoic geology of the southern part of the New England geosyncline. Journal of the Geological Society of Australia 24, 339-355.

KORSCH R.J. 1978a. Stratigraphic and igneous units in the Rockvale - Coffs Harbour region, northeastern N.S.W. Journal & Proceedings of the Royal Society of N.S.W. 111, 13-17.

KORSCH R.J. 1978b. Regional scale thermal metamorphism overprinting low grade regional metamorphism, Coffs Harbour Block, northern N.S.W. Journal & Proceedings of the Royal Society of N.S.W. 111, 89-96.

KORSCH R.J. 1978c. Petrographic variations within thick turbidite sequence, an example from the Late Palaeozoic of eastern Australia. Sedimentology 25, 247-265.

KORSCH R.J. 1981a. Some tectonic implications of sandstone petrofacies in the Coffs Harbour association, New England Orogen, N.S.W. Journal of the Geological Society of Australia 28, 261-269.

KORSCH R.J. 1981b. Deformational history of the Coffs Harbour Block. Journal & Proceedings of the Royal Society of N.S.W. 114, 17-22.

KORSCH R.J. 1981c. Structural geology of the Rockvale Block, northern New South Wales. Journal of the Geological Society of Australia 28, 51-70.

KORSCH R.J. 1982. Early Permian tectonic events in the New England Orogen. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 35-42.

KORSCH R.J. 1984. Sandstone compositions from the New England Orogen, eastern Australia: implications for tectonic setting. Journal of Sedimentary Petrology 54, 192-211.

KORSCH R.J., ARCHER N.R. & MCCONACHY G.W. 1978. The Demon Fault. Journal & Proceedings of the Royal Society of N.S.W. 111, 101-106.

KORSCH R.J. & HARRINGTON H.J. 1981. Stratigraphic and structural synthesis of the New England Orogen. Journal of the Geological Society of Australia 28, 205-226.

KORSCH R.J. & HARRINGTON H.J. 1987. Oroclinal bending, fragmentation and deformation of terranes in the New England Orogen, eastern Australia. In Leitch, E.C. & Scheibner, E. eds. Terrane accretion and orogenic belts. American Geophysical Union Geodynamics Series 19, 129-139.

KORSCH R.J., O’BRIEN P.E., SEXTON M.J., WAKE-DYSTER K.D. & WELLS A.T. 1989. Development of Mesozoic transtensional basins in easternmost Australia. Australian Journal of Earth Sciences 36, 13-28.

LACROIX A. 1890. Sur les enclaves acides des roches volcaniques d’Auvergne. Bull. Serv. Carte Géol. Fr. 2, 25-56.

LANDENBERGER B. 1988. Petrogenesis of two contrasting granitoid suites, Chaelundi Complex, northeastern New South Wales. Unpublished Honours thesis, University of Newcastle, NSW.

LANDENBERGER B. 1991. Origin and significance of mafic enclaves in peraluminous A-type granites: Examples from northeastern New South Wales. In: B.W Chappell (ed.) “Second Hutton symposium on granites and related rocks - Abstracts” pp. 59 Publ. Bureau of Mineral Resources, Geology and Geophysics Record 1991/25, Canberra, 1991. 255

LANDENBERGER B. 1992. Origin and significance of mafic enclaves in peraluminous A-type granites: Examples from northeastern New South Wales. In “The Origin of Granites II”, Transactions of the Royal Society of Edinburgh: Earth Sciences.

LANDENBERGER B. & COLLINS W.J. 1992. Granitoid genesis in the Chaelundi Complex, northeastern New South Wales: Implications for the petrogenesis of A-type granites of the New England Batholith. Geological Society of Australia Abstracts 32, 203.

LANDENBERGER B. & COLLINS W.J. 1995. S-type granites of the Hillgrove Plutonic Suite, New England Batholith, eastern Australia: products of partial melting of an intermediate greywacke source. Proceedings of the third Hutton Symposium on the Origin of Granites, University of Maryland.

LANDENBERGER B. & COLLINS W.J. 1996. Derivation of A-type granites from a dehydrated, charnockitic lower crust: Evidence from the Chaelundi Complex, eastern Australia. Journal of Petrology 37, 145-170.

LANDENBERGER B., COLLINS W.J. & WHITFORD D.J. 1996. Metasedimentary sources for granitoids of the Hillgrove Plutonic Suite, New England Batholith: Applying Nd and Sr isotopes to refine the models. In: P.K. Seccombe & Ashley P.M. (Eds.) Centre for Isotope studies research report 1993-1994. CSIRO Sydney. pp. 98-104.

LANDENBERGER B., FARRELL T.R., OFFLER R., COLLINS W.J. & WHITFORD D.J. 1993. Tectonic implications of Rb/Sr biotite ages from the Hillgrove Plutonic Suite. In: Flood, P.G. & Aitchison, J.C. (eds.) New England Orogen, Eastern Australia Symposium volume, Department of Geology and Geophysics, University of New England, Armidale, NSW, Australia pp. 353-358.

LANDENBERGER B., FARRELL T.R., OFFLER R., COLLINS W.J. & WHITFORD D.J. 1993. Hillgrove Plutonic Suite, New England Batholith: biotite Rb-Sr chronology of its deformational and metamorphic history. In: Centre for Isotope Studies, Research Report 1991-1992, pp.136-140.

LANDENBERGER B., FARRELL T.R., OFFLER R., COLLINS W.J. & WHITFORD D.J. 1995. Tectonic implications of Rb-Sr biotite ages for the Hillgrove Plutonic Suite, New England Fold Belt, N.S.W, Australia. Precambrian Research 71, 251-263.

LAPORTE D. 1994. Wetting behaviour of physical partial melts during crustal anatexis: the distribution of hydrous silicic melts in polycrystalline aggregates of quartz. Contributions to Mineralogy & Petrology 116, 486-499.

LARSEN L.L. & SMITH E.I. 1990. Mafic enclaves in the Wilson Ridge pluton, northwestern Arizona: Implications for the generation of a calc-alkaline intermediate pluton in an extensional environment. Journal of Geophysical Research 95, 17693-17716.

LE BRETON N. & THOMPSON A.B. 1988. Fluid-absent (dehydration) melting of biotite in metapelites in the early stages of crustal anatexis. Contributions to Mineralogy & Petrology 99, 226-237.

LE MAITRE R.W. 1981. Genmix - a generalised petrological mixing program. Computers and Geosciences 7, 229-247.

LEITCH E.C. 1972. The geological development of the Bellinger-Macleay region: a study in the tectonics of the New England Fold Belt. Ph.D thesis (unpublished) University of new England, Armidale NSW, Australia.

LEITCH E.C. 1974. The geological development of the southern part of the New England Fold Belt. Journal of the Geological Society of Australia 21, 133-156.

LEITCH E.C. 1975. Plate tectonic interpretation of the Palaeozoic history of the New England Fold Belt. Bulletin of the Geological Society of America 86, 141-144.

LEITCH E.C. 1978. Structural succession in a Late Palaeozoic slate belt and its tectonic significance. Tectonophysics 47, 311-323. 256

LEITCH E.C. 1982a. Crustal development in New England. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 9-16.

LEITCH E.C. 1982b. Metamorphism and deformation at Hall’s Peak and their regional significance. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 173-177.

LEITCH E.C. 1988. The Barnard Basin and the Early Permian development of the southern part of the New England Fold Belt. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 61-67.

LEITCH E.C. & ASTHANA D. 1985. The geological development of the Thora district, northern margin of the Nambucca Slate Belt, eastern New England Fold Belt. Proceedings of the Linnean Society of N.S.W. 108, 119-140.

LEITCH E.C. & MCDOUGALL I. 1979. The age of orogenesis in the Nambucca slate belt: A K-Ar study of low-grade regional metamorphic rocks. Journal of the Geological Society of Australia 26, 111-119.

LEITCH E.C., NEILSON M.J. & HOBSON E. 1969. Dorrigo - Coffs Harbour 1:250000 geological sheet. Geological Survey of NSW.

LESHER C.E. 1986. Effects of silicate liquid composition on mineral-liquid element partitioning from soret diffusion studies. Journal of Geophysical Research 91, 6123-6141.

LINDSLEY D.H. 1983. Pyroxene thermometry. American Mineralogist 68, 477-493.

LINNEMAN S.R. & MYERS J.D. 1990. Magmatic inclusions in the Holocene rhyolites of Newberry Volcano, Central Oregon. Journal of Geophysical Research 95, 17677-17691.

LISTER G.S. & BALDWIN S.L. 1993. Plutonism and the origin of metamorphic core complexes. Geology 21, 607-610.

LISTER G.S. & SNOKE A.W. 1984. S-C Mylonites. Journal of Structural Geology 6, 617-638.

LOISELLE M.C. & WONES D.R. 1979. Characteristics and origin of anorogenic granites. Geological Society of America, Abstracts with Programs 11, 468.

LUTH W.C., JAHNS R.H. & TUTTLE O.F. 1964. The granite system at pressures of 4-10 kbar. Journal of geophysical Research 69, 759-773.

MAHOOD G. & HILDRETH W. 1983. Large partition coefficients for trace elements in high-silica rhyolites. Geochimica et Cosmochimica Acta 47, 11-30.

MARSHAK S. 1988. Kinematics of orocline and arc formation in thin-skinned orogens. Tectonics 7, 73-86.

MASON D.R. 1968. The Mornington Diorite intrusion. Third year project, University of New England, Australia (unpublished).

MASON D.R. 1986. Magmatic ferromagnesian inclusions in granitoid plagioclase cores, Barrington Tops Granodiorite, N.S.W., Australia. American Mineralogist 71, 1314-1321.

MASON D.R. & KAVALIERIS I. 1984. A preliminary note on the Barrington Tops Granodiorite, N.S.W. The University of Newcastle, N.S.W., Department of Geology Research Report Number 2.

MCCARTHY T.S. & HASTY R.A. 1976. Trace element distribution patterns and their relationship to the crystallization of granitic melts. Geochimica et Cosmochimica Acta 40, 1351-1358.

MCCULLOCH M.T. & CHAPPELL B.W. 1982. Nd isotopic characteristics of S- and I-type granites. Earth & Planetary Science Letters. 58, 51-64. 257

MCCULLOCH M.T. & WASSERBURG G.J. 1978. Sm-Nd and Rb-Sr chronology of continental crust formation. Science 200, 1003-1011.

MCDOUGALL I. & WELLMAN P. 1975. Potassium-argon ages for some Australian Mesozoic igneous rocks. Journal of the Geological Society of Australia 23, 1-9.

MCPHIE J. 1982a. The Coombadjha volcanic complex: a Late Permian cauldron, northeastern New South Wales. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 221-227.

MATTINSON J.M. 1978. Age, origin and thermal history of some plutonic rocks from the Salinian Block of California. Contributions to Mineralogy & Petrology 67, 233-245.

MESCHEDE M. 1986. A method of discriminating between different types of mid-ocean-ridge basalts and continental tholeiites with the Nb-Zr-Y diagram. Chemical Geology 56, 207-218.

METCALF R.V., SMITH E.I., WALKER J.D., REED R.C. & GONZALES D.A. 1995. Isotopic disequilibrium among commingled hybrid magmas: Evidence for a two-stage magma mixing-commingling process in the Mt. Perkins pluton, Arizona. Journal of Geology 103, 509-527.

MICHAEL P.J. 1991. Intrusion of basaltic magma into a crystallizing granitic magma chamber: The Cordillera del Paine pluton in southern Chile. Contributions to Mineralogy & Petrology 108, 396-418.

MILLER C.F. 1985. Are strongly peraluminous magmas derived from pelitic sedimentary sources? Journal of Geology 93, 673-689.

MILLER C.F., WATSON E.B. & HARRISON T.M. 1988. Perspectives on source, segregation and transport of granitoid magmas. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 135-156.

MINSTER J.F., MINSTER J.B., TREUIL M & ALLEGRE C.J. 1977. Systematic use of trace elements in igneous processes. Part III. Inverse problem of fractional crystallization processes in volcanic suites. Contributions to mineralogy & Petrology 61, 49-77.

MONTEL J.M., MARIGNAC C., BARBEY P. & PICHAVANT M. 1992. Thermobarometry and granite genesis: The Hercynian low-P, high-T Velay anatectic dome (French Massif Central). Journal of Metamorphic Geology 10, 1-15.

MONTEL J.M. & VIELZEUF D. 1996. Partial melting of metagreywackes. Part II. Composition of minerals and melts. Contributions to Mineralogy & Petrology (in press).

MORAND V.J. 1982. Structure and metamorphism in the central part of the Tia Complex. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 95-104.

MORRIS P.A. 1988b. Volcanic arc reconstruction using discriminate function analysis of detrital clinopyroxene and amphibole from the New England Fold Belt, eastern Australia. Journal of Geology 96, 299-311.

MULLEN E.D. 1983. MnO-TiO2-P2O5: a minor element discriminant for basaltic rocks of oceanic environments and its implications for petrogenesis. Earth & Planetary Science Letters 62, 53-62.

MUNKSGAARD N.C. 1988. Source of the Cooma Granodiorite, New South Wales - a possible role of fluid-rock interactions. Australian Journal of Earth Sciences 35, 363-378.

MURRAY C.G., FERGUSSON C.L., FLOOD P.G., WHITAKER W.G. & KORSCH R.J. 1987. Plate tectonic model for the Carboniferous evolution of the New England Fold Belt. Australian Journal of Earth Sciences. 34, 213-236.

NAKAMURA N. 1974. Determination of REE, Ba, Fe, Mg, Na and K in carbonaceous and ordinary chondrites. Geochimica et Cosmochimica Acta 38, 757-775. 258

NANO S.C. 1987. Geology of the Mummel River area. B.Sc. Honours thesis (unpublished), University of New England, Armidale, NSW, Australia.

NASH W.P. & CRECRAFT H.R. 1985. Partition coefficients for trace elements in silicic magmas. Geochimica et Cosmochimica Acta 49, 2309-2322.

NEILSON M.J. 1970. The petrology of the New England Batholith near Guyra, New South Wales. Unpub. PhD thesis, University of New England, Armidale, N.S.W.

NELSON D.R., CRAWFORD A.J. & MCCULLOCH, M.T. 1984. Nd-Sr isotopic and geochemical systematics in Cambrian boninites and tholeiites from Victoria, Australia. Contributions to Mineralogy & Petrology 88, 164-172.

NEWTON R.C. 1990. Fluids and melting in the Archaean deep crust of southern India. In: J.R. Ashworth & M. Brown (Eds.) High-temperature Metamorphism and Crustal Anatexis. Unwin Hyman, Boston, Sydney, Wellington. pp. 149-179.

OFFLER R. 1982b. The origin of exotic blocks in serpentinites, Peel Fault System, Glenrock Station, N.S.W. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 43-51.

OFFLER R.O. & HAND M. 1988. Metamorphism in the forearc and subduction complex sequences of the Southern New England Fold Belt. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 78-86.

OLLIER C.D. 1982a. Geomorphology and tectonics of the Armidale region. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 141-147.

OLLIER C.D. 1982b. Geomorphology and tectonics of the Dorrigo Plateau, N.S.W. Journal of the Geological Society of Australia 29, 431-435.

OLLIER C.D. 1982c. The great escarpment of eastern Australia: tectonic and geomorphic significance. Journal of the Geological Society of Australia 29, 13-23.

O’NEIL J.R., SHAW S.E. & FLOOD R.H. 1977. Oxygen and hydrogen isotope compositions as indicators of granite genesis in the New England Batholith, Australia. Contributions to Mineralogy & Petrology 62, 313-328.

ORSINI J.P., COCIRTA C. & ZORPI M.J. 1991. Genesis of mafic microgranular enclaves through differentiation of basic magmas, mingling and chemical exchanges with their host granitoid magmas. In: (Didier J. & Barbarin B., Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New york, Tokyo. pp. 445-464.

OXLEY J. 1826. Journals of two expeditions into the interior. Murray, London.

PANKHURST R.J., HOLE M.J. & BROOK M. 1988. Isotope evidence for the origin of Andean granites. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 123-133.

PATERSON S.R. & TOBISCH O.T. 1988. Using pluton ages to date regional deformations: Problems with commonly used criteria. Geology 16, 1108-1111.

PATIÑO-DOUCE A.E. 1993. Titanium substitution in biotite: an empirical model with applications to

thermobarometry, O2 and H2O barometries, and consequences for biotite stability. Chemical Geology 108, 133-162.

PATIÑO-DOUCE A.E. & BEARD J.S. 1995. Dehydration melting of biotite gneiss and quartz amphibolite from 3 to 15 kbar. Journal of Petrology 36, 707-738. 259

PATIÑO-DOUCE A.E. & JOHNSTON A.D. 1991. Phase equilibria and melt productivity in the pelitic system: implications for the origin of peraluminous granitoids and aluminous granulites. Contributions to Mineralogy & Petrology 107, 202-218.

PAUL A.B. 1984. The petrology, geochemistry and mineralogy of the Sheep Station Creek Complex and associated plutons. B.Sc. Hons thesis. Sydney University. (unpubl.).

PAVLIS T.L., MONTEVERDE D.H., BOWMAN J.R., RUBENSTONE J.L. & REASON M.D. 1988. Early Cretaceous near-trench plutonism in southern Alaska: A tonalite-trondhjemite intrusive complex injected during ductile thrusting along the Border Ranges Fault System. Tectonics 7, 1179-1199.

PEARCE J.A. & CANN J.R. 1973. Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth & Planetary Science Letters 19, 290-300.

PEARCE J.A., HARRIS N.B.W. & TINDLE A.G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology 25, 956-983.

PERCIVAL J.A. 1991. Granulite facies metamorphism and crustal magmatism in the Ashuanipi Complex, Quebec-Labrador, Canada. Journal of Petrology 32, 1261-1297.

PETFORD N., KERR R.C. & LISTER J.R. 1993. Dike transport of granitoid magmas. Geology 21, 845-848.

PHILLIPS G.N. 1980. Water activity changes across an amphibolite-granulite facies transition, Broken Hill, Australia. Contributions to Mineralogy & Petrology 75, 377-386.

PHILLIPS G.N., WALL V.J. & CLEMENS J.D. 1981. Petrology of the Strathbogie Batholith: a cordierite bearing granite suite. Canadian Mineralogist 19, 47-63.

PICHAVANT M., KONTAK D.J., HERRERA J.V. & CLARK A.H. 1988. The Miocene-Pliocene Macusani Volcanics, SE Peru. I. Mineralogy and magmatic evolution of a two-mica aluminosilicate-bearing ignimbrite suite. Contributions to Mineralogy and Petrology 100, 300-324.

PICHAVANT M., KONTAK D.J., HERRERA J.V. & CLARK A.H. 1988. The Miocene-Pliocene Macusani Volcanics, SE Peru. II. Geochemistry and origin of a peraluminous magma. Contributions to Mineralogy and Petrology 100, 300-324.

PICHAVANT M. & MONTEL J.M. 1988. Petrogenesis of a two-mica ignimbrite suite: the Macusani Volcanics, SE Peru. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 197-207.

PIN C. 1991. Sr-Nd isotopic study of igneous and metasedimentary enclaves in some Hercynian granitoids from the Massif Central, France. In: Didier J. & Barbarin B., Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New york, Tokyo. pp. 333-343.

PIN C., BINON M., BELIN J.M., BARBARIN B. & CLEMENS J.D. 1990. Origin of microgranular enclaves in granitoids: Equivocal Sr-Nd evidence from Hercynian rocks in the Massif Central (France). Journal of Geophysical Research 95, 17821-17828.

PITCHER W.J. 1991. Synplutonic dykes and mafic enclaves. In: (Didier J. & Barbarin B., Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New york, Tokyo. pp. 383-391.

PITCHER W.J. 1993. The nature and origin of granite. Blackie London, Glasgow, 321 pp.

PLATEVOET B. & BONIN B. 1991. Enclaves and the mafic-felsic associations in the Permian alkaline province of Corsica, France: Physical and chemical interactions between coeval magmas. In: (Didier J. & Barbarin B., Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New york, Tokyo. pp. 191- 204. 260

POGSON D.J. & HILYARD D. 1980. Results of isotopic age dating related to Geological Survey of New South Wales investigations, 1974-1978. Records of the Geological Survey of New South Wales 20, 251-273.

POGSON D.J. & HITCHINS B.L. 1973. New England 1:500,000 Geological sheet, Geological Survey of NSW, Sydney.

POLI G.E. & TOMMASINI S. 1991. Model for the origin and significance of microgranular enclaves in calc-alkaline granitoids. Journal of Petrology 32, 657-666.

POTTS P.J. 1992. A handbook of silicate rock analysis. Blackie, Glasgow pp. 622.

PRESNALL D.C. & BATEMAN P.C. 1973. Fusion relations in the system NaAlSi3O8 -CaAl2Si2O8 - KAlSi3O8 - H2O and generation of granitic magmas in the Sierra Nevada Batholith. Geological Society of America Bulletin 84, 3181-3202.

PRICE R.C., BROWN W.M., & WOOLARD C.A. 1983. The geology, geochemistry and origin of late-Silurian high-Si igneous rocks of the Upper Murray Valley, NE Victoria. Journal of the Geological Society of Australia 30, 443-459.

PROVOST A. 1990. An improved diagram for isochron data. Chemical Geology (Isotope Geoscience Section) 80, 85-99.

REID J.B., EVANS O.C. & FATES D.G. 1983. Magma mixing in granitic rocks of the central Sierra Nevada, California. Earth & Planetary Science Letters 66, 243-261.

REID J.B., MURRAY D.P., HERMES O.D. & STEIG E.J. 1993. Fractional crystallization in granites of the Sierra Nevada: How important is it? Geology 21, 587-590.

ROBERTS J. & ENGEL B.A. 1987. Depositional and tectonic history of the southern New England Orogen. Australian Journal of Earth Sciences 34, 1-20.

ROBERTS J., CLAOUE-LONG J.C. & JONES P.J. 1991. Calibration of the Carboniferous and Early Permian of the southern New England Orogen by SHRIMP ion microprobe zircon analysis. In: Advances in the study of the Sydney Basin. Twenty-fifth Newcastle symposium pp. 38-43. Department of Geology, University of Newcastle.

ROBERTS M.P. & CLEMENS J.D. 1993. Origin of high-potassium, calc-alkaline, I-type granitoids. Geology 21, 825-828.

ROGERS J.J.W. & GREENBERG J.K. 1990. Late-orogenic, post-orogenic, and anorogenic granites: distinction by major-element and trace-element chemistry and possible origins. Journal of Geology 98, 291-309

RUDNICK R.L. 1992. Restites, Eu anomalies, and the lower continental crust. Geochimica et Cosmochimica Acta 56, 963-970.

RUSHMER T. 1991. Partial melting of two amphibolites: Contrasting experimental results under fluid-absent conditions. Contributions to Mineralogy & Petrology 107,41-59.

RYERSON F.J. & HESS P.C. 1978. Implications of liquid-liquid distribution coefficients to mineral-liquid partitioning. Geochimica et Cosmochimica Acta 42, 921-932.

SANDIFORD M. 1989. Secular trends in the thermal evolution of metamorphic terrains. Earth and Planetary Science Letters 95, 85-96.

SANDIFORD M. & POWELL R. 1990. Some isostatic and thermal consequences of the vertical strain geometry in convergent orogens. Earth and Planetary Science Letters 98, 154-165.

SAWKA W.N. 1988. REE and trace element variations in accessory minerals and hornblende from the strongly zoned McMurry Meadows Pluton, California. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 157-168. 261

SAWKA W.N., BANFIELD J.F. & CHAPPELL B.W. 1986. A weathering-related origin of widespread monazite in S-type granites. Geochimica et Cosmochimica Acta 50, 171-175.

SAWYER, E.W. 1994. Melt segregation in the continental crust. Geology 22, 1019-1022.

SAXENA S.K. & HOLLANDER N.B. 1969. Distribution of iron and magnesium in coexisting biotite, garnet, and cordierite. American Journal of Science 267, 210-216.

SCHEIBNER E. 1985. Suspect terranes in the Tasman Fold Belt System, east Australia. In D.G. Howell ed. Tectonostratigraphic terranes of the Circum - Pacific region. Circum - Pacific Council Energy & Mineral Resources. Earth Science Series 1, 493-514.

SCHMIDT M.W. 1992. Amphibole composition in tonalite as a function of pressure: an experimental calibration of the Al-in-hornblende barometer. Contributions to Mineralogy & Petrology 110, 304-310.

SCHOCK H.H. 1977. Trace element partitioning between phenocrysts of plagioclase, pyroxenes and the host pyroclastic matrix. Journal of Radioanalytical Chemistry 38, 327-340.

SEVIGNY J.H., PARRISH R.R. & GHENT E.D. 1989. Petrogenesis of peraluminous granites, Monashee Mountains, southeastern Canadian Cordillera. Journal of Petrology 30, 557-581.

SHAW S.E. 1964. The petrology of portion of the New England Batholith, New South Wales. Ph.D. thesis (unpublished). University of New England, Armidale, New South Wales, Australia.

SHAW S.E. & FLOOD R.H. 1981. The New England Batholith eastern Australia : geochemical variations in time and space. Journal of Geophysical Research 86, 10530-10544.

SHAW S.E. & FLOOD R.H. 1982. The Bundarra Plutonic Suite: a summary of new data. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 183-192.

SHAW S.E. & FLOOD R.H. 1993. A compilation of Late Permian and Triassic biotite Rb-Sr data from the New England Batholith and areas to the southeast. In: Centre for Isotope Studies, Research Report. Centre For Isotope Studies, CSIRO Mineral Research Laboratories, North Ryde, NSW. Australia pp. 151-155.

SHAW S.E., FLOOD R.H. & VERNON R.H. 1988. Structure of the Dundee Ignimbrite - Dundee, N.S.W. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 150-156.

SILVER L.T. & CHAPPELL B.W. 1988. The Peninsular Ranges Batholith: an insight into the evolution of the Cordilleran batholiths of southwestern North America. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 105-121.

SIMPSON C. 1986. Determination of movement sense in mylonites. Journal of Geological Education 34, 246-261.

SIMPSON C. & SCHMIDT S.M. 1983. Microstructural indicators of sense of shear in shear zones. Geological Society of America Bulletin 94, 1281-1283.

SIVELL W.J. & WATERHOUSE J.B. 1988. Petrogenesis of Gympie Group volcanics: evidence for remnants of an Early Permian volcanic arc in eastern Australia. Lithos 21, 81-95.

SKJERLIE K.P. & JOHNSTON A.D. 1992. Vapor-absent melting at 10kbar of a biotite- and amphibole-bearing tonalitic gneiss: Implications for the generation of A-type granites. Geology 20, 263-266.

SKJERLIE K.P. & JOHNSTON A.D. 1993. Fluid-absent melting behaviour of a F-rich tonalitic gneiss at mid- crustal pressures: Implications for the generation of anorogenic granites. Journal of Petrology 34, 785- 815. 262

SKJERLIE K.P., PATIÑO DOUCE, A.E. & JOHNSTON, A.D. 1993. Fluid absent melting of a layered crustal protolith: implications for the generation of anatectic granites. Contributions to Mineralogy & Petrology 114, 365-378.

SPARKS R.S.J. & MARSHALL L.A. 1986. Thermal and mechanical constraints on mixing between mafic and silicic magmas. Journal of Volcanology & Geothermal Research 29, 99-124.

STEIGER R.H. & JAGER E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmo-chronology. Earth & Planetary Science Letters 36, 359-362.

STEPHENS C.J., SCHÖN R.W. & EWART A. 1993. Mesozoic crustal extension in the northern New England Orogen: Geochemical and isotopic evidence from large scale silicic magmatism. In: Flood, P.G. & Aitchison, J.C. (eds.) New England Orogen, Eastern Australia Symposium volume, Department of Geology and Geophysics, University of New England, Armidale, NSW, Australia pp. 637-642.

STEPHENSON N.C.N. & HENSEL H.D. 1982. Amphibolites and related rocks from the Wongwibinda metamorphic complex, northern New South Wales, Australia. Lithos 15, 59-75.

STEVENS G. & CLEMENS J.D. 1993. Fluid-absent melting and the roles of fluids in the lithosphere: a slanted summary? Geology 108, 1-17.

STIMAC J.A., PEARCE T.H., DONELLY-NOLAN J.M. & HEARN B.C.JR. 1990. The origin and implications of undercooled andesitic inclusions in rhyolites, Clear Lake Volcanics, California. Journal of Geophysical Research 95, 17729-17746.

STRECKHEISEN A.L. 1973. Plutonic rocks: classification and nomenclature recommended by the IUGS Subcommission on the systematics of igneous rocks. Geotimes 18, 26-30.

STÜWE K., SANDIFORD M. & POWELL R. 1993. Episodic metamorphism and deformation in low-pressure, high temperature terranes. Geology 21, 829-832.

SUN S.S. & MCDONOUGH W.E. 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A.D. & Norry, M.J. (eds.) Magmatism in the ocean basins. Geological Society Special Publication 42, 313-345.

SUTTON J. & WATSON J.V. 1986. Architecture of the continental lithosphere. Philosophic Transactions of the Royal Society of London A317, 5-12.

SYLVESTER P.J. 1989. Post-collisional alkaline granites. Journal of Geology. 97, 261-280.

TAIT R.E. & HARLEY S.L. 1988. Local processes involved in the generation of migmatites within mafic granulites. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 209-222.

TARNEY J. & JONES C.E. 1994. Trace element geochemistry of orogenic igneous rocks and crustal growth models. Journal of the Geological Society, London 151, 855-868.

TAYLOR H.P.JR. 1988. Oxygen, hydrogen, and strontium isotope constraints on the origin of granites. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 317-338.

THIRLWALL M.F., SMITH T.E., GRAHAM A.M., THEODOROU N., HOLLINGS P., DAVIDSON J.P. & ARCULUS R.J. 1994. High field strength element anomalies in arc lavas: Source or process? Journal of Petrology 35, 819-838.

THOMPSON A.B. 1982. Dehydration melting of pelitic rocks and the generation of H2O-undersaturated granitic liquids. American Journal of Science 282, 1567-1595. 263

THOMPSON J. 1973. Results of radiometric dating programme, 1971-1973. Records of the Geological Survey of N.S.W. 16, 239-244.

THOMPSON P.H. 1989. Moderate overthickening of thinned sialic crust and the origin of granitic magmatism and regional metamorphism in low-P-high-T terranes. Geology 17, 520-523.

TINDLE A.G. 1991. Trace element behaviour in microgranular enclaves from granitic rocks. In: Didier J. & Barbarin B. (Editors) Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New York, Tokyo. pp. 313-331.

TINDLE A.G. & PEARCE J.A. 1983. Assimilation and partial melting of continental crust: evidence from the mineralogy and geochemistry of autoliths and xenoliths. Lithos 16, 185-202.

TULLIS J. 1983. Deformation of feldspars. In: Ribbe, P.H. (Ed.): Reviews in Mineralogy Vol. 2, Mineralogical Society of America, pp. 297-323.

TURNER F.J. 1981. Metamorphic Petrology - Mineralogical, field and tectonic aspects. 2nd edition. Pub. Hemisphere Publishing Corp.

TURNER S.P., FODEN J.D. & MORRISON R.S. 1992. Derivation of some A-type magmas by fractionation of basaltic magma: An example from the Padthaway Ridge, South Australia. Lithos 28, 151-179.

TUTTLE O.F. & BOWEN N.L. 1958. Origin of granite in the light of experimental studies in the system

NaAlSi3O8-KAlSi3O8-SiO2-H2O. Memoirs of the Geological Society of America. 74, 153 pp.

VAN DER LAAN S.R. & WYLLIE P.J. 1993. Experimental interaction of granitic and basaltic magmas and implications for mafic enclaves. Journal of Petrology 34, 491-517.

VEEVERS J.J. 1989. Middle/Late Triassic (230±5) singularity in the stratigraphic and magmatic history of the Pangaean heat anomaly. Geology 17, 784-787.

VEEVERS J.J., CONAGHAN P.J. & SHAW S.E. 1993. Permian and Triassic New England Orogen/Bowen- Gunnedah-Sydney Basin in the context of Gondwanaland and Pangea. In: Flood P.G. & Aitchison J.C. (Eds.) New England Orogen, Eastern Australia. University of New England, Armidale, NSW. pp.31- 51.

VERNON R.H. 1982. Isobaric cooling of two regional metamorphic complexes related to igneous intrusions in southeastern Australia. Geology 10, 76-81.

VERNON R.H. 1983. Restite, xenoliths and microgranitoid enclaves in granites. Journal and Proceedings of the Royal Society of N.S.W. 116, 77-103.

VERNON R.H. 1990. Crystallization and hybridism in microgranitoid enclave magmas: Microstructural evidence. Journal of Geophysical Research 95, 17849-17859.

VERNON R.H. 1991. Interpretation of microstructures of microgranitoid enclaves. In: Didier J. & Barbarin B. (Editors) Enclaves and granite petrology. Elsevier Amsterdam, Oxford, New York, Tokyo. pp. 277- 291.

VERNON R.H. & COLLINS W.J. 1988. Igneous microstructures in migmatites. Geology 16, 1126-1129.

VERNON R.H., ETHERIDGE M.A. & WALL, V.J. 1988. Shape and microstructure of microgranitoid enclaves: indicators of magma mingling and flow. Lithos 22, 1-11.

VERNON R.H. & FLOOD R.H. 1982. Some problems in the interpretation of microstructures in igneous rocks. In Flood, P.G. & Runnegar, B. eds., New England Geology, Voisey Symposium Volume, University of New England, Armidale, N.S.W., pp. 201-210.

VERNON R.H. & FLOOD R.H. 1988. Contrasting deformation of S- and I-type granitoids in the Lachlan Fold Belt, eastern Australia. Tectonophysics 147, 127-143. 264

VIELZEUF D., CLEMENS J.D., MOINET E. & MONTEL J.M. 1990. Experimental determination of the fluid-absent melting reactions in Al-metagreywackes. Third International Symposium on Experimental Mineralogy, Petrology and Geochemistry, Edinburgh, April 1990.

VIELZEUF D., CLEMENS J.D., PIN C. & MOINET E. 1990. Granites, granulites, and crustal differentiation. In: D. Vielzeuf & Ph. Vidal (Eds.), Granulites and Crustal Evolution. NATO ASI Series1990. Kluwer Academic Publishers, Netherlands. pp. 59-85.

VIELZEUF D. & HOLLOWAY J.R. 1988. Experimental determination of the fluid-absent melting relations in the pelitic system. Contributions to Mineralogy and Petrology 98, 257-276.

VIELZEUF D. & MONTEL J.M. 1994. Partial melting of metagreywackes. Part I. Fluid absent experiments and phase relationships. Contributions to Mineralogy & Petrology 117, 375-393.

VOISEY A.H. 1942. The geology of the County of Sandon, N.S.W. Proceedings of the Linnean Society of N.S.W. 67, 288-293.

VOISEY A.H. 1959. Australian geosynclines. The Australian Journal of Science. 22(5), 188-198.

WAGNER G.A., REIMER G.M. & JÄGER E. 1977. Cooling ages derived by apatite fission-track, mica Rb-Sr and K-Ar dating: the uplift and cooling history of the Central Alps. Memoirs of the Institute of Geology and Mineralogy, University of Padova XXX Volume 28.

WALL V.J., CLEMENS J.D. & CLARKE D.B. 1987. Models for granitoid evolution and source compositions. Journal of Geology 95, 731-749.

WASS S.Y. 1979. Multiple origins of clinopyroxenes in alkali basaltic rocks. Lithos 12, 115-132.

WATANABE T., IWASAKI M., ISHIGA H., ILIZUMI S., HONMA H., KAWACHI Y., MORRIS P., NAKA T. & ITAYA T. 1988. Late Carboniferous Orogeny in the southern New England Fold Belt, NSW, Australia. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 93-98.

WATERHOUSE J.B. 1988. The possibility that various so-called Early Permian rocks and faunas are really Late Carboniferous - Significance for dating volcanic and tectonic events. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 87-92.

WATERHOUSE J.B. & SIVELL W.J. 1987. Permian evidence for trans-Tasman relationships between east Australia, New Caledonia and New Zealand. Tectonophysics 142, 227-240.

WATERHOUSE J.B. & SIVELL W.J. 1988. New England Orogen and the Tasman Sea. In Kleeman, J.D. ed., New England Orogen - Tectonics and Metallogenesis, Symposium Volume, University of New England, Armidale, N.S.W., pp. 199-203.

WATSON E.B. 1979. Zircon saturation in felsic liquids: Experimental results and applications to trace element geochemistry. Contributions to Mineralogy and Petrology 70, 407-419.

WATSON E.B. 1982. Melt infiltration and magma evolution. Geology 10, 236-240.

WATSON E.B. & BRENAN J.M. 1987. Fluids in the lithosphere, 1. Experimentally-determined wetting

characteristics of CO2-H2O fluids and their implications for fluid transport, host-rock physical properties, and fluid inclusion information. Earth & Planetary Science Letters 85, 497-515.

WELLMAN P. 1979. On the Cainozoic uplift of the southeastern Australian highland. Journal of the Geological Society of Australia 26, 1-9.

WETZEL K., REMER M. & HIRSCH K. 1989. Minor element effects of combined partial melting and crystallization. Earth & Planetary Science Letters 93, 142-150. 265

WHALEN J.B., CURRIE K.L. & CHAPPELL B.W. 1987. A-type granites: geochemical characteristics, discrimination and petrogenesis. Contributions to Mineralogy & Petrology 95, 407-419.

WHITE A.J.R. 1979. Sources of granite magmas. Geological Society of America Abstracts with Programs 11, 539.

WHITE A.J.R. & CHAPPELL B.W. 1977. Ultrametamorphism and granitoid genesis. Tectonophysics 43, 7-22.

WHITE A.J.R. & CHAPPELL B.W. 1983. Granitoid types and their distribution in the Lachlan Fold Belt, southeastern Australia. Geological Society of America Memoir 159 pp. 21-34.

WHITE A.J.R. & CHAPPELL B.W. 1988. Some supracrustal (S-type) granites of the Lachlan Fold Belt. In “The Origin of Granites”, Transactions of the Royal Society of Edinburgh: Earth Sciences 79, 169-181.

WHITE S.H., BRETAN P.G. & RUTTER E.H. 1986. Fault-zone reactivation: kinematics and mechanisms. Philosophical Transactions of the Royal Society of London A317, 81-97.

WHITNEY J.A. 1988. The origin of granite: The role of source water in the evolution of granitic magmas. Geological Society of America Bulletin 100, 1886-1897.

WHITNEY P.R. 1992. Charnockites and granites of the western Adirondacks, New York, USA: A differentiated A-type suite. Precambrian Research 57, 1-19.

WICKHAM S.M. 1987. Crustal anatexis and granite petrogenesis during low-pressure regional metamorphism: The Trois Seigneurs Massif, Pyrenees, France. Journal of Petrology 28, 127-169.

WILKINSON J.F.G. 1969. Intrusive rocks: The New England Batholith, The geology of New South Wales. Journal of the Geological Society of Australia 16, 271-278.

WILSON M. 1989. Igneous Petrogenesis. Unwin Hyman. London, Sydney, Boston, Wellington. 466pp.

WISE D.U., DUNN D.E., ENGELDER J.T., GEISER P.A., HATCHER R.D., KISH S.A., ODOM A.L. & SCHAMEL S. 1984. Fault-related rocks: Suggestions for terminology. Geology 12, 391-394.

WORMALD R.J. & PRICE R.C. 1988. Peralkaline granites near Temora, southern New South Wales: Tectonic and petrological implications. Australian Journal of Earth Sciences 35, 209-222.

WYBORN D. 1981. Three S-Type volcanic suites from the Lachlan Fold Belt, southeast Australia. Journal of Geophysical Research 86, 10335-10348.

WYBORN D., TURNER B.S. & CHAPPELL B.W. 1987. The Boggy Plain Supersuite: a distinctive belt of I-type igneous rocks of potential significance in the Lachlan Fold Belt. Australian Journal Earth Science 34, 21-44.

ZEN E-AN. 1986a. Phase relations of peraluminous granitic rocks and their petrogenetic implications. Annual Review Earth and Planetary Sciences 16, 21-51.

ZEN E-AN. 1986b. Aluminium enrichment in silicate melts by fractional crystallization: some mineralogic and petrographic constraints. Journal of Petrology 27, 1095-1117.

ZORPI M.J., COULON C, ORSINI J.B. & COCIRTA C. 1989. Magma mingling, zoning and emplacement in calc- alkaline granitoid plutons. Tectonophysics 157, 315-329. 266

APPENDICES

Appendix A: Analytical methods. Appendix B: Structural data. Appendix C: Modal data for granitoids. Appendix D: Mineral analyses (Electron-microprobe). Appendix E: Whole-rock geochemical analyses (XRF and INAA). Appendix F: Isotopic data. Appendix G: Catalogued sample list and grid references. 267

Appendix A: Analytical methods.

(i) Structural Data

Structural data collected in the field were analysed by both basic mapping techniques (plotting of strike trends on form-surface maps as presented in Chapter 2 and Map 1 in the rear pocket), and by equal area stereographic projection (Schmidt net plots - a presented in Chapter 2). Stereographic plots were produced by the computer program ‘GeOrient’, which was written by Rod Holcombe at the Department of Earth Sciences at the University of Queensland.

(ii) Modal Data for granitoids

Modal data in the case of chapters 5 and 6 were collected by point counting of sawn slabs (1000 points) stained for alkali feldspar (sodium cobaltinitrite) in the case of coarse grained granitoids, which was supplemented by thin section point counts (1000 points) for the mafic mineralogy. Modal data for the Woodlands enclaves (Chapter 6) were collected by thin- section point counts only (1000 point).

Modal data for the Hillgrove Suite granitoids (Chapter 4) were largely determined by mass balance calculations. Whole-rock geochemical analyses, combined with electron microprobe data for biotite, ilmenite and alkali feldspar appropriate for each sample were entered into the program ‘GENMIX’ (Le Maitre 1981), together with calculated quartz and plagioclase end- member compositions (albite and anorthite). The calculated modal compositions were compared to point-counts for a small number of samples, and were found to be in good agreement (modes within 0A5% in most cases).

Although this method may produce erratic results for granites with a complex mineralogy (i.e. a large number of mineral phases), for granitoids with simple mineralogy such a the S- type granitoids of the Hillgrove Suite (where biotite is the sole mafic silicate), the degrees of freedom in such mass balance calculations are high, and results are consistent with observed modes. 268

(iii) Mineral analyses

Mineral analyses for chapter 5 were conducted on a JEOL JSM-840 scanning electron microscope fitted with a Tracor-Northern energy dispersive detector. Mineral data collected after 1990 (chapters 4 and 6) were conducted on the same microscope fitted with a new LINK PCXA EDS System (UK 1992) detector. The EDS software used on this machine is SPEED version 3b by Nick Ware (RSES, ANU, Canberra). Analyses were conducted with an accelerating voltage of 15kV, and a beam current of 3 nA 25-30% DT.

(iv) Whole-rock geochemical analyses

Samples collected for analysis weighed in excess of 10 kg for granitoid samples, and in excess of 5 kg for metasediments. Most enclaves were in the range 265 kg according to their size. All samples were split to 1-2 cm chips on a hydraulic rock splitter fitted with tungsten carbide jaws, with all weathered or altered portions being discarded. Large samples were then divided, with half the sample being crushed to sand particle size in a ‘Tema’ tungsten carbide ringmill. Smaller sized samples were completely crushed. A riffle splitter was then used to divide each sample to a final portion (~100 grams) which was then crushed to a fine powder by running in the ringmill for two minutes.

Major and trace element analyses were determined from low dilution fusion discs on a Philips PW-1404 X-ray fluorescence spectrometer fitted with a dual anode (Mo-Sc) tube. Operating software is the Phillips X44 software package. Standards used are a collection of European CRPG, South African Bureau SARM, and American USGS, with standard values obtained from the Geostandards Newsletter Special Issue Vol XVIII, July 1994. Precision is measured at better than 1% for most elements, with the exception of La, Ce and Cu, for which standard deviation measurements vary up to 10%.

Low dilution fusion discs were used for both major and trace elements, and were prepared according to the method of Eastell & Willis (1990). Glass discs (40 mm diameter) were prepared from powders using a 12:22 lithium metaborate/lithium tetraborate flux in a low dilution mixture of 4 gms powder : 8 gms flux. Fusion was conducted in platinum crucibles at 1000EC in a muffle furnace, and melts were agitated over a butane burner prior to final casting in platinum moulds. Loss of volatiles (LOI) was calculated as a percentage weight 269 loss on ignition, corrected for ignition gain by oxidation of ferrous iron.

Ferrous iron (FeO) was determined for each sample by titrimetry using K2Cr2O7 after dissolution in HF and H2SO4, following the method by Potts (1992). Fluorine analyses for chapters 5 and 6 were measured by fluorine ion-specific electrode after dissolution in molten NaOH, following the method of Potts (1992). Rare-earth-element (REE) analyses were conducted by Becquerel Laboratories (Lucas Heights) by INAA. La and Ce contents quoted for those samples which lack values for other REE are values measured by XRF.

(v) Isotope analyses

All isotopic analyses excepting zircon U-Pb analyses were conducted at the Centre for Isotope Studies (CIS) facility at CSIRO, North Ryde. Samples (0A1 g) were initially dissolved in triple distilled hydrofluoric acid (HF) overnight in open beakers, boiled down, and then enbombed with HF for a week period at 120EC. After debombing, samples were again boiled down with the addition of a few drops of perchloric acid (HClO4). The residue was taken up in HCl and then centrifuged. Molecular separation was then conducted in ion exchange columns, the fractions were then collected and dried. The separates were analysed on a VG354 7-collector mass spectrometer operating in fully automated mode. Standards statistics are as follows. Rb/Sr NBS987 standard 87/86 = 0.710289 ± 0.0070% (2SD). Sm/Nd CIT Nd Beta standard 143/144 = 0A511901 ± 0A0032% (2SD). 270

Appendix B

Structural data. 271

In this appendix, structural data are grouped in terms of deformation events. Foliations are recorded in the strike/dip/sense format, and lineations are recorded in the plunge-trend format.

S3 data for Tia Complex and the Rockvale Subzone of the Wollomombi Zone

135/75SW 115/82SW 118/80SW 110/82SW 112/70SW 124/73SW 120/75SW 128/84SW 116/65SW 115/72SW 054/86W 040/90 035/68NW 035/75SE 110/74SW 106/79SW 110/70W 108/80W 110/82SW 110/87NE 132/70NE 176/82NE 165/90 144/80NE 176/52W 153/35W 102/84N 040/40NW 040/56NW 078/60NW 048/53NW 110/80N 056/48N 168/80W 052/82SE 056/90 155/80SW 014/66W 070/90 016/72W 108/82N 106/82S 062/88S 085/82S 082/70N 045/70NW 052/36NW 090/76N 064/82NW

S5 data for Tia Complex, the Winterbourne Subzone of the Wollomombi Zone and areas north of the Wongwibinda Complex.

012/80W 060/78SE 065/78S 090/84S 068/82S 026/68NW 066/68N 028/62NW 072/80N 054/88NW 062/78N 012/74W 036/78SE 020/78W 178/85W 028/86E 025/80E 028/70W 044/85NW 034/78SE 052/80E 073/90 070/85N 070/85S 060/90 084/90 170/40W 040/82SE 160/72NW 020/80E 036/86SE 056/84NW 025/78W 064/74SE 079/78S 006/76E 153/88W 160/82SW 045/82SE 133/83SW 142/90 145/80SW 022/66W 082/82N 030/90 060/78NW 064/78W 068/84NW 082/77N 010/72E 024/90 166/80SW 020/80W 138/85NE 158/78W 152/88W 168/78E 155/86SW 152/82NE 155/88W 156/84W 000/88E 000/84W 014/82W 000/86E 025/86SE 032/76NW 024/82W 055/82NW 060/82NW 058/84W 060/90 052/82SE 070/80SE 052/85SE 040/78SE 070/82NW 054/86SE 060/90 048/90 024/50W 056/86SE 080/86N 035/90 040/80SE 060/90 060/90 070/82N 048/77NW 060/80SE 078/86NW 066/74W 040/86SE 068/75SE 062/90 070/90 020/75E 000/85W 008/82E 018/86E 058/85SE 030/90 040/90 040/90 018/82E 065/65NW 056/68NW 050/66W 052/54W 060/62NW 048/72W 058/58W 045/64NW 040/90 060/80NW 060/72SE 050/88NW 070/80SE 045/90 065/80NW 090/82N 085/90 070/78S 090/82N 075/90 070/84N 050/85NW 080/86N 090/85N 080/84N 018/86E 135/72NE 132/82NE 135/90 148/82NE 160/80NE 150/83SW 162/86SW 120/86SW 130/86NE 140/79SW 165/70SW 172/84SW 154/80NE 148/88NE 146/86NE 145/78SW 154/84SW 176/80W 160/88SW 178/78SW 016/74W 166/48W 144/68SW 170/82W 038/88W 030/76W 040/86W 012/74W 170/52W 158/78NE 141/76NE 141/78SW 130/86SW 140/80SW 140/90 150/78SW 165/80SW 160/60SW 135/86SW 125/78SW 122/90 120/86SW 128/82SW 172/78W 154/90 056/82N 045/66SE 040/74NW 050/80NW 018/74W 000/82W 170/82W 157/86W 006/77W 006/90 018/77W 018/80W 018/90 155/90 170/78W 156/86SW 160/90 015/86NW 145/90 010/80NW 080/90 272

5 F5 and L5 data for Tia Complex, the Winterbourne Subzone of the Wollomombi Zone and areas north of the Wongwibinda Complex. Fold plunge measurements are marked 5 with the prefix ‘F’. All other measurements are L5 stretching lineations.

F62/346 F58/270 F57/260 F55/222 F56/254 F60/254 F46/220 F52/220 F70/250 40/260 76/320 62/312 78/316 70/294 33/275 70/270 68/260 84/278 44/286 55/285 76/308 77/258 75/110 37/262 48/260 40/286 55/270 80/310 65/302 70/305 68/270 74/245 72/285 60/075 70/008 72/280 80/280 84/280 80/240 74/306 66/076 58/290 80/230 75/254 76/292 42/254 70/250

S7 (mylonite ‘C’ planes) data for all D7 sear zones throughout the Wollomombi Zone

060/70N 078/88N 055/70NW 050/84NW 066/82SE 056/74W 076/76NW 150/65W 160/60W 053/72NW 075/72W 070/70W 104/55W 098/38W 070/60N 070/25N 110/90N 010/36E 136/60W 018/24W 020/24W 010/30W 160/62W 050/78NW 080/68N 014/68E 148/68W 160/70W 178/50E 008/73E 008/82E 006/33W 150/50W 050/40NW 050/40NW 060/42NW 106/60NE 102/42NE 070/40N 050/45NW 100/66N 049/74NW 057/48NW 05050NW 025/40NW 025/68NW 147/83SW 056/82N 045/66SE 040/74NW 050/80NW 000/48W 176/52W

7 L7 (mylonite ‘C’ planes) data for all D7 sear zones throughout the Wollomombi Zone

45/275 47/265 64/310 84/310 52/212 40/238 74/342 65/260 60/270 76/308 62/320 60/302 52/347 38/004 36/336 25/333 75/190 35/115 54/223 24/280 40/260 24/252 60/236 76/324 64/314 70/102 70/284 64/278 88/107 68/116 81/071 32/270 50/240 65/075 40/300 40/344 42/252 58/356 50/323 35/338 45/325 63/017 47/332 46/326 40/314 68/315 70/276 42/254 28/172 40/235 46/235 35/264 40/275

S1 data for early Permian sedimentary sequences to the east and west of the Wollomombi Zone

052/68N 068/70N 035/65NW 055/78SE 030/43NW 054/70NW 060/86NW 040/81NW 060/80NW 056/90 066/78SE 045/85NW 048/70W 055/65NW 032/82SE 048/80NW 060/90 114/64NE 103/90 141/80SW 120/65N 090/72N 120/42N 090/60N 106/70N 100/42N 105/45N 060/60S 064/80SE 045/65SE 085/86S 060/45SE 050/70SE 036/88 060/76SE 060/70SE 060/86S 050/90 060/90 062/86S 065/85S 062/38N 100/68N 118/62N 120/70N 072/38NW 104/62N 090/60N 120/60N 125/40NE 142/58SW 108/52SW 080/70N 064/50N 054/82NW 090/74N 132/80NE 090/58N 156/62NE 128/68N 170/55W 000/50W 000/54W 005/62W 065/70N 030/70NW 112/72N 140/70SW 150/78SW 018/58W 060/58NW 025/48NW 104/76N 034/54W 076/66N 115/55N 120/55NE 090/90 110/72N 092/86N 112/76N 118/85N 090/82N 104/80N 098/88N 102/86N 070/90 060/85W 076/55NW 042/72NW 130/32NE 010/75NW 000/60W 030/86NW 030/74NW 040/70SE 030/75SE 025/85E 020/82W 020/80NW 273

Appendix C

Modal data for granitoids. 274

Modal data for Hillgrove Supersuite granitoids (Chapter 4)

Alkali Sample Pluton Quartz Plagioclase Feldspar Biotite Ilmenite Apatite Other G37 Rockvale 32.5 45.7 0.0 20.7 0.1 0.4 E29 Tobermory 30.4 35.5 13.7 19.5 0.1 0.2 N2 Rockisle 27.2 51.3 6.5 14.2 0.1 0.1 0.3 (actinolite) E27 Rockvale 31.5 35.5 12.8 19.5 0.2 0.1 N3 Rockisle 27.6 48.2 9.9 13.8 0.0 0.1 0.1 (actinolite) A243 Argyle 30.7 35.1 15.8 17.8 0.2 0.2 A274 Enmore 31.6 34.4 15.4 17.6 0.0 0.1 E28 Rockvale 31.0 33.6 17.9 16.6 0.1 0.1 A183 Gara 30.9 33.7 19.4 14.4 0.2 0.2 A246 Hillgrove 30.7 33.7 18.4 16.6 0.1 0.1 A278 Hillgrove 31.5 32.7 18.6 16.4 0.1 0.1 D15 Dundurrabin 30.6 36.4 16.2 15.9 0.1 0.3 A221 Gostwyck 31.1 34.8 18.0 15.2 0.2 0.2 A283 Hillgrove 31.2 32.8 19.7 15.5 0.2 0.1 A285 Hillgrove 31.8 33.7 17.4 16.8 0.2 0.1 A277 Gara 31.4 33.3 19.9 14.1 0.2 0.1 Y94 Kimberley Park 29.8 38.2 16.8 14.1 0.2 0.3 trace actinolite D21 Dundurrabin 31.6 35.2 16.0 16.2 0.2 0.3 A87 Argyle 30.2 34.6 19.6 14.3 0.3 0.3 A185 Gara 32.3 33.4 17.9 15.1 0.2 0.2 A60 Hillgrove 31.2 31.5 21.2 14.9 0.2 0.2 trace actinolite D11 Dundurrabin 31.2 35.9 17.0 15.4 0.2 0.2 D18 Dundurrabin 31.6 34.0 18.5 15.6 0.2 0.1 NFG SSCC 31.8 34.9 17.2 15.3 0.2 0.1 A180 Gara 31.5 34.1 18.7 15.6 0.2 0.2 A184 Gara 32.0 31.6 21.3 13.4 0.2 0.2 G31 Tobermory 31.2 34.1 20.0 14.4 0.2 0.1 Y62 Tia 31.5 36.3 16.9 14.8 0.0 0.1 A273 Enmore 31.5 32.1 22.1 13.6 0.1 0.2 E6 Kookabookra 31.6 30.4 22.4 14.5 0.1 0.1 N42 Rockisle 30.1 41.3 16.7 11.5 0.2 0.2 N43 Kilburnie 30.8 32.6 23.1 12.8 0.3 0.1 A282 Hillgrove 31.9 30.3 24.3 12.7 0.3 0.2 A276 Blue Knobby 33.1 30.6 21.5 14.3 0.2 0.1 C120 Abroi 32.0 31.9 22.1 13.5 0.1 0.1 N4 Rockisle 32.1 40.7 14.8 12.1 0.2 0.1 A284 Hillgrove 32.2 29.9 25.4 12.0 0.2 0.0 N15 Murder Dog 31.6 32.3 24.2 11.1 0.0 0.2 N44 Kilburnie 32.1 30.7 24.9 11.4 0.2 0.0 A7 Winterbourne 32.2 34.0 23.3 8.9 0.1 0.1 0.6 (Garnet) N46 Kilburnie 33.9 30.1 24.3 10.4 0.1 0.2 Trace hornblende GUM.A SSCC 33.1 38.2 18.5 9.5 0.1 0.1 N45 Kilburnie 34.9 26.4 29.5 8.2 0.1 0.1 Trace hornblende A60B Hillgrove 34.9 23.5 35.4 5.8 0.1 0.1 A286 Hillgrove 36.8 30.9 29.4 2.5 0.0 0.0

Modal data for Bundarra Suite granitoids (Chapter 4)

Alkali Sample Pluton Quartz Plagioclase feldspar Biotite Ilmenite Apatite Cordierite BI3 Gwydir River 32.9 29.9 26.0 8.7 0.1 0.2 1.2 BI2 Gwydir River 32.8 29.1 26.6 8.6 0.1 0.3 1.6 B4 Banalasta 29.9 25.1 35.1 4.8 0.2 0.3 4.0 BI7 Gwydir River 33.2 28.6 27.2 8.5 0.1 0.3 1.5 B11 Banalasta 32.9 26.0 30.3 5.5 0.2 0.3 4.2 B2 Banalasta 33.2 26.8 29.7 5.1 0.2 0.3 4.1 BI1 Gwydir River 34.2 26.6 28.7 7.7 0.1 0.3 1.9 BI9 Gwydir River 33.4 26.0 30.0 5.8 0.2 0.3 3.7 CB2 Gwydir River 35.0 30.5 24.5 5.2 0.1 0.1 3.9 B9 Banalasta 35.7 23.3 32.9 4.8 0.2 0.1 2.3 275

Modal data for Chaelundi Complex granitoids (Chapter 5) A-type suite

Sample Quartz Orthoclase Plagioclase Biotite Hornblende Opaques CC131 18.6 29 39 7.3 4 0.7 CC132 15.8 24.3 44.2 6.2 6.2 1.6 CC66 18.9 34 34.8 8.6 2.3 0.5 CC133 28.5 33.2 28.5 6.9 1 1.5 CC134 24.2 28.8 37.6 5.6 1.9 1.4 CC135 29.4 34.8 27.4 5.5 1.7 0.8 CC64 34.8 35.1 24.6 4.9 0 0.2 CC129 43.8 22.6 28.2 4.9 0 0.5 CC128 40.2 33.1 19.8 6 0 0.7 CC57 36.6 30.7 28.6 3.9 0 0.2 CC136 40.6 33 23.6 2.6 0 0

I-type suite (modes presented in Landenberger (1988))

Sample Quartz Orthoclase Plagioclase Biotite Hornblende Opaques CC7 27.2 13.2 46.5 5.5 5.6 1.4 CC12 27.5 13.8 46.4 6.5 4.7 0.7 CC8 26.3 15.9 46.8 6.5 3.1 1.2 CC106 24.6 16.6 44.7 7.7 5.2 1.1 CC114 25.5 19 39.7 9.1 5.5 0.9 CC10 26.6 22.9 39.2 7.3 3.2 0.4 CC100 23.7 18.4 47.8 6.8 2.7 0.5 CC18 23.2 15 50.2 5.3 5.7 0.3 CC119 22.4 16.6 50.6 4.9 4.4 0.9 CC9 24.3 23.1 44.3 6.7 1.1 0.3 CC20 30 23.2 37 7.6 1.3 0.4 CC103 23.4 24.8 43.6 6.7 0.9 0.3 CC101 25.3 24.5 42.6 6.2 1 0.2 276

Modal data for Woodlands Quartz Monzonite and enclaves (Chapter 6)

Host and related granitoids

Sample WQM2 WQM WQM3 ORLA

Hornblende 7.8 4.4 5.2 0 Biotite 8.7 5.8 5.1 6.6 Orthoclase 20.2 32.5 32.3 28.9 Quartz 18.3 20.7 20.9 35.2 Plag 42.8 34.6 34.5 27.9 Zircon trace trace trace trace Allanite trace trace trace trace Apatite trace trace trace trace

Enclaves

Sample WQMX2 WQMX6 WQMX7 WQMX3 WQMX10 WQMX8 WQMX5 Phenocrysts: Plagioclase 39.8 46.5 44.1 40.3 48.1 10.5 11.8 Augite 6.4 3.8 5.1 5.8 7.7 0.9 2.6 Hypersthene 7.7 7.1 1.2 3.1 1.2 1.5 - Fe-Ti oxides 2.0 1.0 1.1 1.1 1.3 1.4 1.5 Hornblende - - 0.2 - 2.2 0.9 0.5 Biotite - - 0.2 - 0.7 0.2 0.6 Xenocrysts: Quartz - - 0.5 - - 0.5 0.6 Orthoclase -----3.74.4 Matrix: Plagioclase 20.7 19.3 17.6 23.3 19.4 39.5 33.5 Hornblende 15.2 13.1 15.5 16.2 7.5 29.7 18.4 Biotite 6.7 7.0 11.1 8.0 2.4 4.2 14.9 Magnetite 1.0 0.4 0.6 0.5 2.0 2.2 2.2 Quartz 0.5 1.8 2.6 1.7 6.9 3.2 7.9 Orthoclase - - 0.2 - 0.5 1.6 1.1 Apatite trace trace trace trace 0.1 trace trace 277

Appendix D

Mineral analyses (Electron-microprobe). 278

Hillgrove Supersuite and Bundarra Suite Mineral Analyses - Chapter 4.

Amphiboles (Rockisle Suite)

Actinolite (secondary)

Sample A60 A60 A60 N4 N4 N4 N4 N46 Analysis 12312341 Pluton ------Hillgrove ------Rockisle ------Kilburnie

SiO2 51.90 53.72 52.89 50.56 52.45 52.27 52.42 49.40 TiO2 0.00 0.00 0.10 0.48 0.00 0.22 0.25 0.30 Al2O3 0.82 0.63 1.70 3.55 1.63 2.22 2.20 3.49 FeO 19.68 17.06 15.91 16.71 16.12 16.44 16.19 21.42 MnO 0.84 0.64 0.61 0.94 0.73 0.84 0.94 0.51 MgO 11.25 12.95 12.98 12.10 12.90 13.04 13.12 10.75 CaO 11.40 12.26 12.56 11.67 12.18 11.97 11.98 10.35

Na2O 0.22 0.19 0.21 0.48 0.00 0.33 0.18 1.07 K2O 0.10 0.00 0.12 0.34 0.12 0.19 0.18 0.43 Total 96.21 97.63 97.20 96.82 96.14 97.62 97.58 97.98 Cations Si 7.870 7.921 7.806 7.546 7.821 7.709 7.723 7.466 AlIV 0.130 0.079 0.194 0.454 0.179 0.291 0.277 0.534 AlVI 0.016 0.030 0.101 0.170 0.108 0.094 0.104 0.087 Fe3+ 0.204 0.076 0.011 0.151 0.098 0.148 0.157 0.395 Fe2+ 2.128 1.967 1.945 1.814 1.834 1.762 1.712 1.996 Mn 0.108 0.080 0.076 0.119 0.092 0.105 0.117 0.065 Mg 2.543 2.847 2.856 2.692 2.868 2.867 2.882 2.422 Ti 0.000 0.000 0.011 0.054 0.000 0.024 0.028 0.034 Ca 1.852 1.937 1.986 1.866 1.946 1.891 1.891 1.676 Na 0.065 0.054 0.060 0.139 0.000 0.094 0.051 0.313 K 0.019 0.000 0.023 0.065 0.023 0.036 0.034 0.083 Total 14.936 14.991 15.068 15.069 14.968 15.021 14.976 15.072 Mg# 51.034 57.286 58.437 56.374 58.626 58.734 59.197 49.649 Na (A site) 0.000 0.000 0.046 0.005 0.000 0.000 0.000 0.000 K (A site) 0.000 0.000 0.023 0.065 0.000 0.021 0.000 0.072 Hornblende (primary)

N46 N46 N46 N46 N46 N46 N45 2345671 ------Kilburnie ------

SiO2 43.99 45.71 47.71 45.05 45.12 45.90 44.65 TiO2 1.68 0.59 0.37 1.25 0.73 0.58 1.8 Al2O3 6.76 8.78 4.18 6.07 5.60 5.13 6.15 FeO 26.34 20.31 26.22 25.99 27.38 27.18 25.04 MnO 0.33 0.38 0.56 0.54 0.56 0.65 0.39 MgO 5.88 5.70 8.17 5.98 5.86 6.36 6.28 CaO 10.19 8.06 8.82 10.14 9.70 9.95 10.5

Na2O 2.19 2.73 1.21 1.26 1.03 1.21 1.69 K2O 0.86 0.81 0.56 0.78 0.62 0.58 0.81 Total 98.21 93.07 97.80 97.08 96.75 97.54 97.44 Cations Si 6.877 7.229 7.368 7.078 7.151 7.196 6.990 AlIV 1.123 0.771 0.632 0.922 0.849 0.804 1.010 AlVI 0.123 0.866 0.129 0.202 0.197 0.144 0.125 Fe3+ 0.223 0.021 0.641 0.296 0.465 0.436 0.166 Fe2+ 3.042 2.648 2.232 2.883 2.792 2.779 2.980 Mn 0.044 0.051 0.073 0.072 0.075 0.086 0.052 Mg 1.370 1.344 1.881 1.401 1.385 1.487 1.466 Ti 0.198 0.070 0.043 0.148 0.087 0.068 0.212 Ca 1.707 1.366 1.459 1.707 1.647 1.671 1.761 Na 0.664 0.837 0.362 0.384 0.316 0.368 0.513 K 0.172 0.163 0.110 0.156 0.125 0.116 0.162 Total 15.542 15.366 14.932 15.246 15.088 15.155 15.436 Mg# 29.288 33.070 38.966 30.118 29.356 31.051 31.433 Na (A site) 0.371 0.203 0.000 0.090 0.000 0.039 0.274 K (A site) 0.172 0.163 0.000 0.156 0.089 0.116 0.162 279

Hillgrove Supersuite and Bundarra Suite Mineral Analyses - Chapter 4.

Biotite Analyses (averaged for each sample)

Suite ------Rockisle ------Hillgrove ------SampleA60 N4 N46 N45 Y62 A60B A180 D15 D18 A7 NFG N15 E29

SiO236.23 36.16 35.75 35.22 35.79 35.65 35.89 35.59 35.14 35.51 35.30 36.11 35.15 TiO2 3.22 3.99 4.37 5.20 3.17 4.23 5.36 4.24 3.75 3.75 4.16 4.46 3.17 Al2O14.753 14.40 13.34 12.60 16.17 15.21 13.37 13.68 14.39 15.19 14.17 13.41 16.86 FeO22.78 21.13 28.25 28.82 21.50 24.53 24.92 24.88 23.76 22.75 25.00 24.91 22.20 MnO0.33 0.40 0.24 0.29 0.53 0.46 0.36 0.66 0.54 0.34 0.67 0.39 0.43 MgO8.41 9.83 5.43 5.16 8.21 5.99 7.41 7.30 7.52 7.99 7.18 7.25 7.92 CaO 0.00 0.03 0.00 0.00 0.00 0.06 0.00 0.00 0.04 0.17 0.10 0.04 0.00

Na2O0.18 0.00 0.00 0.00 0.00 0.08 0.23 0.06 0.24 0.00 0.34 0.00 0.09 K2O 9.76 9.60 9.39 9.05 9.81 9.78 9.42 9.63 9.37 9.51 9.40 9.72 9.74 Total95.65 95.55 96.76 96.33 95.17 95.97 96.94 96.02 94.74 95.22 96.32 96.28 95.56 Cations Si 5.627 5.579 5.635 5.601 5.551 5.572 5.565 5.586 5.556 5.542 5.527 5.640 5.449 AlIV 2.373 2.421 2.365 2.361 2.449 2.428 2.435 2.414 2.444 2.458 2.473 2.360 2.551 AlVI 0.327 0.197 0.113 0.000 0.507 0.373 0.008 0.116 0.237 0.334 0.141 0.109 0.530 Ti 0.376 0.463 0.518 0.622 0.370 0.497 0.625 0.500 0.446 0.440 0.490 0.524 0.369 Fe2+ 2.959 2.726 3.724 3.832 2.788 3.206 3.230 3.265 3.142 2.968 3.273 3.253 2.878 Mn 0.043 0.052 0.032 0.039 0.069 0.060 0.047 0.087 0.072 0.045 0.089 0.052 0.057 Mg 1.948 2.260 1.277 1.223 1.897 1.397 1.712 1.709 1.772 1.860 1.676 1.688 1.831 Na 0.053 0.000 0.000 0.000 0.000 0.025 0.069 0.018 0.074 0.000 0.103 0.000 0.028 K 1.934 1.890 1.888 1.836 1.941 1.949 1.863 1.927 1.889 1.892 1.877 1.937 1.926 Total15.641 15.589 15.552 15.514 15.572 15.507 15.555 15.622 15.632 15.539 15.649 15.563 15.618 Mg#39.70 45.33 25.53 24.20 40.50 30.35 34.64 34.36 36.06 38.52 33.87 34.16 38.89

SampleA285 E28 A276 A278 E27 A87 Bana 1 B11 BI2 B4 BI9 CB2 BI7 Suite ------Hillgrove ------Bundarra ------

SiO235.19 35.11 35.09 35.57 35.23 36.59 34.62 35.10 34.99 34.33 34.42 34.73 34.73 TiO2 4.60 3.46 3.83 3.64 3.39 3.89 3.38 3.27 4.15 3.65 4.04 3.68 4.24 Al2O15.153 16.97 16.86 15.91 16.78 13.92 18.86 18.15 15.73 17.79 17.57 18.52 16.15 FeO23.28 22.04 23.06 23.46 21.88 24.09 23.79 23.81 24.16 24.63 24.58 24.28 24.42 MnO0.39 0.35 0.43 0.33 0.41 0.30 0.31 0.41 0.39 0.31 0.37 0.76 0.47 MgO7.38 7.66 7.42 7.64 8.15 7.41 5.64 5.61 7.15 5.37 5.29 4.15 6.63 CaO 0.04 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.05 0.05 0.00

Na2O0.16 0.00 0.00 0.07 0.00 0.00 0.19 0.05 0.00 0.19 0.13 0.00 0.08 K2O 9.75 9.53 9.69 9.55 9.73 9.69 9.60 9.65 9.36 9.65 9.32 9.46 9.52 Total95.99 95.12 96.38 96.17 95.58 95.87 96.37 96.04 95.91 95.90 95.74 95.60 96.23

Structural Formulae

Si 5.479 5.453 5.412 5.504 5.449 5.699 5.345 5.438 5.453 5.365 5.377 5.421 5.410 AlIV 2.521 2.547 2.588 2.496 2.551 2.301 2.655 2.562 2.547 2.635 2.623 2.579 2.590 AlVI 0.260 0.560 0.477 0.404 0.508 0.253 0.775 0.751 0.342 0.641 0.611 0.828 0.374 Ti 0.539 0.404 0.445 0.424 0.395 0.456 0.392 0.381 0.486 0.428 0.474 0.432 0.497 Fe2+ 3.031 2.862 2.974 3.035 2.830 3.137 3.070 3.084 3.148 3.218 3.211 3.169 3.181 Mn 0.051 0.046 0.056 0.043 0.054 0.040 0.040 0.054 0.051 0.040 0.049 0.100 0.062 Mg 1.714 1.773 1.705 1.763 1.880 1.720 1.297 1.295 1.661 1.250 1.232 0.965 1.539 Na 0.047 0.000 0.000 0.020 0.000 0.000 0.056 0.014 0.000 0.058 0.039 0.000 0.024 K 1.936 1.889 1.907 1.885 1.919 1.925 1.890 1.907 1.860 1.923 1.856 1.884 1.891 Total15.577 15.534 15.564 15.575 15.586 15.530 15.521 15.485 15.547 15.559 15.473 15.378 15.568 Mg#36.12 38.25 36.44 36.74 39.91 35.41 29.71 29.56 34.54 27.98 27.73 23.34 32.60 280

Hillgrove Supersuite and Bundarra Suite Mineral Analyses - Chapter 4 (continued).

Cordierite analyses - Bundarra Suite

Sample ------B1 ------CB1 ------Analysis 12345671234567 Core Core Core Mid Mid Rim Rim Core Core Mid Mid Rim Rim Rim

SiO2 47.10 47.46 47.22 47.14 47.46 47.53 46.90 47.89 46.68 47.66 47.73 46.45 46.70 46.38 TiO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Al2O3 32.17 32.64 32.22 32.47 32.27 32.06 31.70 32.43 32.00 32.24 32.08 31.53 31.68 31.75 FeO 11.40 11.28 11.46 11.28 11.34 10.94 12.22 10.58 9.79 8.84 8.74 11.65 11.72 11.26 MnO 0.82 0.70 0.53 0.83 0.77 0.75 0.78 0.13 0.35 0.17 0.16 1.26 0.95 1.22 MgO 5.00 5.01 5.24 5.22 5.13 5.47 4.82 6.70 6.56 6.62 7.13 4.57 4.53 3.92 CaO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.06

Na2O 0.87 1.06 0.87 1.15 0.78 0.78 0.78 0.61 1.27 1.00 0.90 1.22 0.88 1.69 K2O 0.00 0.00 0.07 0.06 0.05 0.00 0.00 0.00 0.15 0.10 0.00 0.00 0.05 0.09 Total 97.36 98.16 97.61 98.16 97.80 97.54 97.19 98.34 96.79 96.64 96.74 96.67 96.53 96.37

Si 4.999 4.992 4.996 4.969 5.009 5.020 5.005 4.993 4.957 5.024 5.021 4.993 5.014 5.001 Ti 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Al 4.024 4.046 4.017 4.034 4.014 3.990 3.987 3.985 4.004 4.005 3.977 3.995 4.009 4.034 Fe2+ 1.012 0.992 1.014 0.994 1.001 0.966 1.091 0.922 0.869 0.779 0.769 1.047 1.052 1.015 Mn 0.074 0.062 0.047 0.074 0.069 0.067 0.071 0.011 0.031 0.015 0.014 0.115 0.086 0.111 Mg 0.791 0.786 0.827 0.820 0.807 0.861 0.767 1.041 1.038 1.040 1.118 0.732 0.725 0.630 Ca 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.007 Na 0.179 0.216 0.178 0.235 0.160 0.160 0.161 0.123 0.261 0.204 0.184 0.254 0.183 0.353 K 0.000 0.000 0.009 0.008 0.007 0.000 0.000 0.000 0.020 0.013 0.000 0.000 0.007 0.012 Total 11.079 11.094 11.089 11.135 11.067 11.065 11.082 11.076 11.182 11.082 11.082 11.136 11.077 11.165 Mg# 42.16 42.69 43.78 43.43 43.01 45.46 39.78 52.72 53.55 56.70 58.81 38.66 38.91 35.87

Garnet analyses

Sample ------A7 (Hillgrove) ------B11 (Bundarra) ------Analysis 10334 10335 10336 10337 10477 10478 10479 10480 Core Core Rim Core Core Core Rim Rim

SiO2 36.97 36.81 36.65 36.60 36.96 37.07 36.76 37.00 TiO2 0.11 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Al2O3 21.12 20.91 21.02 20.86 21.19 21.04 20.85 20.91 FeO 33.95 33.90 33.71 34.12 36.53 36.11 36.20 36.14 MnO 3.87 4.11 3.57 3.82 3.24 3.22 3.04 3.33 MgO 3.79 3.65 3.66 3.87 2.54 2.58 2.81 2.53 CaO 0.97 0.89 0.97 0.98 0.70 0.69 0.92 0.86

Na2O 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 K2O 0.00 0.00 0.00 0.00 0.07 0.00 0.00 0.06 Total 101.01 100.28 99.58 100.26 101.24 100.71 100.59 100.83

Si 5.923 5.937 5.937 5.909 5.943 5.977 5.945 5.970 AlIV 0.077 0.063 0.063 0.091 0.057 0.023 0.055 0.030 AlVI 3.911 3.911 3.949 3.878 3.959 3.976 3.919 3.947 Ti 0.013 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe2+ 4.548 4.572 4.566 4.606 4.912 4.869 4.896 4.876 Mn 0.525 0.561 0.490 0.522 0.441 0.440 0.416 0.455 Mg 0.905 0.878 0.884 0.931 0.609 0.620 0.678 0.609 Ca 0.166 0.154 0.168 0.169 0.121 0.119 0.159 0.149 Total 16.070 16.076 16.057 16.107 16.042 16.024 16.068 16.035 End members Almandine 73.77 74.16 74.76 73.94 80.75 80.50 79.62 80.09 Grossular 2.70 2.49 2.76 2.72 1.98 1.97 2.59 2.44 Pyrope 14.68 14.24 14.47 14.95 10.01 10.25 11.02 10.00 Spessartine 8.52 9.11 8.02 8.38 7.25 7.27 6.77 7.47 Andradite 0.33 0.00 0.00 0.00 0.00 0.00 0.00 0.00 281

Hillgrove Supersuite and Bundarra Suite Mineral Analyses - Chapter 4 (continued).

Average Ilmenite Analyses

Sample A60 N4 N46 N45 Y62 A60B A180 D15 D18 A7 NFG E6 N15 E29 Suite ------Rockisle ------Hillgrove ------

SiO2 0.14 0.00 0.03 0.00 0.00 0.05 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 TiO2 53.95 53.73 54.19 52.94 53.63 53.74 53.33 54.05 53.93 54.25 54.47 53.17 53.18 54.09 Al2O3 0.12 0.00 0.04 0.00 0.00 0.21 0.11 0.00 0.00 0.00 0.00 0.00 0.00 0.00 FeO 43.00 39.49 44.38 45.22 38.08 41.10 42.08 36.38 39.52 40.90 39.45 40.46 41.54 40.89 MnO 4.73 7.95 3.23 2.81 9.18 6.38 5.39 11.11 7.90 5.53 7.26 5.42 5.76 6.83 MgO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.15 0.00 0.00 Total 102.17 101.41 101.95 101.04 100.88 101.59 100.97 101.79 101.40 100.82 101.19 99.32 100.60 101.86

Si 0.007 0.000 0.002 0.000 0.000 0.002 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Ti 2.002 2.011 2.013 1.994 2.012 2.003 2.003 2.014 2.014 2.032 2.030 2.023 2.006 2.012 Al 0.007 0.000 0.003 0.000 0.000 0.012 0.006 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe2+ 1.774 1.643 1.832 1.893 1.588 1.703 1.757 1.507 1.640 1.703 1.635 1.711 1.742 1.691 Mn 0.197 0.335 0.135 0.119 0.388 0.268 0.228 0.466 0.332 0.233 0.305 0.232 0.245 0.286 Mg 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.011 0.000 0.000 Total 3.988 3.989 3.984 4.006 3.988 3.988 3.994 3.986 3.986 3.968 3.970 3.977 3.994 3.988 Mn # 10.02 16.93 6.87 5.92 19.63 13.59 11.48 23.63 16.84 12.04 15.71 11.95 12.31 14.46 (Mn# = Mn/[Mn+Fe])

Sample A285 E28 A276 A278 E27 A87 G37 A286 B11 BI2 B4 BI9 CB2 BI7 Suite ------Hillgrove ------Bundarra ------

SiO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 TiO2 54.26 54.10 54.19 54.70 54.06 54.02 52.86 54.36 54.27 53.79 54.28 53.98 54.18 54.18 Al2O3 0.00 0.00 0.00 0.00 0.18 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 FeO 42.18 41.60 41.59 42.25 40.78 41.69 39.76 41.10 42.66 41.96 43.29 43.10 39.68 42.12 MnO 5.41 5.92 5.96 5.46 6.60 5.73 6.81 4.72 4.61 5.89 3.65 4.55 7.59 5.65 MgO 0.06 0.11 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Total 102.04 101.76 101.74 102.58 101.66 101.47 99.80 100.18 101.54 101.63 101.22 101.86 101.45 102.01

Si 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Ti 2.014 2.012 2.015 2.019 2.011 2.015 2.013 2.042 2.020 2.006 2.025 2.012 2.019 2.012 Al 0.000 0.000 0.000 0.000 0.010 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe2+ 1.741 1.720 1.720 1.734 1.686 1.729 1.683 1.716 1.766 1.740 1.796 1.786 1.644 1.739 Mn 0.226 0.248 0.250 0.227 0.276 0.241 0.292 0.200 0.193 0.247 0.153 0.191 0.318 0.236 Mg 0.005 0.008 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Total 3.986 3.988 3.985 3.981 3.984 3.985 3.987 3.958 3.980 3.994 3.975 3.988 3.981 3.988 Mn # 11.50 12.60 12.67 11.57 14.08 12.21 14.78 10.42 9.87 12.45 7.87 9.66 16.23 11.96 282

Hillgrove Supersuite and Bundarra Suite Mineral Analyses - Chapter 4 (continued).

Average Microcline analyses

Sample A60 N4 N46 N45 Y62 A60B A180 D15 D18 A7 NFG E6 N15 E29 Suite ------Rockisle ------Hillgrove ------

SiO2 65.11 63.78 64.60 65.09 63.81 65.63 64.61 63.76 64.22 63.97 60.97 63.46 64.11 63.11 Al2O3 18.25 18.75 19.04 19.21 18.54 18.95 18.37 18.39 18.06 18.14 18.45 17.57 18.42 19.02 CaO 0.29 0.05 0.28 0.27 0.17 0.14 0.23 0.22 0.11 0.25 0.34 0.31 0.00 0.07

Na2O 3.36 0.70 3.75 4.98 1.02 4.53 3.14 1.10 0.89 3.27 1.56 2.50 1.09 1.18 K2O 11.85 15.88 11.26 9.72 15.66 10.37 12.06 15.36 15.42 12.17 13.94 12.77 15.27 15.01 Total 99.07 99.20 98.92 99.37 99.32 99.75 98.61 99.02 98.70 97.89 95.49 96.85 98.88 98.54 Cations Si 3.001 2.974 2.971 2.968 2.976 2.984 2.993 2.981 3.001 2.989 2.950 3.005 2.989 2.959 Al 0.991 1.030 1.032 1.032 1.019 1.016 1.003 1.013 0.995 0.999 1.052 0.980 1.012 1.051 Ca 0.014 0.002 0.014 0.013 0.008 0.007 0.011 0.011 0.006 0.013 0.018 0.015 0.000 0.004 Na 0.300 0.063 0.334 0.440 0.092 0.400 0.282 0.100 0.081 0.297 0.146 0.229 0.099 0.107 K 0.696 0.945 0.660 0.565 0.931 0.602 0.713 0.916 0.919 0.725 0.860 0.771 0.908 0.897 Total 5.002 5.015 5.011 5.019 5.027 5.008 5.003 5.020 5.001 5.022 5.027 5.005 5.008 5.018 End-members An 1.393 0.223 1.368 1.295 0.823 0.693 1.135 1.074 0.548 1.210 1.721 1.524 0.000 0.349 Ab 29.674 6.265 33.156 43.212 8.936 39.642 28.030 9.713 8.021 28.666 14.286 22.554 9.787 10.639 Or 68.933 93.513 65.476 55.493 90.241 59.665 70.835 89.213 91.432 70.124 83.994 75.923 90.213 89.012

Sample A285 E28 A276 A278 E27 A87 A286 Bana1 BI2 B11 B4 BI9 CB2 BI7 Suite ------Hillgrove ------Bundarra ------

SiO2 68.86 64.26 65.98 64.54 63.46 64.20 65.07 64.32 63.11 63.24 63.93 63.07 61.34 61.88 Al2O3 16.31 19.01 17.47 18.43 19.11 18.70 19.22 19.09 18.77 19.03 18.63 18.92 20.50 19.36 CaO 0.44 0.32 0.21 0.20 0.22 0.21 0.21 0.415 0.32 0.17 0.27 0.10 0.33 0.28

Na2O 2.57 2.08 2.63 3.19 3.52 1.93 3.85 2.97 1.75 2.12 2.58 0.71 2.72 2.22 K2O 10.70 13.65 10.36 11.93 11.53 13.96 11.20 11.995 13.89 13.36 12.94 15.68 12.43 12.76 Total 99.06 99.40 96.79 98.28 98.27 99.14 99.54 99.02 97.84 98.16 98.35 98.48 97.80 96.60 Cations Si 3.124 2.968 3.065 2.992 2.957 2.978 2.971 2.967 2.965 2.961 2.976 2.961 2.884 2.937 Al 0.872 1.035 0.956 1.007 1.049 1.022 1.034 1.038 1.039 1.050 1.022 1.047 1.136 1.083 Ca 0.021 0.016 0.010 0.010 0.011 0.010 0.010 0.021 0.016 0.009 0.013 0.005 0.017 0.014 Na 0.226 0.186 0.236 0.287 0.318 0.173 0.340 0.266 0.159 0.192 0.233 0.065 0.248 0.204 K 0.619 0.804 0.614 0.705 0.685 0.826 0.652 0.706 0.832 0.798 0.768 0.939 0.746 0.772 Total 4.862 5.009 4.882 5.001 5.020 5.011 5.008 4.999 5.012 5.009 5.013 5.017 5.044 5.010 End-members An 2.442 1.550 1.214 0.992 1.059 1.018 1.024 2.068 1.573 0.854 1.327 0.499 1.646 1.437 Ab 26.052 18.477 27.465 28.618 31.367 17.166 33.936 26.777 15.818 19.271 22.947 6.407 24.547 20.612 Or 71.506 79.973 71.321 70.390 67.574 81.817 65.040 71.155 82.608 79.875 75.726 93.095 73.807 77.951 283

Hillgrove Supersuite and Bundarra Suite Mineral Analyses - Chapter 4 (continued). Plagioclase analyses (most calcic from each sample)

Sample A60 N4 N46 N45 Y62 A60B A180 D15 D18 A7 NFG N15 E29 Suite ------Rockisle ------Hillgrove ------

SiO2 59.10 56.98 55.27 63.66 58.32 59.21 57.03 59.87 58.07 58.10 59.33 61.00 58.35 Al2O3 25.37 26.55 27.92 22.54 25.92 25.48 26.63 25.02 25.99 26.13 24.87 24.54 26.13 FeO 0.00 0.29 0.00 0.00 0.00 0.00 0.00 0.13 0.00 0.00 0.00 0.00 0.00 CaO 7.63 8.25 10.48 4.00 7.85 7.15 8.92 6.73 8.11 8.15 7.42 6.33 7.78

Na2O 7.26 6.27 5.74 9.19 7.05 7.42 6.48 7.53 6.57 7.11 7.34 8.16 7.22 K2O 0.17 0.75 0.24 0.68 0.18 0.24 0.41 0.21 0.24 0.27 0.12 0.20 0.20 Total 99.53 99.21 99.74 100.17 99.32 99.87 99.56 99.48 98.97 99.87 99.18 100.44 99.79

Si 2.651 2.582 2.500 2.817 2.624 2.654 2.574 2.680 2.620 2.608 2.670 2.708 2.617 Al 1.341 1.418 1.489 1.175 1.374 1.346 1.417 1.320 1.382 1.383 1.319 1.284 1.381 Fe 0.000 0.011 0.000 0.000 0.000 0.000 0.000 0.005 0.000 0.000 0.000 0.000 0.000 Ca 0.367 0.401 0.508 0.190 0.378 0.343 0.431 0.323 0.392 0.392 0.358 0.301 0.374 Na 0.631 0.551 0.503 0.788 0.615 0.645 0.567 0.653 0.575 0.619 0.640 0.702 0.628 K 0.010 0.043 0.014 0.038 0.010 0.014 0.024 0.012 0.014 0.015 0.007 0.011 0.011 Total 4.999 5.006 5.014 5.009 5.002 5.002 5.013 4.993 4.983 5.017 4.994 5.007 5.012

An 36.38 40.26 49.54 18.66 37.70 34.27 42.21 32.66 39.98 38.20 35.60 29.67 36.90 Ab 62.65 55.38 49.11 77.57 61.27 64.36 55.48 66.13 58.61 60.30 63.72 69.21 61.97 Or 0.97 4.36 1.35 3.78 1.03 1.37 2.31 1.21 1.41 1.51 0.69 1.12 1.13

Sample A285 E28 A276 A278 E27 G37 A286 BANA1 B11 BI2 B4 CB2 BI7

SiO2 58.80 56.65 59.52 59.38 58.59 59.38 62.73 59.44 59.62 60.04 61.75 61.29 59.73 Al2O3 25.69 27.37 25.72 25.49 25.66 25.60 22.77 25.22 24.88 24.72 23.53 24.24 25.28 FeO 0.00 0.00 0.00 0.10 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 CaO 7.94 9.61 7.01 7.33 7.67 7.59 3.07 6.69 6.78 6.41 5.02 5.46 7.21

Na2O 7.32 6.17 7.62 7.10 7.15 7.48 8.98 7.64 7.61 7.42 8.40 8.11 7.55 K2O 0.11 0.13 0.23 0.54 0.16 0.18 0.55 0.27 0.59 0.87 0.36 0.25 0.45 Total 99.96 99.93 100.10 100.04 99.43 100.43 98.11 99.39 99.85 99.45 99.06 99.34 100.33

Si 2.633 2.545 2.652 2.654 2.636 2.646 2.818 2.668 2.676 2.692 2.762 2.735 2.662 Al 1.355 1.449 1.351 1.343 1.361 1.345 1.206 1.334 1.316 1.306 1.240 1.275 1.328 Fe 0.000 0.000 0.000 0.004 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Ca 0.381 0.463 0.335 0.351 0.370 0.362 0.148 0.322 0.326 0.308 0.241 0.261 0.344 Na 0.635 0.537 0.658 0.615 0.624 0.646 0.782 0.665 0.662 0.645 0.728 0.702 0.652 K 0.006 0.007 0.013 0.031 0.009 0.010 0.032 0.015 0.034 0.050 0.021 0.014 0.026 Total 5.011 5.002 5.008 4.998 5.000 5.010 4.986 5.005 5.014 5.002 4.992 4.986 5.013 An 37.25 45.91 33.27 35.20 36.88 35.57 15.37 32.11 31.90 30.71 24.31 26.72 33.68 Ab 62.14 53.35 65.44 61.71 62.21 63.43 81.35 66.35 64.79 64.33 73.61 71.82 63.82 Or 0.61 0.74 1.30 3.09 0.92 1.00 3.28 1.54 3.31 4.96 2.08 1.46 2.50 284

Chaelundi Complex A-type Suite Mineral Analyses - Chapter 5.

Pyroxenes and Cummingtonite CPX OPX Cummingtonite Sample CC131 CC66 CC132 CC132 Sample CC66 CC131 CC132

SiO2 51.83 50.33 52.52 49.94 SiO2 50.20 50.89 49.74 TiO2 0.12 0.60 0.34 0.00 TiO2 0.00 0.19 0.18 Al2O3 0.19 1.70 0.87 0.00 Al2O3 0.83 0.43 0.81 FeO 11.65 11.33 19.60 34.43 FeO 32.99 30.67 33.83 MnO 0.38 0.44 0.48 1.04 MnO 1.76 1.52 1.49 MgO 12.73 13.37 24.13 13.69 MgO 10.16 11.32 9.23 CaO 21.52 20.52 1.36 1.51 CaO 1.55 1.39 1.51

Na2O 0.00 0.61 0.00 0.00 Na2O 0.64 0.00 0.21 Total 98.42 98.90 99.30 100.61 Total 98.13 96.41 97.00 Cations Cations Si 1.990 1.924 1.953 1.977 Si 7.782 7.902 7.819 AlIV 0.009 0.076 0.038 0.000 AlIV 0.152 0.079 0.150 AlVI 0.000 0.001 0.000 0.000 AlVI 0.000 0.000 0.000 Fe3+ 0.003 0.054 0.023 0.030 Fe3+ 0.299 0.046 0.147 Fe2+ 0.369 0.266 0.569 1.087 Fe2+ 3.739 3.899 4.183 Mn 0.012 0.014 0.015 0.035 Mn 0.231 0.200 0.198 Mg 0.729 0.762 1.338 0.808 Mg 2.348 2.620 2.163 Ti 0.003 0.017 0.010 0.000 Ti 0.000 0.022 0.021 Ca 0.885 0.841 0.054 0.064 Ca 0.257 0.231 0.254 Na 0.000 0.045 0.000 0.000 Na 0.192 0.000 0.064 K 0.000 0.000 0.000 0.000 K 0.000 0.000 0.000 Total 4.000 4.000 4.000 4.000 Total 15.000 15.000 15.000 En 36.700 39.650 67.447 40.639 Mg# 35.486 38.731 32.328 Fs 18.717 16.620 29.821 56.140 Wo 44.583 43.730 2.732 3.221 Mg# 65.489 69.547 68.802 41.244

Biotite

Sample CC128 CC129 CC131 CC132 CC133 CC134 CC135 CC136 CC57 CC64 CC66

SiO2 36.37 35.87 35.29 35.46 35.32 35.57 35.92 37.55 35.48 34.08 34.50 TiO2 2.99 2.64 4.57 3.44 3.39 3.20 4.00 1.98 2.81 3.94 4.14 Al2O3 13.53 14.58 12.30 12.23 12.54 12.73 12.74 20.01 13.26 12.71 12.59 FeO 26.96 26.33 25.99 28.02 27.02 27.96 27.86 22.31 27.32 29.61 27.62 MnO 1.21 1.28 0.27 0.26 0.53 0.44 0.38 1.75 1.38 0.60 0.29 MgO 5.47 4.45 7.25 7.03 7.13 6.60 6.53 1.53 4.91 4.65 6.66 CaO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.18

Na2O 0.26 0.26 0.00 0.00 0.27 0.32 0.20 0.54 0.36 0.20 0.45 K2O 9.45 9.49 9.24 9.87 9.32 9.43 9.38 9.82 9.47 9.33 8.96 Total 96.24 94.90 94.91 96.31 95.52 96.25 97.01 95.49 94.99 95.12 95.39 Cations Si 5.750 5.738 5.632 5.647 5.640 5.655 5.648 5.791 5.725 5.558 5.536 AlIV 2.250 2.262 2.313 2.295 2.360 2.345 2.352 2.209 2.275 2.442 2.381 AlVI 0.271 0.487 0.000 0.000 0.000 0.040 0.009 1.428 0.246 0.001 0.000 Ti 0.356 0.318 0.549 0.412 0.407 0.383 0.473 0.230 0.341 0.483 0.500 Fe2+ 3.564 3.522 3.468 3.731 3.608 3.717 3.663 2.877 3.686 4.038 3.706 Mn 0.162 0.173 0.036 0.035 0.072 0.059 0.051 0.229 0.189 0.083 0.039 Mg 1.289 1.061 1.725 1.669 1.697 1.564 1.531 0.352 1.181 1.131 1.593 Na 0.080 0.081 0.000 0.000 0.084 0.099 0.061 0.161 0.113 0.063 0.140 K 1.906 1.936 1.881 2.005 1.898 1.912 1.881 1.932 1.949 1.941 1.834 Total 15.627 15.578 15.604 15.795 15.764 15.775 15.669 15.208 15.704 15.739 15.730 Mg# 26.564 23.155 33.215 30.906 31.994 29.620 29.472 10.895 24.267 21.874 30.065 Mn# 4.35 4.69 1.04 0.93 1.95 1.57 1.36 7.36 4.87 2.01 1.05 285

Chaelundi Complex A-type Suite Mineral Analyses - Chapter 5 (continued).

Hornblende Sample CC131 CC132 CC133 CC134 CC135 CC66

SiO2 45.68 44.64 45.40 44.83 47.49 44.83 TiO2 1.21 1.15 1.23 1.57 0.42 1.12 Al2O3 5.19 5.79 5.97 6.09 4.55 5.52 FeO 22.85 25.38 25.62 24.99 23.37 25.06 MnO 0.51 0.46 0.25 0.85 1.12 0.76 MgO 8.68 7.17 6.60 6.39 8.22 7.51 CaO 10.20 9.90 10.00 9.89 10.53 10.13

Na2O 1.47 1.38 1.60 1.86 1.23 1.64 K2O 0.58 0.68 0.68 0.63 0.51 0.51 Total 96.37 96.55 97.35 97.10 97.44 97.08 Cations Si 7.120 7.036 7.085 7.027 7.321 7.030 AlIV 0.880 0.964 0.915 0.973 0.679 0.970 AlVI 0.074 0.111 0.183 0.152 0.148 0.050 Fe3+ 0.348 0.424 0.300 0.274 0.306 0.407 Fe2+ 2.352 2.582 2.803 2.782 2.463 2.554 Mn 0.067 0.061 0.033 0.113 0.146 0.101 Mg 2.017 1.685 1.536 1.493 1.889 1.756 Ti 0.142 0.136 0.144 0.185 0.049 0.132 Ca 1.703 1.672 1.672 1.661 1.739 1.702 Na 0.444 0.422 0.484 0.565 0.368 0.499 K 0.115 0.137 0.135 0.126 0.100 0.102 Total 15.263 15.230 15.291 15.352 15.207 15.302 Mg# 42.159 35.453 32.867 32.027 39.328 36.446 Mn# 2.433 2.002 1.053 3.560 5.017 3.297 Na (A site) 0.148 0.093 0.156 0.226 0.107 0.200 K (A site) 0.115 0.137 0.135 0.126 0.100 0.102 Plagioclase Cores Sample CC128 CC129 CC131 CC132 CC133 CC134 CC135 CC136 CC57 CC64 CC66

SiO2 65.81 65.87 61.02 56.62 56.51 58.35 60.39 67.06 67.25 64.13 60.63 Al2O3 21.09 20.98 23.28 26.58 27.08 26.11 24.90 20.43 20.63 22.08 23.34 FeO 0.15 0.00 0.19 0.00 0.39 0.00 0.25 0.00 0.00 0.38 0.21 CaO 2.08 1.91 5.38 8.41 9.17 8.03 6.24 0.79 1.56 3.81 6.26

Na2O 9.97 10.22 8.00 6.44 6.03 6.73 7.84 10.87 10.92 9.38 7.59 K2O 0.78 0.42 0.47 0.42 0.26 0.22 0.23 0.27 0.19 0.37 0.69 Total 99.88 99.40 98.34 98.47 99.44 99.44 99.96 99.42 100.66 100.15 98.90 Cations Si 2.902 2.911 2.755 2.577 2.553 2.621 2.692 2.951 2.934 2.834 2.737 Al 1.096 1.093 1.239 1.426 1.442 1.382 1.308 1.060 1.061 1.150 1.242 Fe 0.006 0.000 0.007 0.000 0.015 0.000 0.009 0.000 0.000 0.014 0.008 Ca 0.098 0.090 0.260 0.410 0.444 0.386 0.298 0.037 0.073 0.180 0.303 Na 0.852 0.875 0.700 0.568 0.528 0.586 0.678 0.927 0.924 0.804 0.664 K 0.044 0.024 0.027 0.024 0.015 0.013 0.013 0.015 0.011 0.021 0.040 Total 4.998 4.993 4.989 5.006 4.997 4.988 4.999 4.990 5.002 5.003 4.994 An 9.881 9.137 26.351 40.897 44.970 39.227 30.143 3.801 7.240 17.951 30.072 Ab 85.707 88.471 70.908 56.671 53.512 59.494 68.534 94.652 91.710 79.974 65.981 Or 4.412 2.392 2.741 2.432 1.518 1.280 1.323 1.547 1.050 2.076 3.947 286

Chaelundi Complex A-type Suite Mineral Analyses - Chapter 5 (continued).

Alkali Feldspar Sample CC128 CC129 CC131 CC132 CC134 CC135 CC136 CC57 CC64 CC66

SiO2 64.98 63.88 65.35 63.69 64.97 65.48 65.16 64.23 62.34 64.64 Al2O3 18.43 19.30 18.16 18.70 18.40 19.07 18.31 17.72 17.44 18.56 FeO 0.13 0.00 0.00 0.00 0.00 0.00 0.13 0.00 0.13 0.46 CaO 0.13 0.00 0.18 0.12 0.00 0.14 0.00 0.10 0.15 0.46

Na2O 2.95 1.37 2.72 3.27 2.22 4.88 1.18 1.47 0.56 3.70 K2O 12.40 14.68 12.86 12.36 13.85 10.03 15.93 15.14 15.81 11.29 Total 99.02 99.23 99.51 98.77 99.44 99.60 100.82 98.81 96.68 99.11 Cations Si 2.995 2.961 3.007 2.969 2.997 2.977 2.994 3.005 2.997 2.976 Al 1.001 1.054 0.985 1.027 1.000 1.022 0.992 0.977 0.988 1.007 Fe 0.005 0.000 0.000 0.000 0.000 0.000 0.005 0.000 0.005 0.018 Ca 0.006 0.000 0.009 0.006 0.000 0.007 0.000 0.005 0.008 0.023 Na 0.264 0.123 0.243 0.296 0.199 0.430 0.105 0.133 0.052 0.330 K 0.729 0.868 0.755 0.735 0.815 0.582 0.934 0.904 0.970 0.663 Total 5.001 5.007 4.999 5.033 5.010 5.018 5.030 5.025 5.020 5.017 An 0.643 0.000 0.882 0.578 0.000 0.669 0.000 0.481 0.750 2.233 Ab 26.385 12.422 24.112 28.512 19.589 42.227 10.119 12.797 5.070 32.506 Or 72.972 87.578 75.007 70.909 80.411 57.104 89.881 86.722 94.179 65.261 Ilmenite Sample CC131 CC132 CC133 CC134 CC66

SiO2 0.46 0.32 0.39 0.50 0.24 TiO2 52.31 53.08 53.48 53.83 50.65 Al2O3 0.00 0.00 0.00 0.00 0.00 FeO 43.46 45.04 41.73 43.44 44.65 MnO 3.90 1.91 5.75 4.43 3.38 MgO 0.00 0.00 0.00 0.00 0.00 Total 100.13 100.35 101.35 102.20 98.92 Cations Si 0.023 0.016 0.019 0.025 0.012 Ti 1.980 2.000 1.995 1.990 1.956 Al 0.000 0.000 0.000 0.000 0.000 Fe2+ 1.828 1.887 1.730 1.786 1.917 Mn 0.166 0.081 0.241 0.184 0.147 Mg 0.000 0.000 0.000 0.000 0.000 Total 3.997 3.984 3.986 3.985 4.032 287

Mineral Analyses - Woodlands Quartz Monzonite & enclaves - Chapter 6.

Pyroxenes Clinopyroxene Orthopyroxene WQMWQMX10 WQMX2 WQMX3 WQMX5 WQMX6 WQMX7 WQMX8 WQMX2 WQMX6

SiO2 51.05 51.15 50.75 50.62 50.13 51.40 51.24 49.95 53.28 53.55 TiO2 0.62 0.52 0.88 0.88 1.05 0.79 0.59 1.01 0.35 0.36 Al2O3 2.33 1.77 2.47 2.62 2.44 2.69 2.56 2.82 1.27 1.16 FeO 8.69 13.80 9.31 9.25 8.65 9.16 9.27 8.36 16.99 16.67 MnO 0.38 0.59 0.29 0.32 0.24 0.25 0.46 0.26 0.51 0.50 MgO 15.18 12.89 15.37 14.92 15.71 15.23 14.92 14.89 26.10 26.54 CaO 20.60 19.09 20.67 20.80 19.83 20.49 20.49 20.87 1.49 1.52

Na2O 0.41 0.29 0.00 0.46 0.00 0.29 0.37 0.00 0.00 0.00 Total 99.26 100.10 99.74 99.87 98.05 100.30 99.90 98.16 99.99 100.30 Cations Si 1.916 1.940 1.900 1.896 1.901 1.909 1.914 1.895 1.944 1.944 AlIV 0.084 0.060 0.100 0.104 0.099 0.091 0.086 0.105 0.055 0.050 AlVI 0.019 0.019 0.009 0.012 0.010 0.027 0.027 0.021 0.000 0.000 Fe3+ 0.037 0.020 0.026 0.047 0.019 0.026 0.033 0.017 0.024 0.026 Fe2+ 0.206 0.402 0.245 0.205 0.240 0.238 0.230 0.235 0.475 0.459 Mn 0.012 0.019 0.009 0.010 0.008 0.008 0.015 0.008 0.016 0.015 Mg 0.849 0.729 0.858 0.833 0.888 0.843 0.831 0.842 1.419 1.437 Ti 0.018 0.015 0.025 0.025 0.030 0.022 0.017 0.029 0.010 0.010 Ca 0.828 0.776 0.829 0.835 0.805 0.815 0.820 0.848 0.058 0.059 Na 0.030 0.021 0.000 0.033 0.000 0.021 0.027 0.000 0.000 0.000 K 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Total 3.999 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 4.000 En 44.22 37.84 43.82 43.39 45.48 43.87 43.41 43.37 71.81 72.52 Fs 12.66 21.89 13.83 13.13 13.27 13.72 13.75 12.96 25.25 24.50 Wo 43.12 40.27 42.35 43.47 41.25 42.41 42.84 43.68 2.95 2.98 Mg# 76.89 62.32 75.40 76.06 76.90 75.64 74.95 76.41 73.38 74.15

Secondary amphiboles (after pyroxenes) Actinolite after CPX Cummingtonite (after OPX) Sample WQMX6 WQMX5 CC15 CC15 CC133 CC133 CC134 WQMX10 WQMX2 WQMX6

SiO2 53.33 51.07 51.33 50.30 52.87 53.16 50.32 52.01 52.82 53.34 TiO2 0.15 0.12 0.12 0.37 0.21 0.17 0.33 0.00 0.32 0.00 Al2O3 0.73 1.32 0.39 1.14 0.83 0.62 1.38 0.60 1.14 0.66 FeO 19.10 20.13 11.42 14.19 16.33 17.37 23.21 27.26 20.56 22.07 MnO 0.51 0.81 0.30 0.62 0.84 0.64 1.12 1.40 0.85 0.90 MgO 14.09 10.25 12.50 11.33 13.30 13.26 8.92 13.63 18.05 17.20 CaO 9.36 11.24 21.87 18.94 11.41 11.73 10.54 1.47 1.84 2.49

Na2O 0.21 0.28 0.18 0.23 0.40 0.55 0.63 0.33 0.00 0.00 K2O 0.00 0.19 0.00 0.07 0.07 0.13 0.23 0.00 0.00 0.00 Total 97.48 95.41 98.11 97.19 96.26 97.63 96.68 96.70 95.58 96.66 Cations Si 7.872 7.837 7.589 7.561 7.877 7.853 7.753 7.908 7.841 7.896 AlIV 0.127 0.163 0.068 0.202 0.123 0.108 0.247 0.092 0.159 0.104 AlVI 0.000 0.075 0.000 0.000 0.023 0.000 0.004 0.015 0.041 0.011 Fe3+ 0.166 0.166 0.166 0.166 0.166 0.166 0.166 0.109 0.029 0.058 Fe2+ 2.980 2.980 2.980 2.980 2.980 2.980 2.980 3.270 2.500 2.627 Mn 0.064 0.105 0.038 0.079 0.106 0.080 0.146 0.180 0.107 0.113 Mg 3.101 2.345 2.755 2.539 2.954 2.920 2.049 3.089 3.995 3.796 Ti 0.017 0.014 0.013 0.042 0.024 0.019 0.038 0.000 0.036 0.000 Ca 1.480 1.848 3.464 3.050 1.821 1.856 1.740 0.239 0.293 0.35 Na 0.060 0.083 0.052 0.067 0.116 0.158 0.188 0.097 0.000 0.000 K 0.000 0.037 0.000 0.013 0.013 0.024 0.045 0.000 0.000 0.000 Total 15.436 15.436 15.436 15.436 15.436 15.436 15.436 15.000 15.000 15.000 Mg# 31.433 31.433 31.433 31.433 31.433 31.433 31.433 46.47 60.25 57.56 Na (A site) 0.000 0.000 1.515 1.117 0.000 0.014 0.000 0.000 0.000 0.000 K (A site) 0.000 0.000 0.000 0.013 0.000 0.024 0.000 0.000 0.000 0.000 288

Mineral Analyses - Woodlands Quartz Monzonite & enclaves - Chapter 6 (continued).

Biotite

Sample WQM WQMX10 WQMX2 WQMX3 WQMX5 WQMX6

SiO2 33.62 35.46 35.57 35.93 34.68 35.59 TiO2 3.95 4.21 3.15 3.28 4.06 3.64 Al2O3 12.45 12.37 12.02 12.35 13.19 12.36 FeO 31.12 27.55 27.34 27.83 25.34 27.02 MnO 0.75 0.35 0.18 0.25 0.43 0.15 MgO 3.93 6.63 6.92 7.20 7.69 7.61 CaO 0.00 0.00 0.00 0.00 0.00 0.00

Na2O 0.28 0.31 0.00 0.00 0.00 0.00 K2O 8.84 9.34 9.04 8.87 9.52 9.10 Total 94.94 96.22 94.22 95.71 94.91 95.47 Cations Si 5.54 5.63 5.74 5.71 5.53 5.66 AlIV 2.42 2.31 2.26 2.29 2.47 2.32 AlVI 0.00 0.00 0.03 0.02 0.01 0.00 Ti 0.49 0.50 0.38 0.39 0.49 0.44 Fe2+ 4.28 3.66 3.69 3.70 3.38 3.59 Mn 0.10 0.05 0.02 0.03 0.06 0.02 Mg 0.96 1.57 1.67 1.71 1.83 1.80 Na 0.09 0.10 0.00 0.00 0.00 0.00 K 1.86 1.89 1.86 1.80 1.94 1.85 Total 15.74 15.70 15.66 15.64 15.71 15.67 Mg# 18.38 30.02 31.09 31.57 35.11 33.43

Hornblende Sample WQM WQMX10 WQMX2 WQMX3 WQMX5 WQMX6 WQMX7 WQMX8 WQMX10

SiO2 42.99 45.88 42.88 44.98 46.79 46.21 45.63 46.16 42.34 TiO2 1.20 0.93 3.02 1.36 1.15 1.64 1.39 1.12 2.78 Al2O3 6.53 4.71 7.26 5.65 4.38 5.93 5.59 4.71 9.55 FeO 28.39 27.26 24.62 24.55 19.96 21.95 23.40 20.62 19.81 MnO 1.23 1.03 0.47 0.79 0.91 0.37 0.57 0.79 0.37 MgO 4.26 6.08 6.78 6.72 10.49 8.74 7.87 9.21 9.00 CaO 10.24 9.60 10.50 10.67 11.01 10.95 10.57 10.80 10.54

Na2O 1.57 1.20 1.64 1.30 1.07 0.94 1.62 1.35 2.40 K2O 1.02 0.49 0.58 0.65 0.59 0.00 0.61 0.60 0.69 Total 97.43 97.18 97.74 96.67 96.35 96.73 97.25 95.36 97.48 Cations Si 6.885 7.229 6.689 7.071 7.205 7.099 7.077 7.211 6.488 AlIV 1.115 0.771 1.311 0.929 0.795 0.901 0.923 0.789 1.512 AlVI 0.118 0.103 0.024 0.117 0.000 0.173 0.099 0.078 0.213 Fe3+ 0.311 0.463 0.286 0.231 0.288 0.290 0.237 0.190 0.219 Fe2+ 3.242 2.758 2.696 2.811 2.052 2.298 2.608 2.352 2.144 Mn 0.167 0.137 0.062 0.105 0.119 0.048 0.075 0.105 0.048 Mg 1.017 1.428 1.577 1.575 2.408 2.002 1.820 2.145 2.056 Ti 0.145 0.110 0.354 0.161 0.133 0.190 0.162 0.132 0.320 Ca 1.757 1.620 1.754 1.797 1.816 1.802 1.756 1.807 1.730 Na 0.487 0.367 0.495 0.396 0.319 0.280 0.487 0.409 0.713 K 0.208 0.098 0.116 0.130 0.116 0.000 0.121 0.120 0.135 Total 15.453 15.086 15.366 15.324 15.252 15.082 15.364 15.336 15.578 Mg# 21.47 29.83 34.13 33.35 49.48 43.16 38.39 44.77 46.02 Na (A site) 0.245 0.000 0.250 0.193 0.136 0.082 0.244 0.217 0.443 K (A site) 0.208 0.086 0.116 0.130 0.116 0.000 0.121 0.120 0.135 289

Mineral Analyses - Woodlands Quartz Monzonite & enclaves - Chapter 6 (continued).

Plagioclase

Host ------Enclave Matrix------Phenocrysts------Xenocryst Sample WQM WQMX10 WQMX3 WQMX8 WQMX10 WQMX2 WQMX3 WQMX5 WQMX6 WQMX7 WQMX8

SiO2 58.86 61.26 64.91 64.02 54.02 46.42 51.96 51.27 45.38 46.88 60.16 Al2O3 24.55 23.31 21.49 21.37 28.36 33.91 30.24 29.98 33.84 32.88 23.22 FeO 0.29 0.35 0.32 0.28 0.24 0.60 0.16 0.35 0.61 0.62 0.31 CaO 6.84 4.55 2.52 2.58 11.00 17.46 13.19 12.64 18.25 16.86 5.26

Na2O 7.13 8.81 10.00 10.17 5.33 1.46 4.10 4.00 1.24 2.01 8.21 K2O 0.44 0.25 0.14 0.13 0.17 0.00 0.09 0.13 0.00 0.00 0.12 Total 98.11 98.53 99.50 98.55 99.12 99.85 99.87 98.37 99.32 99.25 97.28 Cations Si 2.677 2.760 2.876 2.865 2.463 2.142 2.366 2.366 2.113 2.175 2.746 Al 1.316 1.237 1.122 1.127 1.524 1.844 1.623 1.631 1.857 1.798 1.249 Fe 0.011 0.013 0.012 0.010 0.009 0.023 0.006 0.014 0.024 0.024 0.012 Ca 0.333 0.220 0.120 0.124 0.537 0.863 0.644 0.625 0.910 0.838 0.257 Na 0.629 0.769 0.859 0.882 0.471 0.131 0.362 0.358 0.112 0.181 0.726 K 0.026 0.014 0.008 0.007 0.010 0.000 0.005 0.008 0.000 0.000 0.007 Total 4.992 5.014 4.996 5.016 5.015 5.002 5.006 5.001 5.015 5.016 4.997 An 33.75 21.89 12.13 12.21 52.76 86.86 63.67 63.10 89.05 82.25 25.96 Ab 63.66 76.68 87.07 87.06 46.27 13.14 35.81 36.13 10.95 17.75 73.33 Or 2.59 1.43 0.80 0.73 0.97 0.00 0.52 0.77 0.00 0.00 0.71 Alkali Feldspar Fe-Ti Oxides Host ----Xenocrysts------Ilmenite------Titanomagnetite Sample WQM WQMX5 WQMX8 Sample WQMX10 WQMX3 WQMX5 WQMX6 WQMX10 WQMX3

SiO2 65.16 65.11 63.74 SiO2 0.30 0.42 0.39 0.35 0.49 0.53 Al2O3 18.98 18.09 18.48 TiO2 52.58 52.62 53.12 50.38 11.10 25.24 FeO 0.00 0.14 0.13 Al2O3 0.00 0.15 0.00 0.30 0.14 0.27 CaO 0.13 0.00 0.21 FeO 46.60 44.81 42.35 48.09 82.71 69.90

Na2O 3.31 5.14 3.18 MnO 2.19 2.75 4.20 1.54 0.68 1.36 K2O 11.75 9.42 12.32 MgO 0.00 0.00 0.26 0.15 0.00 0.26 Total 99.65 98.10 98.43 Total 101.67 100.75 100.32 100.81 95.12 97.56 Cations Si 2.984 3.006 2.975 Si 0.015 0.021 0.019 0.018 0.034 0.032 Al 1.024 0.984 1.016 Ti 1.969 1.978 1.997 1.914 0.571 1.145 Fe 0.000 0.005 0.005 Al 0.000 0.009 0.000 0.018 0.011 0.019 Ca 0.006 0.000 0.010 Fe3+ 0.020 0.000 0.000 0.074 1.737 1.017 Na 0.294 0.460 0.288 Fe2+ 1.904 1.873 1.770 1.899 1.608 1.695 K 0.686 0.555 0.733 Mn 0.092 0.116 0.178 0.066 0.039 0.069 Total 4.994 5.010 5.028 Mg 0.000 0.000 0.019 0.011 0.000 0.023 An 0.65 0.00 1.02 Total 4.000 3.997 3.984 4.000 4.000 4.000 Ab 29.79 45.33 27.89 Or 69.57 54.67 71.09

Broad beam analyses of enclave matrix

Sample WQMX2 WQMX5 WQMX8 WQMX8 WQMX7 (margin)

SiO2 55.91 57.90 45.51 59.25 59.32 TiO2 0.70 0.98 1.15 1.12 0.55 Al2O3 14.05 15.85 12.31 14.87 16.35 FeO 10.49 5.11 4.47 5.87 5.79 MnO 0.29 0.12 0.11 0.31 0.13 MgO 3.18 0.93 1.06 1.18 1.30 CaO 6.96 2.97 3.36 3.96 3.70

Na2O 4.02 4.19 4.12 4.23 6.59 K2O 2.77 5.89 2.69 4.65 3.22 Total 98.37 93.94 74.78 95.44 96.95 290

Appendix E

Whole-rock geochemical analyses. 291

Whole-rock analyses - Hillgrove Supersuite - Chapter 4.

Sample C.GR. N2 N3 A60 N42 N4 N46 N45 G37 E29 E27 A243 Grid ref 513629 511338 517327 915213 486321 483317 477578 474573 036367 039627 048353 974713 Pluton SSCC* Rockisle Rockisle Hillgrove Rockisle Rockisle Kilburnie Kilburnie RockvaleTobermory Rockvale Argyle Suite ------Rockisle ------Hillgrove ------

SiO2 67.39 67.89 68.52 69.47 70.38 70.88 72.37 73.62 67.22 67.66 67.96 68.71 TiO2 0.59 0.44 0.40 0.60 0.36 0.39 0.48 0.38 0.70 0.69 0.67 0.62 Al2O3 15.43 16.09 15.77 14.11 14.88 14.59 13.04 12.82 14.96 14.78 14.78 14.55 Fe2O3 0.69 0.42 0.42 0.50 0.36 0.50 0.53 0.41 0.87 0.64 0.46 0.51 FeO 2.84 2.40 2.32 3.03 2.03 2.00 2.53 1.99 3.76 3.74 3.89 3.49 MnO 0.07 0.06 0.06 0.07 0.05 0.06 0.05 0.03 0.11 0.09 0.10 0.11 MgO 1.22 1.63 1.52 1.28 1.13 1.22 0.46 0.36 1.63 1.62 1.63 1.32 CaO 3.49 3.33 3.00 2.28 2.57 2.54 1.66 1.33 3.45 2.40 2.49 2.34

Na2O 4.25 4.25 4.18 3.04 3.86 3.67 3.32 3.11 3.52 3.13 2.97 3.23 K2O 2.29 2.18 2.59 4.21 3.27 3.06 4.16 4.60 2.00 3.65 3.52 3.79 P2O5 0.13 0.11 0.11 0.13 0.09 0.10 0.09 0.07 0.16 0.16 0.16 0.13 LOI 1.27 0.81 0.72 0.79 0.61 0.57 0.67 0.74 1.20 0.92 0.87 0.65 REST 0.21 0.22 0.23 0.30 0.27 0.19 0.31 0.40 0.25 0.26 0.31 0.26 Total 99.88 99.82 99.85 99.78 99.87 99.77 99.67 99.85 99.81 99.74 99.81 99.72

Trace Elements (ppm) Ba 312 432 441 609 491 415 679 741 400 606 683 701 Rb 92 71 88 182 111 117 175 164 92 145 138 141 Sr 162 418 338 141 243 220 104 97 352 183 187 227 Pb 12 10 16 25 15 13 21 20 16 24 23 22 Th 13 12 11 21 12 14 16 16 15 16 16 14 U 105010203441 Zr 205 172 169 226 152 151 343 293 228 215 205 208 Hf 533754965655 Nb 9 9 914109151116161116 Y 30 19 22 29 21 21 53 152 34 35 35 29 La 15 32 4 35 9 52 6 68.10 25.50 15 17 32 Ce 35 38 31 37 32 57 29 139.00 55.00 48 43 34 Nd ------76.30 27.90 - - - Sm ------19.106.32--- Eu ------1.361.08--- Tb ------3.390.84--- Ho ------4.381.17--- Yb ------9.252.87--- Lu ------1.310.45--- V 61 51 36 58 34 44 20 19 75 60 59 61 Cr 92826242323121328323331 Ni 056225004653 Zn 46 46 46 57 37 61 61 48 55 73 72 70 Ga 16 17 18 16 17 15 18 15 16 17 16 17 F 732 477 739 1123 1169 424 1017 1642 641 789 1145 604 Cl 103 61 70 46 54 31 234 184 91 40 125 65

CIPW NORMS Q 23.49 24.13 24.6 27.65 27.52 29.75 32.04 34.13 27.54 26.18 27.63 26.65 C 0.06 0.98 1.01 0.93 0.76 0.9 0.49 0.85 1.19 1.81 2.16 1.25 Or 13.61 12.96 15.4 25.07 19.45 18.2 24.81 27.39 11.9 21.74 20.96 22.57 Ab 36.01 36.06 35.46 25.82 32.76 31.16 28.1 26.32 29.84 26.59 25.16 27.43 An 16.21 15.86 13.99 9.99 11.63 11.95 7.23 5.24 16.01 10.67 10.88 10.73 Hy 6.85 7.52 7.17 7.5 5.76 5.78 4.68 3.66 9.32 9.44 9.94 8.49 Mt 1.02 0.62 0.62 0.74 0.53 0.74 0.78 0.6 1.28 0.95 0.68 0.76 Il 1.12 0.84 0.76 1.14 0.69 0.74 0.92 0.73 1.33 1.32 1.28 1.18 Ap 0.31 0.26 0.26 0.31 0.22 0.24 0.22 0.17 0.38 0.39 0.39 0.31 Mg# 43.36 54.76 53.86 42.95 49.8 52.08 24.47 24.38 43.58 43.56 42.75 40.26

*SSCC = Sheep Station Creek Complex 292

Whole-rock analyses - Hillgrove Supersuite - Chapter 4 (continued).

Sample A274 E28 A183 A246 A278 D15 A221 A283 A285 A277 Y94 D21 Grid ref 823978 046357 818145 954189 942175 600543 693069 932205 977184 826144 923684 558598 Pluton Enmore Rockvale Gara Hillgrove Hillgrove Dund.Gostwyck Hillgrove Hillgrove Gara Kim. Pk. Dund. Suite------Hillgrove ------

SiO2 68.77 68.83 68.94 68.97 69.14 69.17 69.27 69.31 69.32 69.34 69.34 69.41 TiO2 0.64 0.56 0.64 0.62 0.62 0.63 0.68 0.61 0.63 0.62 0.66 0.63 Al2O3 14.49 14.53 14.14 14.44 14.36 14.10 14.11 14.45 14.40 14.24 14.38 14.09 Fe2O3 0.87 0.45 0.46 0.54 0.40 0.41 0.59 0.54 0.42 0.50 0.47 0.44 FeO 3.10 3.17 3.24 3.29 3.44 3.54 3.34 3.13 3.36 3.06 3.24 3.46 MnO 0.09 0.08 0.08 0.08 0.08 0.11 0.08 0.08 0.08 0.08 0.09 0.10 MgO 1.36 1.33 1.14 1.31 1.35 1.19 1.08 1.26 1.32 1.09 0.83 1.13 CaO 2.21 2.11 2.43 2.27 2.22 2.32 2.39 2.20 2.27 2.37 2.39 2.24

Na2O 3.11 3.11 3.08 3.09 3.00 3.47 3.23 3.03 2.99 3.05 3.67 3.34 K2O 3.65 3.87 3.86 3.94 3.95 3.64 3.75 4.00 3.85 3.88 3.55 3.65 P2O5 0.14 0.12 0.14 0.14 0.14 0.15 0.15 0.13 0.13 0.13 0.14 0.14 LOI 1.13 0.96 1.17 0.87 0.86 0.80 0.67 0.75 0.68 1.13 0.61 0.90 REST 0.28 0.26 0.35 0.31 0.27 0.27 0.32 0.28 0.29 0.32 0.41 0.26 Total 99.82 99.40 99.68 99.87 99.83 99.79 99.67 99.77 99.73 99.81 99.78 99.79

Trace Elements (ppm) Ba 661 590 795 603 557 586 706 626 585 815 615 561 Rb 152 163 159 165 165 142 144 166 156 156 140 141 Sr 184 165 161 154 153 159 165 152 152 159 240 159 Pb 22 30 19 24 22 31 19 20 23 23 20 22 Th 18 22 19 15 18 19 13 15 16 17 17 18 U 256423442234 Zr 222 204 250 228 229 239 298 225 221 238 230 233 Hf 344576878778 Nb 14138141212131512121411 Y 35 37 35 36 35 39 41 37 36 32 32 43 La 19 24 37 24 20 10 27 44 17 18 32 32 Ce 43 52 36 53 42 45 44 51 36 45 51 41 Nd ------Sm ------Eu ------Tb ------Ho ------Yb ------Lu ------V 51 63 51 57 72 58 48 59 60 46 46 53 Cr 30 30 22 33 27 26 22 50 29 37 11 14 Ni 4700420811100 Zn 68 68 60 64 64 65 63 64 64 57 63 65 Ga 15 16 17 17 18 16 15 16 16 16 16 17 F 836 734 1454 1241 836 787 1194 860 1142 1162 2251 798 Cl 80 101 82 63 73 101 49 60 48 70 34 133

CIPW NORMS Q 28.75 28.01 28.13 27.73 28.14 26.58 28.11 28.42 28.87 28.81 27.06 27.88 C 1.87 1.8 1.1 1.54 1.59 0.71 0.97 1.59 1.72 1.23 0.99 1.04 Or 21.72 23.13 23.03 23.45 23.51 21.67 22.36 23.82 22.94 23.11 21.17 21.73 Ab 26.37 26.47 26.18 26.21 25.44 29.42 27.48 25.72 25.4 25.89 31.23 28.3 An 9.81 9.53 10.51 9.79 9.83 10.31 10.42 9.79 9.93 10.47 9.7 9.96 Hy 7.49 8.05 7.54 8.01 8.5 8.32 7.4 7.61 8.25 7.07 6.74 7.98 Mt 1.28 0.67 0.68 0.8 0.6 0.61 0.87 0.8 0.63 0.74 0.7 0.65 Il 1.22 1.07 1.22 1.18 1.18 1.2 1.3 1.16 1.2 1.18 1.26 1.2 Ap 0.34 0.29 0.34 0.34 0.34 0.36 0.36 0.31 0.31 0.31 0.34 0.34 Mg# 43.88 42.78 38.54 41.5 41.15 37.46 36.56 41.77 41.18 38.83 31.34 36.79

Kim. Pk. = Kimberley Park Dund. = Dundurrabin 293

Whole-rock analyses - Hillgrove Supersuite - Chapter 4 (continued).

Sample A87 A185 D11 D18 NFG A180 A184 G31 Y62 A273 E6 N43 Grid ref 983718 846146 653507 656549 519622 868164 822145 950563 760456 815978 077782 500549 Pluton Argyle Gara Dund. Dund. SSCC Gara GaraTobermory Tia Enmore Kook. Kilburnie Suite------Hillgrove ------

SiO2 69.42 69.44 69.74 69.79 69.79 69.81 69.92 69.96 70.00 70.15 70.24 70.43 TiO2 0.70 0.63 0.61 0.60 0.62 0.68 0.57 0.54 0.50 0.60 0.56 0.55 Al2O3 14.14 14.24 14.06 14.20 14.01 14.16 13.94 14.52 14.50 14.15 13.81 14.05 Fe2O3 0.77 0.66 0.42 0.44 0.68 0.63 0.37 0.35 0.40 0.42 0.51 0.63 FeO 2.88 3.02 3.34 3.33 3.18 3.20 2.88 2.89 2.82 2.98 2.95 2.53 MnO 0.08 0.08 0.10 0.11 0.10 0.08 0.07 0.07 0.09 0.07 0.08 0.06 MgO 1.05 1.12 1.06 1.12 1.12 1.15 0.99 1.11 1.22 1.06 1.09 0.92 CaO 2.26 2.28 2.21 2.00 2.05 2.37 2.12 2.13 2.25 2.01 1.62 1.95

Na2O 3.38 3.01 3.41 3.15 3.31 3.17 3.07 3.16 3.40 3.08 3.18 3.27 K2O 3.94 3.74 3.63 3.79 3.62 3.87 4.07 3.94 3.62 4.09 4.26 4.18 P2O5 0.14 0.13 0.14 0.13 0.13 0.14 0.12 0.14 0.13 0.13 0.12 0.13 LOI 0.76 1.12 0.76 0.84 0.85 1.42 1.33 0.78 0.68 0.83 1.09 0.89 REST 0.30 0.29 0.30 0.28 0.32 0.33 0.34 0.29 0.27 0.28 0.35 0.28 Total 99.82 99.78 99.77 99.78 99.80 101.00 99.79 99.88 99.88 99.84 99.83 99.87

Trace Elements (ppm) Ba 716 641 601 640 660 730 629 616 541 653 590 659 Rb 153 147 140 153 140 164 169 166 139 167 167 174 Sr 174 159 154 169 156 155 143 156 181 153 146 141 Pb 23 23 24 23 19 26 25 19 22 26 21 22 Th 17 17 13 16 16 17 17 12 16 19 18 19 U 574205101241 Zr 281 232 232 230 235 267 242 221 196 253 233 243 Hf 886777866676 Nb 14128131113121212101013 Y 38 35 39 40 144 38 33 37 31 36 39 36 La 14 30 27 26 48 26.90 32 21 7 30 10 13 Ce 40 52 46 46 32 60.60 46 38 40 41 49 44 Nd - ----30.40------Sm -----6.58------Eu -----0.89------Tb -----1.02------Ho -----1.45------Yb -----3.84------Lu -----0.56------V 57 48 57 64 58 56 49 41 50 47 52 52 Cr 16 20 18 20 20 21 22 20 24 23 24 20 Ni 000013033000 Zn 49 59 61 67 69 59 52 55 56 57 58 48 Ga 17 17 17 14 16 16 16 19 18 14 15 17 F 911 1091 1194 824 1092 1258 1548 1116 975 835 1761 962 Cl 147 43 110 111 102 43 58 105 64 83 15 44

CIPW NORMS Q 27.33 29.93 28.26 29.42 29.36 28.22 29.43 29.1 28.49 29.39 29.31 29 C 0.69 1.61 1.07 1.73 1.44 1.01 1.24 1.73 1.38 1.45 1.69 1.07 Or 23.45 22.28 21.62 22.57 21.56 22.78 24.25 23.44 21.53 24.34 25.36 24.87 Ab 28.63 25.57 28.92 26.71 28.08 26.61 26.08 26.77 28.83 26.11 27.03 27.75 An 10.02 10.03 9.53 8.83 8.89 10.19 8.95 9.18 9.93 8.86 6.28 8.47 Hy 6.28 6.93 7.63 7.78 7.26 7.19 6.66 7.04 7.25 6.92 6.96 5.63 Mt 1.13 0.97 0.62 0.65 1 0.92 0.55 0.52 0.59 0.62 0.75 0.93 Il 1.34 1.2 1.16 1.15 1.18 1.28 1.09 1.03 0.95 1.14 1.07 1.05 Ap 0.34 0.31 0.34 0.31 0.31 0.33 0.29 0.34 0.31 0.31 0.29 0.31 Mg# 39.38 39.79 36.12 37.47 38.56 39.04 37.99 40.63 43.53 38.8 39.7 39.32

Dund. = Dundurrabin 294

Whole-rock analyses - Hillgrove Supersuite - Chapter 4 (continued).

Sample A282 A276 C120 A284 N15 N44 A7 GUM.A A60B A286 Grid ref 900211 870017 063232 922182 563153 493546 845874 508641 915213 953208 Pluton Hillgrove Blue K. Abroi HillgroveMurder D. Kilburnie Winterb. SSCC Hillgrove Hillgrove Suite ------Hillgrove ------

SiO2 70.61 70.79 70.86 71.05 71.23 71.31 71.90 72.79 74.86 76.88 TiO2 0.53 0.53 0.44 0.50 0.52 0.49 0.40 0.34 0.22 0.10 Al2O3 14.08 14.10 14.22 13.99 13.92 13.96 14.12 13.62 13.07 12.79 Fe2O3 0.34 0.39 0.44 0.29 0.37 0.42 0.41 0.59 0.33 0.55 FeO 2.71 2.90 2.58 2.56 2.44 2.37 1.87 1.86 1.10 0.18 MnO 0.06 0.09 0.06 0.06 0.06 0.06 0.05 0.04 0.03 0.01 MgO 1.03 1.05 1.00 0.99 0.81 0.79 0.79 0.52 0.33 0.03 CaO 2.17 1.71 1.65 2.07 2.00 1.82 1.81 1.66 1.03 0.86

Na2O 2.99 2.91 3.15 2.92 3.29 3.15 3.44 3.96 2.97 3.74 K2O 4.39 4.10 4.12 4.40 4.21 4.33 3.83 3.31 5.14 4.03 P2O5 0.11 0.13 0.13 0.10 0.13 0.11 0.09 0.07 0.05 0.03 LOI 0.64 0.75 0.97 0.60 0.58 0.73 0.81 0.85 0.52 0.49 REST 0.24 0.28 0.22 0.27 0.23 0.27 0.29 0.24 0.20 0.25 Total 99.90 99.74 99.84 99.79 99.78 99.80 99.82 99.84 99.84 99.93

Trace Elements (ppm) Ba 587 643 610 594 479 633 654 487 260 221 Rb 198 173 141 199 170 171 139 104 280 205 Sr 128 163 183 120 122 145 213 83 60 123 Pb 24 23 25 29 21 24 18 9 28 20 Th 23 18 12 21 19 17 15 18 36 39 U 34224103128 Zr 217 204 174 211 227 221 175 215 139 127 Hf 5655563623 Nb 91310111212139128 Y 36 35 30 40 33 36 28 37 54 32 La 18 31 40 27.70 30 14 4 32 28 31 Ce 53 57 31 62.50 40 44 36 40 52 49 Nd ---31.70------Sm ---6.66------Eu ---0.74------Tb ---1.09------Ho ---1.45------Yb ---3.60------Lu ---0.54------V 44 50 51 41 46 44 33 26 24 6 Cr 25 22 22 18 17 17 16 8 7 4 Ni 0400000000 Zn 50 71 53 47 43 46 35 35 29 6 Ga 15 17 14 15 15 16 18 15 13 13 F 679 884 500 953 739 959 1073 983 727 1473 Cl 50 52 20 44 26 27 187 49 27 20

CIPW NORMS Q 28.94 31.67 30.53 30.21 29.53 30.45 31.72 32.19 34.95 38.4 C 0.82 2.23 1.93 1.07 0.74 1.21 1.46 0.88 1.01 1.13 Or 26.11 24.43 24.49 26.2 25.05 25.77 22.8 19.68 30.59 23.97 Ab 25.35 24.72 26.74 24.79 27.94 26.76 29.11 33.6 25.2 31.73 An 9.86 7.34 7.31 9.23 8.83 7.94 7.95 7.29 4.39 3.12 Hy 6.51 6.94 6.26 6.23 5.46 5.29 4.51 3.75 2.27 0.07 Mt 0.5 0.58 0.65 0.43 0.55 0.62 0.6 0.86 0.49 0.32 Il 1.01 1.01 0.84 0.95 0.99 0.93 0.76 0.65 0.42 0.19 Ap 0.26 0.31 0.31 0.24 0.31 0.27 0.22 0.17 0.12 0.07 Mg# 40.38 39.22 40.85 40.8 37.17 37.26 42.95 33.25 34.84 22.9

Blue K. = Blue Knobby, Murder D. = Murder Dog, Winterb. = Winterbourne 295

Whole-rock analyses - Metasediments - Chapter 4.

Sample Y44 A30 A79 A4 A281 A3 W443 Y29 A181 A96 Lithology Greywacke Greywacke Greywacke Greywacke Greywacke Greywacke Pelite Greywacke Greywacke Argillite Grid ref 937471 654991 797710 765806 806199 739800 823622 797156 844200

SiO2 59.51 63.22 63.39 64.47 64.62 64.64 64.73 65.21 65.73 65.84 TiO2 0.74 0.91 0.79 0.61 0.62 0.73 0.73 0.59 0.63 0.73 Al2O3 17.48 14.37 14.29 16.08 15.74 15.07 17.14 16.67 15.56 14.61 Fe2O3 1.32 1.59 0.96 1.12 0.70 1.17 0.50 0.90 0.93 0.97 FeO 4.45 5.32 5.33 2.82 3.54 3.99 5.49 3.60 3.44 4.12 MnO 0.10 0.10 0.10 0.07 0.07 0.08 0.08 0.06 0.07 0.07 MgO 3.19 2.74 2.44 1.67 1.58 1.96 2.75 1.60 1.82 1.91 CaO 3.72 3.28 3.09 3.65 3.33 2.73 0.70 3.05 3.18 3.62

Na2O 5.13 3.02 3.65 4.07 3.79 3.88 1.23 3.95 4.25 3.06 K2O 0.74 1.77 1.45 2.81 2.28 2.57 4.71 2.48 1.77 2.33 P2O5 0.16 0.16 0.17 0.08 0.11 0.13 0.13 0.15 0.12 0.14 LOI 3.23 2.89 3.81 2.37 3.20 2.37 1.30 1.07 2.00 2.07 REST 0.29 0.30 0.34 0.37 0.28 0.26 0.33 0.25 0.35 0.30 Total 100.05 99.67 99.82 100.19 99.86 99.59 99.83 99.56 99.86 99.75

Trace Elements (ppm) **** Ba 258 377 391 675 563 451 652 455 470 409 Rb 19 56 27 93 70 74 193 64 54 78 Sr 909 426 501 498 551 350 93 405 552 415 Pb 1411518161327151212 Th 5969511257710 U 2112012010 Zr 125 187 151 192 166 198 176 193 172 172 Hf 4644665553 Nb 91181310820101111 Y 18 28 24 17 22 24 30 23 20 22 La 5 30 7 35 36 15 54 11 22 41 Ce 16 33 27 37 42 36 47 36 31 30 V 122 137 138 71 82 99 105 83 87 102 Cr 60 48 25 20 23 28 89 19 25 26 Ni 16 13 7 3 2 3 31 3 1 4 Zn 74 86 91 65 64 70 104 62 66 74 Ga 19 14 16 18 18 16 21 18 18 14 F 762 1163 1610 1505 757 835 1237 759 1551 1165 Cl 18 23 15 22 22 18 53 4 29 32

CIPW NORMS Q 11.89 25.35 23.73 18.76 22.55 21.69 31 22.2 23.39 26.7 C 1.91 2.11 1.82 0.17 1.32 1.38 9.4 2.4 1.45 1.01 Or 4.39 10.55 8.62 16.67 13.56 15.32 28.22 14.78 10.53 13.87 Ab 43.49 25.7 31.04 34.48 32.19 33.04 10.49 33.65 36.12 26.01 An 17.4 14.83 13.51 16.93 15.74 12.5 2.02 14.06 14.34 16.63 Hy 14 14.03 14.01 7.54 8.99 10.24 15.61 9.04 9.2 10.48 Mt 1.94 2.35 1.43 1.64 1.04 1.73 0.76 1.33 1.37 1.44 Il 1.41 1.74 1.51 1.16 1.18 1.4 1.4 1.13 1.2 1.39 Ap 0.38 0.38 0.41 0.19 0.26 0.31 0.32 0.36 0.29 0.34 Mg# 56.09 47.86 44.93 51.35 44.3 46.68 47.16 44.2 48.53 45.24 296

Whole-rock analyses - Metasediments - Chapter 4 (continued).

Sample A57 A56 A49 A275 A18 Y60 A138 A161 A12 Lithology Greywacke Greywacke Pelite Pelite Greywacke Argillite Pelite Greywacke Argillite Grid ref 843201 833203 868987 808007 718887 735538 834739 986074 746857

SiO2 65.91 66.47 67.81 69.19 69.34 69.49 70.81 70.95 72.40 TiO2 0.68 0.59 0.56 0.59 0.41 0.63 0.49 0.49 0.44 Al2O3 14.99 15.87 16.31 15.26 15.24 14.62 14.13 14.56 13.87 Fe2O3 0.71 0.92 2.43 3.15 0.75 0.67 0.91 0.53 1.95 FeO 3.91 3.02 2.03 1.35 1.89 3.34 2.77 2.73 1.10 MnO 0.07 0.06 0.07 0.04 0.03 0.08 0.09 0.05 0.05 MgO 1.96 1.53 1.40 1.01 0.80 1.17 1.17 1.06 0.67 CaO 3.34 2.86 0.06 0.13 1.84 1.69 1.17 1.89 0.11

Na2O 3.06 3.79 0.78 1.82 4.03 3.25 2.60 4.05 1.73 K2O 2.49 2.58 4.58 3.84 3.90 3.03 3.47 2.24 5.45 P2O5 0.14 0.11 0.05 0.08 0.07 0.14 0.10 0.11 0.05 LOI 2.28 1.78 3.36 3.18 1.39 1.41 1.86 0.91 1.73 REST 0.32 0.30 0.33 0.24 0.28 0.33 0.23 0.27 0.29 Total 99.86 99.89 99.77 99.87 99.96 99.85 99.77 99.84 99.85

Trace Elements (ppm) Ba 626 599 978 518 631 583 442 574 619 Rb 73 74 200 163 134 107 148 75 207 Sr 416 510 72 79 326 238 178 358 191 Pb 15 12 30 19 9 20 14 19 19 Th 91315181111171117 U 310255340 Zr 177 165 165 199 166 236 172 194 188 Hf 555665574 Nb 1091612111112714 Y 25 21 26 27 22 31 28 23 29 La 22869211422374229 Ce 32 26 80 37 37 32 42 43 50 V 99 77 102 84 39 71 87 68 58 Cr 22 20 40 24 26 30 51 24 32 Ni 3015003414 Zn 70 62 94 73 53 78 76 67 88 Ga 18 20 21 21 17 18 17 13 18 F 1197 1035 959 756 989 1470 608 806 1057 Cl 20 29 13 12 34 16 9 10 12

CIPW NORMS Q 26.69 24.94 43.44 42.51 25.27 32.35 36.98 31.99 40.61 C 1.72 2.06 10.12 8.13 1.38 3.56 4.31 2.41 5.14 Or 14.81 15.34 27.29 22.84 23.17 18.03 20.65 13.32 32.43 Ab 26 32.18 6.63 15.45 34.18 27.62 22.09 34.41 14.7 An 15.26 13.19 0 0 8.31 6.76 4.97 8.46 - Hy 10.51 7.75 4.42 2.52 4.23 7.63 6.62 6.51 1.68 Mt 1.06 1.36 3.56 2.78 1.1 0.99 1.34 0.79 2.44 Il 1.3 1.13 1.07 1.12 0.78 1.2 0.93 0.93 0.84 Ap 0.34 0.27 0.17 0.22 0.17 0.34 0.24 0.27 0.14 Mg# 47.18 47.45 55.14 57.14 43 38.43 42.94 40.9 52.05 297

Whole-rock analyses - Bundarra Suite - Chapter 4.

Sample BI3 BI2 B4 BI7 B11 B2 BI1 BI9 CB2 B9 Grid ref 005884 975911 190913 986902 301888 319903 963924 953803 991736 300123 Pluton Gwydir R. Gwydir R. Banalasta Gwydir R. Banalasta Banalasta Gwydir R. Gwydir R. Gwydir R. Banalasta

SiO2 71.78 71.83 71.98 72.21 72.85 72.90 72.90 73.07 73.91 74.67 TiO2 0.41 0.38 0.27 0.38 0.28 0.28 0.35 0.30 0.22 0.28 Al2O3 13.91 13.90 14.67 13.86 14.21 14.27 13.61 13.93 13.87 13.10 Fe2O3 0.24 0.22 0.20 0.20 0.22 0.19 0.20 0.11 0.25 0.17 FeO 2.09 2.10 1.54 2.09 1.71 1.62 1.96 1.80 1.52 1.40 MnO 0.05 0.06 0.04 0.05 0.04 0.04 0.05 0.05 0.08 0.03 MgO 0.60 0.65 0.37 0.58 0.38 0.38 0.53 0.43 0.29 0.27 CaO 1.91 1.89 1.50 1.81 1.51 1.63 1.65 1.46 1.43 1.13

Na2O 3.18 3.16 3.14 3.16 3.09 3.12 3.05 3.15 3.46 3.00 K2O 4.20 4.28 5.03 4.34 4.47 4.35 4.47 4.45 3.67 4.74 P2O5 0.10 0.15 0.15 0.14 0.15 0.16 0.13 0.15 0.10 0.07 LOI 1.07 0.71 0.69 0.69 0.62 0.55 0.66 0.73 0.83 0.79 REST 0.28 0.28 0.29 0.31 0.25 0.25 0.33 0.28 0.23 0.23 Total 99.82 99.60 99.86 99.82 99.79 99.74 99.88 99.90 99.85 99.88

Trace Elements (ppm) Ba 622 516 716 548 470 535 496 475 151 394 Rb 170 185 213 197 204 193 204 216 259 186 Sr 142 129 140 118 123 140 112 115 76 105 Pb 22 24 26 26 20 20 24 23 15 20 Th 20 18 14 20 20 17 20 20 19 22 U 56469446146 Zr 229 194 147 214 153 144 198 153 129 165 Hf 6426535445 Nb 15 11 13 13 13 10 13 13 11 9 Y 35 36 34 41 36 38 36 37 33 34 La 15 23.00 19 22 13 18 35 19 15.40 48 Ce 48 51.70 26 41 31 28 32 32 35.00 44 Nd - 26.50------17.70- Sm -5.89------4.28- Eu -0.67------0.31- Tb -0.98------0.80- Ho -1.31------1.16- Yb -2.93------3.17- Lu -0.42------0.49- V 25 36 11 30 16 18 19 28 16 18 Cr 11 18 10 20 10 8 15 10 11 10 Ni 0000000000 Zn 40 46 34 44 39 37 39 39 42 30 Ga 16 15 16 14 15 16 15 14 15 14 F 1107 1184 1203 1418 1079 1005 1790 1362 1253 974 Cl 28 57 53 54 42 44 55 50 28 26

CIPW NORMS Q 31.47 31.49 30.44 31.92 33.63 33.74 33.34 33.52 35.87 35.86 C 1.11 1.22 1.91 1.3 2.16 2.03 1.47 1.92 2.13 1.35 Z 0.05 0.04 0.03 0.04 0.03 0.03 0.04 0.03 0.03 0.03 Or 25 25.53 29.93 25.85 26.61 25.91 26.61 26.48 21.87 28.18 Ab 27.01 26.88 26.64 26.83 26.24 26.5 25.88 26.72 29.37 25.45 An 8.33 7.87 5.94 7.34 5.91 6.64 6.31 5.55 5.68 4.64 Hy 4.57 4.8 3.22 4.6 3.53 3.39 4.28 3.89 3.1 2.7 Mt 0.35 0.33 0.29 0.3 0.33 0.28 0.3 0.17 0.37 0.25 Il 0.78 0.73 0.51 0.73 0.53 0.53 0.67 0.57 0.42 0.53 Ap 0.24 0.36 0.36 0.34 0.39 0.39 0.31 0.36 0.24 0.17 Mg# 33.84 35.55 29.98 33.09 28.37 29.48 32.52 29.86 25.37 25.58

Gwydir R. = Gwydir River 298

Whole-rock analyses - Hillgrove & Bundarra Suite enclaves - Chapter 4.

Sample E27X D18X A186X A183X CB2X BI2X BI6X BI9X BI9X1 BI1X Grid ref 048353 656549 834147 818145 991736 975911 014889 953803 953803 963924 Pluton Rockvale Dund. Gara Gara Gwydir R. Gwydir R. Gwydir R. Gwydir R. Gwydir R. Gwydir R. Suite ------Hillgrove ------Bundarra ------

SiO2 64.56 65.90 68.16 68.29 66.29 69.47 70.14 70.27 70.39 70.54 TiO2 0.72 0.84 0.64 0.67 1.01 0.76 0.67 0.66 0.58 0.63 Al2O3 15.62 14.75 14.45 14.52 14.51 14.08 14.00 14.07 14.19 13.91 Fe2O3 0.72 0.58 0.66 0.44 0.60 0.42 0.39 0.61 0.36 0.40 FeO 5.13 5.33 3.44 3.45 5.51 3.83 3.58 3.24 3.24 3.49 MnO 0.15 0.17 0.09 0.09 0.11 0.09 0.08 0.08 0.07 0.09 MgO 2.50 1.95 1.66 1.52 1.51 1.39 1.23 1.05 0.98 1.10 CaO 3.66 2.77 2.67 2.70 2.81 2.78 2.59 2.24 2.20 2.42 Na2O 3.09 3.48 3.13 3.22 3.38 3.73 3.56 3.37 3.57 3.59 K2O 2.17 2.38 3.47 3.22 2.18 1.83 2.26 2.51 2.69 2.34 P2O5 0.17 0.15 0.14 0.14 0.27 0.20 0.18 0.18 0.16 0.18 LOI 1.06 1.16 0.99 1.24 1.12 0.98 0.86 1.04 1.00 0.81 REST 0.27 0.45 0.34 0.33 0.56 0.37 0.36 0.44 0.42 0.33 Total 99.81 99.89 99.86 99.83 99.83 99.91 99.90 99.76 99.84 99.84

Trace Elements (ppm) Ba 396 436 651 674 157 187 317 308 443 189 Rb 101 151 139 143 312 174 195 240 241 213 Sr 216 196 151 154 89 102 91 125 115 83 Pb 13 18 20 13 9 11 20 15 14 14 Th 14 16 16 16 19 21 23 24 23 19 U 0153350006 Zr 187 197 202 217 253 439 381 355 340 362 Hf 654771097108 Nb 11 13 14 14 16 20 13 15 14 16 Y 30 39 31 35 41 54 44 46 42 40 La 9202149223333424059 Ce 45 38 40 32 37 63 54 56 50 53 V 98 99 68 53 62 45 38 43 36 40 Cr 74 36 43 30 21 40 33 24 22 31 Ni 19841010000 Zn 98 96 71 66 120 79 72 83 72 73 Ga 18 19 18 15 21 17 18 18 17 15 F 880 2596 1557 1373 3922 1920 1868 2549 2283 1698 Cl 133 157 63 77 171 108 108 93 96 117

CIPW NORMS Q 23.9 25.27 27.15 27.74 28.4 31.4 32.13 34.07 32.19 32.63 C 2.1 2.4 1.35 1.44 3.11 1.84 1.85 2.78 2.33 1.9 Z 0.04 0.04 0.04 0.04 0.05 0.09 0.08 0.07 0.07 0.07 Or 12.92 14.2 20.66 19.18 13.09 10.93 13.49 15.03 16.08 13.98 Ab 26.16 29.49 26.57 27.33 28.67 31.62 30.18 28.64 30.31 30.44 An 16.77 11.21 11.57 11.86 9.64 11.35 10.59 8.38 8.51 9.84 Hy 14.18 13.15 9.05 8.85 11.99 9.09 8.39 7.16 7.3 7.97 Mt 1.07 0.87 0.98 0.65 0.89 0.62 0.58 0.9 0.53 0.59 Il 1.37 1.6 1.22 1.28 1.93 1.45 1.28 1.26 1.11 1.2 Ap 0.41 0.36 0.34 0.34 0.65 0.48 0.43 0.43 0.38 0.43 Mg# 46.48 39.47 46.23 43.98 32.81 39.27 37.97 36.61 35.02 35.97

Dund. = Dundurrabin, Gwydir R. = Gwydir River 299

Whole-rock analyses - Bakers Creek Suite & accretion complex metabasalts - Chapter 4.

Sample G39 CC26A CCD CCD2 A279 D7 A270 A156 UN3-3 UN5-1 Lithology ---- Gabbro ------Diorite ------Metabasalt --- Grid ref 994603 479667 534613 536603 977149 757371 676877 912905 826205 872505 Pluton Days C. SSCC CCD CCD Bakers C. DMC WC CC Tia C. Tia C. SiO2 49.48 50.16 55.31 55.32 58.57 58.92 61.31 62.04 44.85 47.44 TiO2 0.31 0.92 1.69 1.62 0.80 0.83 0.88 0.99 4.59 2.78 Al2O3 20.82 17.01 14.75 14.74 15.61 15.27 14.76 15.18 13.70 12.69 Fe2O3 0.67 1.46 1.30 1.18 0.70 1.15 1.00 0.69 3.24 2.50 FeO 4.33 5.80 7.99 8.03 6.08 4.51 5.38 5.47 10.67 10.68 MnO 0.10 0.14 0.20 0.21 0.15 0.11 0.14 0.12 0.31 0.20 MgO 7.24 8.51 4.65 4.73 5.26 3.42 3.96 3.23 7.33 5.84 CaO 13.49 12.36 6.96 7.67 6.95 7.45 5.94 5.01 6.53 9.86 Na2O 2.29 2.72 3.65 3.35 2.67 3.65 2.77 3.35 3.76 3.37 K2O 0.16 0.19 1.49 1.16 1.87 1.67 2.06 1.50 0.57 0.39 P2O5 0.04 0.10 0.22 0.21 0.14 0.13 0.15 0.16 0.40 0.31 LOI 1.89 1.34 1.70 1.23 0.84 2.46 1.20 1.79 3.76 3.62 REST 0.22 0.20 0.27 0.36 0.27 0.30 0.31 0.36 0.33 0.37 Total 101.05 100.92 100.18 99.81 99.89 99.86 99.87 99.88 100.03 100.05

Trace Elements (ppm) Ba 36 57 223 213 391 330 361 483 296 157 Rb 6 4 66 50 70 54 77 43 12 6 Sr 172 163 178 244 155 98 162 240 302 265 Pb 14613112016138 5 1 Th 14661012914 85 U 00001323 10 Zr 34 80 201 207 136 183 154 187 226 212 Hf 11363325 87 Nb 4 5 12 11 12 12 12 9 31 25 Y 11 22 45 45 30 30 32 34 40 38 La 50072941615 224 Ce 11 15 23 21 26 25 29 45 22 29 V 119 195 212 211 149 124 119 113 351 318 Cr 150 537 70 104 175 115 110 78 0 46 Ni 46 96 21 23 31 20 14 15 1 35 Zn 47 51 100 95 80 66 74 59 110 113 Ga 17 17 19 20 17 18 16 19 22 19 F 1252 109 934 1760 949 1488 1578 1539 1384 1922 Cl 73 121 147 138 10 77 40 309 9 3

CIPW NORMS Q - - 5.16 6.8 11.17 11.46 17.26 18.83 - - Or 0.94 1.12 8.84 6.91 11.11 9.93 12.26 8.92 3.38 2.31 Ab 19.16 22.75 30.8 28.4 22.66 30.96 23.51 28.25 31.9 28.6 An 45.71 33.43 19.54 21.94 25.19 20.47 21.88 22.21 18.9 18.41 Di 15.75 21.57 10.88 11.52 6.59 12.21 4.79 0.75 8.45 22.74 Hy 8.1 7.69 17.37 17.66 19.46 8.74 15.42 15.79 5.81 8.07 Ol 6.637.94------13.296.34 Mt 0.99 2.14 1.93 1.77 1.05 1.7 1.48 1.03 4.79 3.71 Il 0.58 1.73 3.21 3.09 1.52 1.58 1.68 1.89 8.74 5.3 Ap 0.09 0.24 0.52 0.5 0.33 0.31 0.36 0.38 0.95 0.74 Mg# 74.87 72.33 50.91 51.21 60.66 57.47 56.74 51.27 55.04 49.35

Days C. = Days Creek Gabbro, SSCC = Sheep Station creek Complex, CCD = Charon Creek Diorite Bakers C. = Bakers Creek Diorite, DMC = Dorrigo Mountain Complex, WC = Woodburn Complex CC = Cheyenne Complex, Tia C. = Tia Metamorphic Complex. 300

Whole-rock analyses - Chaelundi Complex I-type suite - Chapter 5.

(Note that these samples were collected as part of an honours project and presented in Landenberger 1988. However, the samples have been re-analysed with the low-dilution fusion technique.)

Sample CC106 CC12 CC114 CC7 CC8 CC100 CC18 CC10 CC119 CC9 CC103 CC101 CC20 Grid Ref. 436678 424707 409692 437666 456680 364718 393714 449696 447700 452692 433665 383686 361698

SiO2 67.56 67.61 67.66 67.66 67.80 68.17 68.39 68.47 68.58 68.90 70.36 70.84 71.15 TiO2 0.56 0.56 0.56 0.57 0.54 0.54 0.53 0.50 0.50 0.44 0.42 0.39 0.40 Al2O3 15.10 15.10 15.14 15.28 14.97 15.05 15.01 15.13 14.95 14.64 14.42 14.51 14.40 Fe2O3 0.95 0.99 1.15 0.84 1.09 0.98 0.98 1.09 0.95 0.82 0.64 0.65 0.49 FeO 2.40 2.39 2.26 2.49 2.25 2.26 2.26 2.00 2.14 1.86 2.03 1.76 1.91 MnO 0.07 0.07 0.07 0.06 0.07 0.07 0.07 0.07 0.07 0.05 0.06 0.05 0.05 MgO 1.43 1.43 1.43 1.43 1.50 1.31 1.33 1.28 1.31 1.14 0.87 0.84 0.81 CaO 3.31 3.30 3.24 3.17 2.85 3.14 3.19 3.00 2.94 2.67 2.29 2.46 2.23

Na2O 3.85 3.90 3.86 4.01 3.91 3.90 3.82 3.90 3.78 3.86 3.81 3.79 3.89 K2O 3.13 3.08 3.17 3.07 3.31 3.24 3.21 3.25 3.31 3.33 3.74 3.53 3.56 P2O5 0.17 0.17 0.17 0.16 0.16 0.16 0.16 0.18 0.16 0.13 0.12 0.12 0.10 LOI 0.90 0.79 0.82 0.65 1.07 0.66 0.56 0.69 0.91 1.91 0.71 0.49 0.59 Rest 0.24 0.23 0.22 0.25 0.23 0.23 0.22 0.22 0.22 0.30 0.27 0.23 0.23 Total 99.67 99.61 99.76 99.62 99.75 99.71 99.73 99.77 99.82 100.06 99.73 99.66 99.82 Trace elements (ppm)

Ba 496 526 476 482 513 515 469 536 495 472 524 470 509 Rb 99 98 97 90 97 103 104 119 101 129 102 117 116 Sr 326 329 327 334 335 304 314 307 288 269 232 258 248 Pb 1016159181516181511161918 Th 12 12 10 6 14 15 14 9 13 12 15 13 11 U 34021<15223521 Zr 188 187 182 188 186 197 180 170 165 159 199 170 128 Hf 435555554<2554 Nb 10 15 13 12 11 11 14 11 11 9 15 12 14 Y 32 31 30 26 28 31 27 21 28 24 31 29 23 La 16 18 20 19 17 16 20 19 20 20 24 19 29 Ce 35 41 42 42 32 34 38 40 34 31 40 40 41 Nd ---20---21---21- Sm - - - 4.48 - - - 3.98 - - - 4.27 - Eu - - - 1.11 - - - 0.91 - - - 0.86 - Tb - - - 0.81 - - - 0.74 - - - 0.67 - Ho - - - 1.24 - - - 0.93 - - - 0.97 - Yb - - - 3.27 - - - 2.23 - - - 2.56 - Lu - - - 0.48 - - - 0.31 - - - 0.38 - V 61 63 60 63 58 63 64 53 57 48 45 37 36 Cr 18 14 18 29 16 12 13 17 22 15 8 11 10 Ni <2 <2 <2 <2 <2 <2 <2 <2 <2 <2 <2 <2 <2 Cu 361089611861932479776 Zn 38 42 50 36 33 44 42 44 47 33 37 39 44 Ga 18 14 15 14 15 16 15 16 14 14 14 14 16 F 496 454 463 511 434 498 484 370 442 - 528 381 499 Cl 196 116 105 204 175 105 102 183 138 930 127 121 74 Modes Quartz 24.6 27.5 25.5 27.2 26.3 23.7 23.2 26.6 22.4 24.3 23.4 25.3 30.0 Alkali feldspar16.6 13.8 19.0 13.2 15.9 18.4 15.0 22.9 16.6 23.1 24.8 24.5 23.2 Plagioclase 44.7 46.4 39.7 46.5 46.8 47.8 50.2 39.2 50.6 44.3 43.6 42.6 37.0 Biotite 7.7 6.5 9.1 5.5 6.5 6.8 5.3 7.3 4.9 6.7 6.7 6.2 7.6 Hornblende 5.2 4.7 5.5 5.6 3.1 2.7 5.7 3.2 4.4 1.1 0.9 1.0 1.3 Opaques 1.1 0.7 0.9 1.4 1.2 0.5 0.3 0.4 0.9 0.3 0.3 0.2 0.4

* Grid references are Australian Map Grid (AMG) and refer to the Ebor 1:100000 topographic sheet (sheet m 9337) 301

Whole-rock analyses - Chaelundi Complex A-type suite - Chapter 5.

(Note that those samples marked * were collected as part of an honours project and presented in Landenberger (1988). All other analyses are new data. All samples have been re-analysed with the low-dilution fusion technique.)

Sample *CC66 CC131 CC132 CC133 CC134 CC135 *CC64 CC128 *CC57 CC129 CC136 CC130 Grid Ref.*393747 393774 393776 392769 394758 396754 396759 383734 378745 382752 397747 375800

SiO2 66.46 66.75 66.84 70.55 70.75 71.50 74.23 75.63 75.69 75.71 76.85 76.89 TiO2 0.54 0.62 0.53 0.38 0.38 0.34 0.18 0.11 0.10 0.12 0.05 0.05 Al2O3 16.18 15.63 16.23 14.80 14.78 14.67 13.36 13.01 13.00 13.14 12.91 12.83 Fe2O3 1.09 0.53 0.73 0.35 0.35 0.43 0.28 0.34 0.34 0.22 0.38 0.16 FeO 2.58 3.19 2.82 2.23 2.17 1.73 1.33 0.81 0.75 0.76 0.37 0.19 MnO 0.09 0.08 0.08 0.05 0.06 0.05 0.05 0.04 0.04 0.04 0.03 <0.01 MgO 0.79 1.05 0.80 0.53 0.51 0.48 0.14 0.11 0.11 0.11 <0.01 0.01 CaO 2.37 2.70 2.56 1.71 1.68 1.63 0.77 0.74 0.64 0.73 0.17 0.37

Na2O 4.56 4.26 4.69 4.01 4.01 4.01 3.53 4.02 4.10 4.16 3.70 3.88 K2O 3.82 3.63 3.86 4.38 4.36 4.47 4.92 4.34 4.45 4.05 4.60 4.40 P2O5 0.17 0.19 0.15 0.10 0.10 0.08 0.04 0.03 0.03 0.04 0.01 0.01 LOI 0.74 1.00 0.68 0.25 0.33 0.42 0.64 0.35 0.37 0.44 0.57 0.66 Rest 0.32 0.31 0.33 0.29 0.32 0.26 0.24 0.27 0.26 0.34 0.18 0.40 Total 99.69 99.92 100.30 99.63 99.79 100.07 99.71 99.79 99.88 99.86 99.82 99.86 Trace elements (ppm) Ba 1079 908 1115 889 887 781 708 174 138 194 77 92 Rb 77 93 70 119 118 126 154 290 307 308 458 292 Sr 277 281 282 177 181 167 69 59 46 51 4 14 Pb 23 22 11 21 24 20 28 42 44 45 58 71 Th 9 11 6 19 14 13 21 32 31 24 46 37 U 4453522659518 Zr 443 429 444 335 312 255 207 139 144 135 127 119 Hf 77<267<2456474 Nb 16161010126122929264431 Y 29 33 26 33 36 31 33 76 80 118 44 342 La 49 30 37 31 30 31 34 16 14 20 42 41 Ce 40 62 78 47 66 54 74 43 42 53 47 127 Nd -3237-33-34262534-105 Sm - 6.33 6.34 - 6.16 - 6.35 7.78 7.82 11.00 - 41.90 Eu - 1.52 1.67 - 1.18 - 0.36 0.26 0.18 0.33 - 0.14 Tb - 0.86 0.77 - 0.84 - 0.70 1.82 1.78 2.63 - 8.24 Ho - 1.22 0.98 - 1.16 - 0.95 2.82 2.84 3.32 - 9.03 Yb - 2.94 2.63 - 3.01 - 2.83 8.71 8.87 9.73 - 17.30 Lu - 0.42 0.41 - 0.46 - 0.45 1.39 1.40 1.55 - 2.67 V 42 58 41 44 38 25 11 10 15 13 8 5 Cr 8152794742745 Ni <2<22<2<2<2<2<2<2<2<2<2 Cu 91011261335<2<2111 Zn 66 62 68 38 59 50 27 40 38 52 29 73 Ga 17 18 17 18 17 17 14 18 20 18 22 20 F 429 510 458 648 877 576 686 1500 1407 2115 593 2444 Cl 113 117 140 62 138 137 68 32 34 33 25 13 Modes Quartz 18.9 18.6 15.8 28.5 24.2 29.4 34.8 40.2 36.6 43.8 40.6 - K-feldspar 34.0 29.0 24.3 33.2 28.8 34.8 35.1 33.1 30.7 22.6 33.0 - Plagioclase 34.8 39.0 44.2 28.5 37.6 27.4 24.6 19.8 28.6 28.2 23.6 - Biotite 8.6 7.3 6.2 6.9 5.6 5.5 4.9 6.0 3.9 4.9 2.6 - Hornblende 2.34.06.21.01.91.7---- - Opaques 0.5 0.7 1.6 1.5 1.4 0.8 0.2 0.7 0.2 0.5 - -

* Grid references are Australian Map Grid (AMG) and refer to the Ebor 1:100000 topographic sheet (sheet m 9337) 302

Whole-rock analyses - Woodlands Quartz Monzonite & Enclaves - Chapter 6

Sample WQM2 WQM WQM3 ORLAWQMX2WQMX6WQMX7WQMX3WQMX10WQMX8WQMX5WQMX11

SiO2 64.94 69.03 69.00 74.47 52.94 52.99 53.51 55.15 57.33 58.99 59.23 77.32 TiO2 0.55 0.30 0.33 0.18 1.23 1.24 1.17 1.18 1.31 1.38 1.36 0.06 Al2O3 16.06 15.11 15.08 13.11 18.27 18.24 18.15 17.89 17.28 15.43 15.47 12.17 Fe2O3 1.32 1.06 0.95 0.29 1.42 1.23 1.29 1.34 1.36 1.15 1.21 0.72 FeO 3.70 2.12 2.21 1.15 7.05 7.20 6.84 6.52 6.51 5.92 5.84 0.43 MnO 0.15 0.11 0.10 0.04 0.22 0.22 0.24 0.23 0.28 0.27 0.26 0.01 MgO 0.83 0.34 0.44 0.34 3.66 3.60 3.32 3.01 2.21 2.53 2.40 0.00 CaO 3.27 1.63 1.81 1.29 8.27 8.34 7.79 7.5 5.76 4.74 4.61 0.38

Na2O 4.44 4.55 4.50 3.08 3.58 3.56 3.62 3.85 4.44 4.71 4.63 3.75 K2O 3.22 4.37 4.30 4.92 1.51 1.48 1.79 1.85 1.81 2.87 2.92 4.72 P2O5 0.15 0.06 0.07 0.06 0.41 0.41 0.38 0.4 0.49 0.31 0.31 0.01 LOI 0.25 0.38 0.43 0.61 0.57 0.27 0.27 0.31 0.32 0.20 0.32 0.16 REST 0.46 0.46 0.33 0.22 0.50 0.49 0.57 0.55 0.50 0.58 0.50 0.08 Total 99.73 99.77 99.79 99.87 100.42 100.06 99.69 100.51 100.32 99.75 99.69 99.85 Trace elements (ppm) Ba 1344 1175 1106 296 658 624 679 663 828 903 888 108 Rb 93 107 103 251 85 78 121 116 126 126 143 96 Sr 342 142 156 117 580 584 543 534 671 262 257 5 Pb 21181829167131614201724 Th 1419184168865181620 U 4311020103004 Zr 614 525 491 134 229 215 225 326 365 296 308 229 Hf 111110597513887 Nb 13910151011111210161612 Y 30 36 34 27 23 23 25 22 31 40 35 36 La 7281.7822020.323373029.4414919 Ce 163 165 189 43 47.3 45 52 73 63.4 90 97 46 Nd - 64.1 - - 21.5 - - - 36.7 48.9 - - Sm - 9.42 - - 4.98 - - - 8.11 8.84 - - Eu - 2.05 - - 1.7 - - - 2.92 1.89 - - Tb - 1.17 - - 0.72 - - - 1.18 1.17 - - Ho - 1.7 - - 0.99 - - - 1.36 1.55 - - Yb - 4.9 - - 2.16 - - - 2.96 3.34 - - Lu - 0.74 - - 0.31 - - - 0.45 0.58 - - V 46 16 27 15 188 185 184 160 47 149 140 1 Cr 6 8 22 11 21 21 25 18 23 30 27 5 Ni 000011065000 Cu 00000002114000 Zn 84 75 74 26 158 147 183 175 209 262 248 14 Ga 15 16 16 11 21 19 17 18 20 20 18 16 Cl 347 195 213 64 660 710 642 744 473 432 463 21 F 901 750 374 863 735 719 1582 1175 1025 1767 1726 -

Hornblende 7.2 3.0 4.4 0 Biotite 9.7 5.8 5.4 5.1 Alkali feldspar 15.2 26.1 25.7 30.1 Quartz 19.7 22.7 22.6 35.8 Plag 48.0 42.0 41.7 28.8 Opaques 0.3 0.3 0.1 trace 303

Whole-rock analyses - Smokey Cape Adamellite and enclaves - Chapter 6.

Sample SCA1 SCA2 SCA4 SCA5 SC5X3 SC3X2 SC5X4 SCXX3 SC5X1 SCA1X SC5X2 SC4X2 SC4X3 SCXX2 SCAX1 Lithology ------Adamellite ------Enclave ------

SiO2 66.99 68.84 69.64 73.76 57.45 58.95 58.98 59.29 59.85 59.87 60.01 60.43 60.53 60.82 62.57 TiO2 0.49 0.36 0.30 0.23 1.03 1.05 1.04 0.98 0.94 1.05 0.95 0.95 0.94 0.92 0.77 Al2O3 16.16 15.83 15.17 13.70 15.99 16.34 16.38 16.37 16.18 16.44 16.12 16.22 16.15 16.16 16.14 Fe2O3 0.76 0.37 0.48 0.29 1.10 1.09 0.80 0.62 0.82 0.78 0.91 0.86 0.91 0.68 0.71 FeO 2.66 2.23 2.04 1.63 7.20 5.91 5.59 5.37 5.63 5.45 5.41 5.17 5.59 5.20 4.44 MnO 0.08 0.05 0.05 0.04 0.19 0.14 0.14 0.12 0.15 0.16 0.13 0.12 0.14 0.11 0.10 MgO 0.64 0.38 0.26 0.17 3.06 3.31 3.13 3.14 2.71 3.06 2.87 2.89 2.81 2.85 2.09 CaO 2.32 1.65 1.14 0.80 4.30 4.67 4.98 5.02 4.32 5.12 4.49 4.24 4.01 4.45 3.48

Na2O 3.65 4.11 3.44 3.87 4.21 4.56 3.89 4.38 4.28 4.32 4.63 4.28 4.42 3.79 4.72 K2O 3.90 4.44 4.58 4.55 1.86 1.45 2.52 2.50 2.53 1.79 2.46 2.05 2.23 2.61 2.52 P2O5 0.11 0.08 0.05 0.04 0.30 0.31 0.30 0.28 0.28 0.31 0.28 0.26 0.27 0.26 0.22 LOI 0.66 0.65 1.45 1.04 2.91 2.03 1.95 1.85 2.01 1.75 1.83 1.86 2.17 1.82 1.67 Rest 0.65 0.36 0.28 0.31 0.42 0.47 0.43 0.46 0.45 0.23 0.47 0.45 0.38 0.45 0.36 Total 99.07 99.35 98.88 100.43 100.02 100.28 100.13 100.38 100.15 100.33 100.56 99.78 100.55 100.12 99.79 Trace elements (ppm) Ba 1217 1140 1012 647 382 378 348 354 360 325 392 443 474 333 665 Rb 64 74 74 81 85 59 78 84 89 72 99 67 66 87 69 Sr 223 153 118 82 284 358 376 372 321 366 328 320 326 334 283 Pb 47 14 16 15 16 10 14 12 14 18 15 10 11 14 14 Th 9 9 10 11 9 8 9 8 12 8 10 10 12 9 11 U 123221332434231 Zr 411 337 321 220 245 229 237 243 238 257 245 255 240 240 282 Hf 1198564754656677 Nb 13 9 10 7 12 15 12 14 10 10 12 13 11 9 12 Y 31 31 33 27 50 40 42 39 42 38 43 41 49 40 41 La 31.740184543241720.540252036303340 Ce 62.194828670576044.772845855833570 Nd 30.8------24.1------Sm 6.07------6.45------Eu 1.99------1.5------Tb 0.82------1.18------Ho 0.96------1.4------Yb 2.74------3.47------Lu 0.48------0.52------V 47 32 24 16 140 144 145 135 126 138 122 128 133 129 97 Cr 44432224202017201922211715 Ni 52514854564945484244454031 Cu 4433282130312642618182712 Zn 85 49 61 40 208 196 170 142 155 154 152 174 190 133 142 Ga 21 21 20 18 22 23 21 23 21 23 22 20 21 21 22 S 3341 171 25 130 140 150 177 195 136 89 143 132 132 164 94 F 404 847 610 1320 1457 2632 2155 2413 2110 261 2593 2285 1519 2507 1426 Cl 24 165 34 131 681 69 135 114 409 100 142 59 72 100 41 n.b. Trace elements (except REE and fluorine) for all samples in this table were analysed from pressed-powder pellets and not from LDF disks. 304

Appendix F

Isotopic data 305

Isotope results for Hillgrove Supersuite, Bakers Creek Suite and metasediments - Chapter 4.

87 86 87 86 87 86 147 144 143 144 143 144 Sample Suite Rb Sr Rb/ Sr Sr/ Sr Sr/ Sri Sm Nd Sm/ Nd Nd/ Nd Nd/ Nd gNd N42 Rockisle 109 242 1.3 0.709340 0.70379 3.45 16.3 0.127875 0.512705 0.512454 3.94 D15 Hillgrove 141 159 2.57 0.717034 0.70606 6.54 30.2 0.13083 0.512581 0.512324 1.41 A7 Hillgrove 137 210 1.89 0.722198 0.70490 4.84 23 0.127098 0.512567 0.512317 1.28 A60 Hillgrove 182 141 3.74 0.712945 0.70620 5.69 27.8 0.123662 0.512511 0.512268 0.32 A180 Hillgrove 164 155 3.06 0.719232 0.70617 6.33 29.4 0.130287 0.512537 0.512281 0.57 E28 Hillgrove 162 162 2.9 0.719173 0.70679 6.8 32.6 0.126267 0.512529 0.512281 0.57 N46 Hillgrove 170 102 4.83 0.727107 0.70649 9.13 32.8 0.168592 0.512692 0.512361 2.13 E29 Hillgrove 145 183 2.29 0.716389 0.70661 6.31 29.8 0.127929 0.512483 0.512232 -0.39 Y62 Hillgrove 133 168 2.29 0.715260 0.70548 5.33 24.7 0.13079 0.512617 0.51236 2.11 W443 Sediment 193 92.7 6.04 0.742547 0.71676 7.15 36.4 0.11879 0.512083 0.51185 -7.85 A30 Sediment 56 426 0.38 0.706511 0.70489 4.77 22.2 0.129943 0.512589 0.512334 1.6 A4 Sediment 93.2 498 0.541 0.707217 0.70491 3.96 21.2 0.112951 0.51247 0.512248 -0.07 Y44 Sediment 20.8 909 0.063 0.704676 0.70441 3.42 16.3 0.127001 0.512795 0.512546 5.73 A138 Sediment 144 181 2.3 0.718430 0.70861 5.35 26.7 0.12124 0.512348 0.51211 -2.77 A18 Sediment 134 326 1.19 0.710470 0.70539 4.3 22.9 0.113485 0.51252 0.512297 0.88 CC26A Bakers C. 4.4 163 0.0781 0.703093 0.70276 2.23 6.6 0.205598 0.513145 0.512741 9.55

Initial ratios and gNd calculated at 300 Ma.

Isotope results for Woodlands Quartz Monzonite & enclaves - Chapter 6.

87 86 87 86 87 86 147 144 143 144 143 144 Sample Suite Rb Sr Rb/ Sr Sr/ Sr Sr/ Sri Sm Nd Sm/ Nd Nd/ Nd Nd/ Nd gNd WQM Woodlands 107 142 2.18 0.713818 0.70610 9.18 63.2 0.088345 0.512475 0.512331 0.26 WQMBbiotite sep. 164 11.2 43.08 0.863431 0.70610 biotite - whole rock age = 249Ma WQMX2 enclave 85 580 0.424 0.707265 0.70576 4.77 23.6 0.122229 0.512546 0.512347 0.57 WQMX5 enclave 138 249 1.6 0.710896 0.70523 8.39 43.7 0.11625 0.512569 0.51238 1.21 WQMX2Penclave p. 129 111 3.36 0.715723 0.70382 4.2 18.4 0.137819 0.512628 0.512403 1.69

Initial ratios and gNd calculated at 249 Ma. enclave p. = enclave pyroxene (augite) separate 306

Appendix G

Catalogued sample list with grid references. 307

Sample 1:100000 Catalogue Number Lithology Suite GridRef Pluton Sheet Name number A12 Argillite Sediment 746857 N/A Armidale 14939 A138 Pelite Sediment 834739 N/A Armidale 14940 A156 Diorite Bakers Creek 912905 Cheyenne Complex Armidale 14941 A161 Greywacke Sediment 986074 N/A Armidale 14942 A18 Greywacke Sediment 718887 N/A Armidale 14943 A180 Granite Hillgrove 868164 Gara Armidale 14944 A181 Greywacke Sediment 797156 N/A Armidale 14945 A183 Granite Hillgrove 818145 Gara Armidale 14946 A183X Enclave Hillgrove 818145 Gara Armidale 14947 A184 Granite Hillgrove 822145 Gara Armidale 14948 A185 Granite Hillgrove 846146 Gara Armidale 14949 A186X Enclave Hillgrove 834147 Gara Armidale 14950 A221 Granite Hillgrove 693069 Gostwyck Armidale 14951 A243 Granite Hillgrove 974713 Argyle Armidale 14952 A246 Granite Hillgrove 954189 Hillgrove Armidale 14953 A270 Diorite Bakers Creek 676877 Woodburn Armidale 14954 A273 Granite Hillgrove 815978 Enmore Armidale 14955 A274 Granite Hillgrove 823978 Enmore Armidale 14956 A275 Pelite Sediment 808007 N/A Armidale 14957 A276 Granite Hillgrove 870017 Blue Knobby Armidale 14958 A277 Granite Hillgrove 826144 Gara Armidale 14959 A278 Granite Hillgrove 942175 Hillgrove Armidale 14960 A279 Diorite Bakers Creek 977149 Baker’s Creek Armidale 14961 A281 Greywacke Sediment 806199 N/A Armidale 14962 A282 Granite Hillgrove 900211 Hillgrove Armidale 14963 A283 Granite Hillgrove 932205 Hillgrove Armidale 14964 A284 Granite Hillgrove 922182 Hillgrove Armidale 14965 A285 Granite Hillgrove 977184 Hillgrove Armidale 14966 A286 Granite Hillgrove 953208 Hillgrove Armidale 14967 A3 Greywacke Sediment 739800 N/A Armidale 14968 A30 Greywacke Sediment 654991 N/A Armidale 14969 A4 Greywacke Sediment 765806 N/A Armidale 14970 A49 Pelite Sediment 868987 N/A Armidale 14971 A56 Greywacke Sediment 833203 N/A Armidale 14972 A57 Greywacke Sediment 843201 N/A Armidale 14973 A60 Granite Rockisle 915213 Hillgrove Armidale 14974 A60B Granite Hillgrove 915213 Hillgrove Armidale 14975 A65 Granite Hillgrove 000172 Hillgrove Armidale 14976 A7 Granite Hillgrove 845874 Winterbourne Armidale 14977 A79 Greywacke Sediment 797710 N/A Armidale 14978 A87 Granite Hillgrove 983718 Argyle Armidale 14979 A96 Argillite Sediment 844200 N/A Armidale 14980 B11 Granite Bundarra 301888 Banalasta Bingara 14981 B2 Granite Bundarra 319903 Banalasta Bundarra 14982 B4 Granite Bundarra 190913 Banalasta Bundarra 14983 B9 Granite Bundarra 300123 Banalasta Bundarra 14984 BI1 Granite Bundarra 963924 Gwydir River Bingara 14985 BI1X Enclave Bundarra 963924 Gwydir River Bingara 14986 BI2 Granite Bundarra 975911 Gwydir River Bingara 14987 BI2X Enclave Bundarra 975911 Gwydir River Bingara 14988 BI3 Granite Bundarra 005884 Gwydir River Bingara 14989 BI6X Enclave Bundarra 014889 Gwydir River Bingara 14990 BI7 Granite Bundarra 986902 Gwydir River Bingara 14991 BI9 Granite Bundarra 953803 Gwydir River Bingara 14992 BI9X Enclave Bundarra 953803 Gwydir River Bingara 14993 BI9X1 Enclave Bundarra 953803 Gwydir River Bingara 14994 C.GR. Granite Hillgrove 513629 SSCC Dorrigo 13713 C120 Granite Hillgrove 063232 Abroi Carrai 14995 C22 Granite Hillgrove 048205 Abroi Carrai 14996 CB2 Granite Bundarra 991736 Gwydir River Cobbadah 14997 CB2X Enclave Bundarra 991736 Gwydir River Cobbadah 14998 CC10 Granodiorite Chaelundi I 449696 Chaelundi Complex Ebor 13681 CC100 Granodiorite Chaelundi I 364718 Chaelundi Complex Ebor 13701 CC101 Adamellite Chaelundi I 383686 Chaelundi Complex Ebor 13702 CC103 Adamellite Chaelundi I 433665 Chaelundi Complex Ebor 13703 CC106 Granodiorite Chaelundi I 436678 Chaelundi Complex Ebor 13706 CC114 Granodiorite Chaelundi I 409692 Chaelundi Complex Ebor 13707 308

Sample 1:100000 Catalogue Number Lithology Suite GridRef Pluton Sheet Name number CC119 Granodiorite Chaelundi I 447700 Chaelundi Complex Ebor 13709 CC12 Granodiorite Chaelundi I 424707 Chaelundi Complex Ebor 13682 CC128 Leucoadamellite Chaelundi A 383734 Chaelundi Complex Ebor 14999 CC129 Leucoadamellite Chaelundi A 382752 Chaelundi Complex Ebor 15000 CC130 Carapace Chaelundi A 375800 Chaelundi Complex Ebor 15001 CC131 Quartz Monzonite Chaelundi A 393774 Chaelundi Complex Ebor 15002 CC132 Quartz Monzonite Chaelundi A 393776 Chaelundi Complex Ebor 15003 CC133 Quartz Monzonite Chaelundi A 392769 Chaelundi Complex Ebor 15004 CC134 Adamellite Chaelundi A 394758 Chaelundi Complex Ebor 15005 CC135 Adamellite Chaelundi A 396754 Chaelundi Complex Ebor 15006 CC136 Microgranite Chaelundi A 397747 Chaelundi Complex Ebor 15007 CC18 Granodiorite Chaelundi I 393714 Chaelundi Complex Ebor 13685 CC20 Adamellite Chaelundi I 361698 Chaelundi Complex Ebor 13687 CC26A Gabbro Bakers Creek 479667 SSCC Dorrigo 13554 CC57 Leucoadamellite Chaelundi A 378745 Chaelundi Complex Ebor 13693 CC64 Adamellite Chaelundi A 396759 Chaelundi Complex Ebor 13695 CC66 Quartz Monzonite Chaelundi A 393747 Chaelundi Complex Ebor 13696 CC7 Granodiorite Chaelundi I 437666 Chaelundi Complex Ebor 13677 CC8 Granodiorite Chaelundi I 456680 Chaelundi Complex Ebor 13678 CC9 Granodiorite Chaelundi I 452692 Chaelundi Complex Ebor 13679 CCD Diorite Bakers Creek 534613 Charon Creek Diorite Dorrigo 13710 CCD2 Diorite Bakers Creek 536603 Charon Creek Diorite Dorrigo 15008 D11 Granite Hillgrove 653507 Dundurrabin Dorrigo 15009 D15 Granite Hillgrove 600543 Dundurrabin Dorrigo 15010 D18 Granite Hillgrove 656549 Dundurrabin Dorrigo 15011 D18X Enclave Hillgrove 656549 Dundurrabin Dorrigo 15012 D21 Granite Hillgrove 558598 Dundurrabin Dorrigo 15013 D7 Diorite Bakers Creek 757371 Dorrigo Mtn Dorrigo 15014 E27 Granite Hillgrove 048353 Rockvale Ebor 15015 E27X Enclave Hillgrove 048353 Rockvale Ebor 15016 E28 Granite Hillgrove 046357 Rockvale Ebor 15017 E29 Granite Hillgrove 039627 Tobermory Ebor 15018 E6 Granite Hillgrove 077782 Kookabookra Ebor 15019 G31 Granite Hillgrove 950563 Tobermory Guyra 15020 G37 Granite Hillgrove 036367 Rockvale Guyra 15021 G39 Gabbro Bakers Creek 994603 Day’s Creek Guyra 15022 GUM.A Granite Hillgrove 508641 SSCC Dorrigo 13712 N1 Granite Hillgrove 505547 Kilburnie Nundle 15023 N15 Granite Hillgrove 563153 Murder Dog Nundle 15024 N2 Granite Rockisle 511338 Rockisle Nundle 15025 N3 Granite Rockisle 517327 Rockisle Nundle 15026 N4 Granite Rockisle 483317 Rockisle Nundle 15027 N42 Granite Rockisle 486321 Rockisle Nundle 15028 N43 Granite Hillgrove 500549 Kilburnie Nundle 15029 N44 Granite Hillgrove 493546 Kilburnie Nundle 15030 N45 Granite Rockisle 474573 Kilburnie Nundle 15031 N46 Granite Rockisle 477578 Kilburnie Nundle 15032 NFG Granite Hillgrove 519622 SSCC Dorrigo 13711 ORLA Leucoadamellite Oban River 993773 Oban River Guyra 15033 SC3X2 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15034 SC4X2 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15035 SC4X3 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15036 SC5X1 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15037 SC5X2 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15038 SC5X3 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15039 SC5X4 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15040 SCA1 Quartz monzonite Smokey Cape 065795 Smokey Cape Nambucca Heads 15041 SCA1X Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15042 SCA2 Adamellite Smokey Cape 070845 Smokey Cape Nambucca Heads 15043 SCA4 Adamellite Smokey Cape 070845 Smokey Cape Nambucca Heads 15044 SCA5 Adamellite Smokey Cape 070845 Smokey Cape Nambucca Heads 15045 SCAX1 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15046 SCXX2 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15047 SCXX3 Enclave Smokey Cape 070845 Smokey Cape Nambucca Heads 15048 UN3-3 Metabasalt Tia Complex 826205 N/A Yarrowitch 15049 UN5-1 Metabasalt Tia Complex 872505 N/A Yarrowitch 15050 WQM Quartz Monzonite Woodlands 068796 Woodlands Ebor 15051 WQM2 Quartz Monzonite Woodlands 067797 Woodlands Ebor 15052 309

Sample 1:100000 Catalogue Number Lithology Suite GridRef Pluton Sheet Name number WQM3 Quartz Monzonite Woodlands 068797 Woodlands Ebor 15053 WQMX10 Enclave Woodlands 068796 Woodlands Ebor 15054 WQMX11 Enclave Woodlands 068797 Woodlands Ebor 15055 WQMX2 Enclave Woodlands 068796 Woodlands Ebor 15056 WQMX3 Enclave Woodlands 068796 Woodlands Ebor 15057 WQMX5 Enclave Woodlands 068796 Woodlands Ebor 15058 WQMX6 Enclave Woodlands 068796 Woodlands Ebor 15059 WQMX7 Enclave Woodlands 068796 Woodlands Ebor 15060 WQMX8 Enclave Woodlands 068796 Woodlands Ebor 15061 Y29 Greywacke Sediment 823622 N/A Yarrowitch 15062 Y44 Greywacke Sediment 937471 N/A Yarrowitch 15063 Y60 Psammopelite Sediment 735538 N/A Yarrowitch 15064 Y62 Granite Hillgrove 760456 Tia Yarrowitch 15065 Y94 Granite Rockisle 923684 Kimberley Park Yarrowitch 15066

1:100000 Map sheet reference Map Name Sheet # Armidale 9236 Bendemeer 9136 Bingara 9038 Carrai 9336 Cobbadah 9037 Dorrigo 9437 Ebor 9337 Guyra 9237 Glen Innes 9238 Manilla 9036 Nundle 9135 Newton Boyd 9338 Yarrowitch 9235