Plutonic petrogenesis and mineralisation in southwest

Plutonic petrogenesis and mineralisation in southwest Fiordland

Malcolm Gollan MSc. thesis, Department of Geology, University of Otago, Dunedin, New Zealand

1

I

II

Abstract Southwest Fiordland contains an important record of early Cretaceous plutonism, which is not complicated by a polyphase metamorphic overprint as in central and northern Fiordland. Recent mapping as part of the GNS Science Qmap Fiordland programme has identified multiple Paleozoic through Mesozoic plutons composing the previously undivided plutonic units in southwest Fiordland. Field, petrographic, geochemical, and geochronologic data reveal a transition in time and composition between early calc-alkaline Darran Suite and later ‘adakitic-like’ Separation Point Suite. The Revolver Pluton (Rp) of c.200 km2 is the largest unit. It is a calc-alkaline I-type, coarse grained biotite granite with distinctive pink alkali feldspar megacrysts. LA-ICP-MS zircon UPb dating of Rp from Revolver Bay yields an age of 132.4 ± 1.0 Ma. Granodiorite from the northeast sector of Rp, here named the Long Scarp Granodiorite (Lsg), gives an age of 133.1 ± 0.8 Ma and has characteristics requiring high fH2O during melting similar to the Separation Point Suite. Treble Mountain Granite (Tmg) is medium to coarse grained biotite granite that intrudes Rp on the west side of Isthmus Sound, and yields an age of 130.4 ± 0.9 Ma. Tmg locally exhibits severe hydrothermal alteration and hosts epithermal base-metal vein mineralisation which was worked in the historic Tarawera Mine. Trevaccoon Diorite (Tdi) is a hornblende gabbro that outcrops along the western shore of Long Sound north of Lady Bay and has an age of 128.7 ± 1.0 Ma. A younger granite, here named Upper Blacklock Granite (Ubg), occurs between Rp and Lsg. Ubg is petrographically similar to Rp, but can be distinguished on the basis of its geochemistry and age of 124.7 ± 1.0 Ma. Hornblende-rich HiSY diorite to gabbro, here named Only Island Diorite (Oid), outcrops south from Only Island where it intrudes Ubg. The composition and age of 122.1 ± 0.9 Ma of Oid suggest it is a higher level, unmetamorphosed equivalent of the Western Fiordland Orthogneiss. Nd isotopes (εNd = 0 - +1), oxygen isotopes (whole rock δ18O = 6-9, quartz δ18O = 8-10) for all units other than Tmg, are consistent with a isotopically depleted uniform mafic source, while a rarity of older inherited zircons, limit contamination by the surrounding Ordovician metasediments (δ18O whole rock = 12-18). The presence of 15% inherited 149-136 Ma zircons in all units except Oid suggests a shared contamination by a component comparable in age to Darran Suite rocks in eastern Fiordland. Tmg quartz and feldspar δ18O values are strongly depleted indicating that hydrothermal alteration was caused by circulation of meteoric water. This supports the interpretation that plutons in the region were intruded at shallow crustal levels (< 10 km). Three Paleozoic ages were obtained; one from Big Pluton (Bp, formally Kakapo Granite) which has to date not yielded a well resolved age. The sample located in the eastern portion of Bp yields an age of 367 ± 2 Ma. This sample is geochemically distinct from western portions of Bp where previous dating has been attempted, which is attributed to different melting conditions of a shared source. Two samples of the Tine Peak Tonalite (Tpt) yielded ages of 347 ± 3 Ma and 352 ± 2 Ma. Tpt shows extreme depletion in alkalis, while showing often extreme enrichment in HFS elements. Tpt most likely represents a previously unrecognised extremely high temperature A-type melt of a dominantly restitic metasedimentary source, which has previously yielded an S-type melt. Initial Nd (εNd = -4.6 and -1.9) and Sr (Sri = 0.707537 and 0.707136) isotopes and oxygen isotopes (whole rock δ18O = 12.3) are consistent with such a source.

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Long Sound (Narrow Bend) Revolver Hill

Revolver Bay

Kisbee Bay

Useless Bay

Isthmus Sound

Solitary Peak

Plate 1. (Top) View looking east from Treble Mountain. Most of the middle foreground is Revolver Pluton. The more distant mountains are composed of Big Pluton. (Bottom) Looking south across Lake Monk from near Rugged Mount. Right of Lake Monk and foreground is Tine Peak Tonalite. Monk Granite forms the south end of Lake Monk. Distant Hills are Big Pluton, previously named the Kakapo Granite. Solitary Peak is where Muir et al. (1998) dated a sample of the then Kakapo granite.

IV

Acknowledgements Firstly I would like to thank my supervisor Dr Mike Palin for always being available for discussion. In addition Mike endured more than most supervisors after a small boating incident during field work that left us stranded on an island the size of a house. This leads to special thanks to Doc Sutherland from Southwest Helicopters who came to our rescue. In addition Wayne and Sam from Southwest Helicopters showed much ability in getting us into and out of the field during what seemed to be more than our fare share of adverse weather conditions on the three field sessions. The project was encouraged by Dr Ian Turnbull from GNS Science, who provided much in the way of data, discussion, and support for which I am very grateful. Also special thanks to Department of Conservation, Invercargill for allowing the collection of rock samples. This thesis involved the collection of a lot of geochemical data for which I am indebted to the following people. Oxygen and hydrogen isotope analysis was done by Prof Chris Harris, University of Capetown, and Dr Kevin Faure of GNS Science. Acquisition of some REE data was obtained by Dr Candace Martin, University of Otago. Dr Charlotte Allen assisted greatly with use of the LA-ICP-MS at ANU for dating and geochemical analysis, while Dr Lorraine Paterson and Damien Walls at the University of Otago helped immensely with obtaining EMP and XRF data. Isotopic data for Nd and Sr was supplied by Dr Marc Norman of PRISE at the ANU. Also special thanks to Stephen Read, Department of Geology, for help with graphics. Financial assistance made life a lot more manageable for which I am very grateful to the University of Otago for a postgraduate scholarship. Funding was also made available for field expenses through scholarships from IGNS and the Society of Economic Geologists. The Geological Society of New Zealand also helped financially with attendance to two of its annual conferences.

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Table of contents 1

2

Introduction ....................................................................................................................... 1 1.1

Location .................................................................................................................................1

1.2

Aim of Research ....................................................................................................................2

1.3

Outline of Thesis ...................................................................................................................2

Background........................................................................................................................ 3 2.1 Regional Geology .........................................................................................................................3 2.1 2.1.1 2.1.2

2.2

Western Province Terranes .................................................................................................5 Northwest Nelson and Westland ..................................................................................................... 5 Fiordland.......................................................................................................................................... 8

New Zealand Intrusive Rocks ............................................................................................10

2.2.1 Median Batholith [incorporating the Median Tectonic Zone (MTZ)] ........................................... 11 2.2.2 Paleozoic Suites............................................................................................................................. 12 2.2.2.1 Karamea and Ridge suites.................................................................................................... 12 2.2.2.2 Paringa and Tobin suites ...................................................................................................... 13 2.2.2.3 Foulwind Suite ..................................................................................................................... 14 2.2.2.4 Correlations of Karamea Batholith granitoids with Australia and Antarctica...................... 14 2.2.3 Mesozoic Suites............................................................................................................................. 15 2.2.3.1 Darran (Median) Suite ......................................................................................................... 15 2.2.3.2 Separation Point Suite (SPS)................................................................................................ 18 2.2.3.3 Rahu Suite............................................................................................................................ 20 2.2.3.4 French Creek Suite (post orogenic alkaline magmatism)..................................................... 21 2.2.4 Model for the formation of HiSY magma ..................................................................................... 22

2.3 2.3.1 2.3.2 2.3.3 2.3.4

3

Southwest Fiordland...........................................................................................................23 Previous work................................................................................................................................ 23 Paleozoic metasediments ............................................................................................................... 24 Plutonic rocks ................................................................................................................................ 24 Cretaceous and Tertiary sediments ................................................................................................ 27

Lithological Units ............................................................................................................ 29 3.1

Preservation Formation (Fanny Bay Group) ...................................................................29

3.2

Intrusive Rocks ...................................................................................................................29

3.2.1 Paleozoic Intrusive Rocks.............................................................................................................. 30 3.2.1.1 Tine Peak Tonalite (tpt) ....................................................................................................... 30 3.2.1.2 Big Pluton (bp)..................................................................................................................... 33 3.2.2 Preservation Intrusive Complex (PIC)........................................................................................... 35 3.2.2.1 Revolver Pluton (rp)............................................................................................................. 35 3.2.2.2 Treble Mountain Granite (tmg) ............................................................................................ 37 3.2.2.3 Long Scarp Granodiorite (lsg) ............................................................................................. 39 3.2.2.4 Trevaccoon Diorite (tdi)....................................................................................................... 41 3.2.2.5 Upper Blacklock Granite (ubg)............................................................................................ 44 3.2.2.6 Only Island Diorite (oid)...................................................................................................... 44 3.2.2.7 Minor Intrusives and Dyke Rocks........................................................................................ 45

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Geochronology................................................................................................................. 47 4.1

ELA-ICP-MS methodology................................................................................................47

4.2

Paleozoic geochronology.....................................................................................................49

4.2.1 Tine Peak Tonalite......................................................................................................................... 49 4.2.1.1 Tine Peak Tonalite (OU 75184) ........................................................................................... 50 4.2.1.2 Tine Peak Tonalite (OU 75193) ........................................................................................... 52 4.2.2 Big Pluton (P 70791) ..................................................................................................................... 52 4.2.3 Discussion...................................................................................................................................... 53

4.3

Mesozoic Geochronology....................................................................................................55

VI 4.3.1 4.3.2 4.3.3 4.3.4 4.3.5 4.3.6

4.4

Long Scarp Granodiorite (OU 75175) ...........................................................................................55 Revolver Pluton (OU 75143) .........................................................................................................56 Treble Mountain Granite (OU 75108) ...........................................................................................58 Trevaccoon Diorite (OU 75154) ....................................................................................................59 Upper Blacklock Granite (OU 75170) ...........................................................................................61 Only Island Diorite (OU 75171) ....................................................................................................63

Discussion............................................................................................................................ 64

4.4.1 PIC Inheritance ..............................................................................................................................64 4.4.2 Zircon core study ...........................................................................................................................64 4.4.2.1 Upper Blacklock Granite......................................................................................................65 4.4.2.2 Treble Mountain Granite ......................................................................................................65 4.4.3 Zircon thermometry and survival of inherited zircon ....................................................................68 4.4.4 Summary........................................................................................................................................70

5

Igneous Geochemistry..................................................................................................... 75 5.1

Introduction........................................................................................................................ 75

5.2

Analytical techniques ......................................................................................................... 75

5.2.1 Mineral analysis (electron microprobe) .........................................................................................75 5.2.2 Major and trace element chemistry................................................................................................76 5.2.2.1 XRF ......................................................................................................................................76 5.2.2.2 LA-ICP-MS..........................................................................................................................76 5.2.3 Oxygen isotopes.............................................................................................................................77 5.2.3.1 Oxygen (laser assisted fluorination).....................................................................................77 5.2.3.2 Oxygen (conventional) .........................................................................................................77

5.3

Paleozoic geochemistry ...................................................................................................... 77

5.3.1 Previous work on NZ Paleozoic suites ..........................................................................................78 5.3.2 Geochemistry of Paleozoic granitoids in southwest Fiordland ......................................................81 5.3.2.1 Major elements.....................................................................................................................81 5.3.2.2 Trace elements......................................................................................................................85 5.3.2.3 Zircon saturation temperatures (TZr).....................................................................................88 5.3.2.4 Zircon REE chemistry ..........................................................................................................88 5.3.2.5 Oxygen isotopes ...................................................................................................................90 5.3.2.6 Neodymium and strontium isotopes.....................................................................................90 5.3.3 Discussion and Petrogenesis ..........................................................................................................91

5.4

Mesozoic Geochemistry ..................................................................................................... 95

5.4.1 Previous work on NZ Mesozoic suites ..........................................................................................95 5.4.2 PIC geochemistry...........................................................................................................................97 5.4.2.1 Major elements.....................................................................................................................97 5.4.2.2 Trace elements....................................................................................................................100 5.4.3 Compositional variation within PIC granites...............................................................................104 5.4.4 Oxidation state of PIC granites ....................................................................................................106 5.4.5 HFSE decoupling in Revolver Pluton..........................................................................................107 5.4.6 Oxygen isotopes...........................................................................................................................108 5.4.6.1 Long Scarp Granodiorite ....................................................................................................108 5.4.6.2 Revolver Pluton..................................................................................................................109 5.4.6.3 Treble Mountain Granite ....................................................................................................109 5.4.6.4 Upper Blacklock Granite....................................................................................................109 5.4.6.5 Trevaccoon Diorite.............................................................................................................109 5.4.6.6 Only Island Diorite.............................................................................................................110 5.4.7 Neodymium isotopes ...................................................................................................................110 5.4.8 Emplacement Pressure .................................................................................................................111 5.4.8.1 Qualitative constraints........................................................................................................111 5.4.8.2 PIC analyses in the H2O-saturated Ab-Or-Qtz system – potential constraints ...................111 5.4.8.3 Phengite geobarometry.......................................................................................................112 5.4.9 Discussion....................................................................................................................................114 5.4.9.1 Granite major and trace chemistry .....................................................................................114 5.4.9.2 Mafic rock major and trace chemistry................................................................................117 5.4.9.3 Oxygen isotopes .................................................................................................................118 5.4.9.4 Neodymium isotopes..........................................................................................................121 5.4.10 Petrogenesis ............................................................................................................................121

VII 5.4.10.1 5.4.10.2 5.4.10.3

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Tarawera Mine .............................................................................................................. 126 6.1

Analytical techniques........................................................................................................126

6.1.1 6.1.2

Mineral analysis (X-ray diffraction) ............................................................................................ 126 Hydrogen and carbon stable isotopic analysis ............................................................................. 126

6.2

History ...............................................................................................................................127

6.3

Site description..................................................................................................................127

6.4

Vein mineralisation...........................................................................................................129

6.5

Sulphide mineral textures and chemistry .......................................................................130

6.6

Timing of mineralisation ..................................................................................................132

6.7

Wall rock alteration..........................................................................................................132

6.8

Oxygen and hydrogen isotopes ........................................................................................134

6.9

Carbon isotopes.................................................................................................................134

6.10

Fluid inclusions and geothermobarometry.....................................................................135

6.11

Sulphide compositions and pressure constraints ...........................................................136

6.12

Discussion ..........................................................................................................................137

6.12.1 6.12.2 6.12.3

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Current petrogenetic models for granites ........................................................................... 121 PIC petrogenesis ................................................................................................................ 122 Potential sources ................................................................................................................ 124

Origin of the Tarawera mineralisation .................................................................................... 137 Potential linkages in alteration of host Tmg and Tarawera mineralisation ............................. 138 A genetic classification ........................................................................................................... 139

Structure......................................................................................................................... 142 7.1

Metasediment structure ...................................................................................................142

7.2

Intrusive rock structure ...................................................................................................142

7.3

Shear Zones .......................................................................................................................143

7.4

Joints and dykes ................................................................................................................143

Synthesis......................................................................................................................... 145 8.1

Paleozoic magmatism: from calc-alkaline to calcic A-types .........................................145

8.2

Mesozoic magmatism: PIC source considerations.........................................................146

8.3

Tectonic implications for the MTZ .................................................................................148

8.4

Cretaceous magmatic-hydrothermal base-metal mineralisation .................................152

References ...................................................................................................................... 154

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Appendices ................................................................................................................. 163 10.1

Whole rock geochemical data ..........................................................................................163

10.1.1 10.1.2 10.1.3 10.1.4 10.1.5 10.1.6 10.1.7 10.1.8 10.1.9 10.1.10 10.1.11 10.1.12

Paleozoic southwest Fiordland ............................................................................................... 164 Karamea Suite......................................................................................................................... 166 Paringa Suite........................................................................................................................... 167 Ridge Suite.............................................................................................................................. 168 Tobin Suite.............................................................................................................................. 169 Foulwind Suite........................................................................................................................ 170 Mesozoic southwest Fiordland ............................................................................................... 171 Rahu Suite............................................................................................................................... 174 Darran Suite felsic .................................................................................................................. 178 Darran Suite mafic .................................................................................................................. 180 Separation Point Suite felsic ................................................................................................... 182 Separation Point Suite mafic................................................................................................... 184

VIII 10.1.13

10.2

Western Fiordland Orthogneiss...............................................................................................185

Mineral X-ray Diffraction data....................................................................................... 187

10.2.1 Sample Measurement Conditions: OU 75102.........................................................................187 10.2.1.1 Main Graphics, Analyze View ...........................................................................................187 10.2.1.2 Pattern List .........................................................................................................................188 10.2.2 Sample Measurement Conditions: OU 75104.........................................................................188 10.2.2.1 Main Graphics, Analyze View ...........................................................................................189 10.2.2.2 Pattern List .........................................................................................................................189

10.3

Mineral geochemical data................................................................................................ 189

10.3.1 10.3.2 10.3.3 10.3.4

10.4

Treble Mountain Granite.........................................................................................................189 Revolver Pluton.......................................................................................................................190 Hornfelsed metasediment........................................................................................................190 Tarawera sulphides .................................................................................................................190

Geochronology - U-Th-Pb isotope data.......................................................................... 192

10.4.1 10.4.2 10.4.3 10.4.4 10.4.5 10.4.6 10.4.7 10.4.8 10.4.9 10.4.10 10.4.11

Tine Peak Tonalite OU 75184.................................................................................................192 Tine Peak Tonalite OU 75193.................................................................................................193 Big Pluton P70791 ..................................................................................................................194 Long Scarp Granodiorite OU 75175 .......................................................................................196 Upper Blacklock Granite OU 75176.......................................................................................198 Upper Blacklock Granite OU 75176 (zircon cores) ................................................................199 Revolver Pluton OU 75143.....................................................................................................200 Treble Mountain Granite OU 75108 .......................................................................................201 Treble Mountain Granite OU 75108 (zircon cores) ................................................................202 Trevaccoon Diorite OU 75154................................................................................................203 Only Island Diorite OU 75171 ................................................................................................204

10.5

Sample Catalogue............................................................................................................. 205

10.6

Abstracts ........................................................................................................................... 208

10.6.1 10.6.2 10.6.3

2004 Geological Society of New Zealand Annual Conference...............................................208 2005 Geological Society of New Zealand Annual Conference...............................................209 2005 New Zealand Minerals Conference................................................................................210

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Chapter One

1 INTRODUCTION Southwest Fiordland is an informal region south of Dusky Sound and west of Lake Poteriteri in the South Island, New Zealand (fig.1). The southwest Fiordland block has until recently been composed of a large proportion of undifferentiated plutonic rocks, which intrude Ordovician metasediments, all formally part of the Gondwanan continental margin. Fiordland has historically attracted only intermittent study due largely to its isolation, ruggedness and complexity. A greater understanding of the area became possible with the advent of plate tectonics and more specifically the recognition of a 500 km dextral offset on the Alpine Fault, which forms the boundary between the Pacific and Australian Plates. The recognition of the Alpine fault made it possible to correlate rocks of Fiordland with those in west Nelson, a region in the northwest of the South Island. More recent publications (Muir et al., 1996a, Cooper and Tulloch, 1992, Gibson, 1992, Gibson and Ireland, 1996) have highlighted a correlation with the Lachlan Fold Belt of southeastern Australia, and the Marie Byrd Land and northern Victoria Land regions of Antarctica, where all fragments are thought to have originally been juxtaposed as part of Gondwana. A long recognised correlation between the Lachlan Fold Belt and New Zealand’s Western Province has recently resulted in a renewed focus on whether world renowned Lachlan mineralisation is represented in correlative rocks within the Western Province. Most study to date has concentrated on Northwest Nelson and Westland, with little comprehensive work conducted in the correlative rocks of southwest Fiordland. Although a correlation between southwest Fiordland and Ordovician metasedimentary basement in Northwest Nelson was well established by Park (1922), higher grade metasediments north of Preservation Inlet have remained enigmatic with respect to metasediments north of the Dusky fault. Plutonic rocks have had even less attention and as Ward (1984) states; “much would be gained by a thorough geochemical study of the plutons in southwest Fiordland”.

1.1 Location Due to time constraints and the size of the region, fieldwork concentrated on the peninsula separating Preservation Inlet and Chalky Inlet, the shores of Preservation Inlet, and the lower reaches of Long Sound. Two subsequent trips focused on the southern Cameron Mountains; the first on the tops north of Blacklock Stream, and the second on the tops separating the west branch of Big River and Lake Monk.

2 A French exploration vessel captained by M. de Blosseville conducted the first survey of the area, and in 1826 published the first detailed account of the region. The peninsula separating Preservation and Chalky Inlets was named Presqu’lle Bréauté, translating appropriately as beautiful peninsula, although sadly the name has not endured on modern maps.

1.2 Aim of Research GNS Science (a government funded research institution in New Zealand) at the time of this research were conducting a mapping programme in southwest Fiordland as part of the Qmap project. This project aims to complete mapping of New Zealand geology at a scale of 1: 250000. Qmap at the time were providing encouragement for students to work in areas undergoing mapping from which the opportunity arose to work in this isolated area with logistical support from Qmap Fiordland; at this time led by Drs Ian Turnbull and Andrew Allibone (a draft of the southwest Fiordland sector of Qmap Fiordland that is now largely complete and is enclosed in the map pocket). A detailed geochemical and geochronological study was carried out with the aim of elucidating the petrogenesis of the plutonic units; the results of which are reported in the following chapters. In addition, granite hosted base-metal mineralisation at the historic Tarawera Mine, Isthmus Sound was investigated.

1.3 Outline of Thesis This thesis covers three distinct aspects of geology in southwest Fiordland. Dating work carried out as part of this thesis allowed plutons to be separated into either Cretaceous or Paleozoic. The main focus is petrogenesis of the Cretaceous Long Scarp Granodiorite, Revolver Pluton, Treble Mountain Granite, Trevaccoon Diorite, Upper Blacklock Granite, and Only Island Diorite, which occur in a region concentrated about Preservation Inlet and Long Sound (see map in back cover). The petrogenesis of the Paleozoic Tine Peak Tonalite and Big Pluton in the lower Cameron Mountains, along with other Paleozoic plutons in the region are discussed separately, as is mineralisation of the Tarawera Mine. The plutonic rocks are divided into chapters discussing, lithology, geochronology, and geochemistry, which also includes a discussion of petrogenesis. The plutonic rocks are discussed in chronological order, thus the Paleozoic plutons are discussed first although they are not the main focus. The Tarawera Mine is contained in a separate chapter, although its occurrence in Cretaceous granite suggests some obvious genetic links. A brief synopsis at the end looks to tie together the findings of this work, and put it into a broader context.

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Chapter Two

2 BACKGROUND 2.1 Regional Geology Fiordland along with Northwest Nelson and Westland constitute the Western Province of New Zealand, long recognised as the only remnant of Gondwanan continental crust exposed in New Zealand (Cooper, 1989, Cooper et al., 1982b, Gibson and Ireland, 1997). The Western Province is Cambrian to Carboniferous in age, while Permian to Cretaceous terranes to the east constitutes the Eastern Province. The Western and Eastern Province were first separated by the Median Tectonic Line (MTL, Landis and Coombs, 1967), based on the then popular concept of paired metamorphic belts. The MTL was later renamed the Median Tectonic Zone (Bradshaw, 1993, Kimbrough et al., 1993) (refer to section 2.2.1 for a more detailed discussion). A reorganization of plates resulted in the development of the plate boundary through New Zealand in the Oligocene (Alpine Fault) giving rise to the c.500 km dextral offset of Fiordland relative to Northwest Nelson (Bishop et al., 1985). However, the total displacement between Northwest Nelson and Fiordland may be up to double this as the Alpine Fault is thought to have only accommodated half of the total deformation (Sutherland et al., 2000), which makes locating the original positions of the two regions relative to each other difficult. Since the Miocene, migration of the Pacific Plate Euler pole southward has resulted in a significant component of convergence on the Alpine Fault giving rise to the Southern Alps running along the length of the South Island (Sutherland, 1995). The Western Province was first separated into three distinct allochthonous sedimentary belts, which were referred to as the Western (Buller Terrane), Central and Eastern Belts based on observations in Northwest Nelson (Grindley, 1978, Grindley, 1961). The allochthonous Central Belt model of Grindley (1961) survived for almost 30 years based on the interpretation that the central belt had been thrust northwards onto the Western and Eastern miogeosyncline (named the Buller Miogeosyncline by Grindley 1961). Cooper (1989) proposed the now widely accepted two terrane model for the Western Province after recognising a correlation between the sedimentary successions of the Central and Eastern Belts. Thus, the Central and Eastern belts were renamed the Takaka Terrane with the implication that the Central Belt was now autochthonous relative to the Eastern Belt (Cooper, 1989). The name Buller was retained for the western belt. In contrast to the Eastern Province, the Western Province is intruded by major batholiths dominated by granitoids.

4

Figure 1. Simplified geological map of Fiordland mainly showing distribution of Darran Suite rocks. Map taken from Muir et al. (1998). B, Loch Burn Formation; WMF, Wall Mountain Fault, GB, Grebe Fault; HKR, Hauroko Fault; LF, Livingston Fault; HF, Hollyford Fault; EF, Eglington Fault; GF, Glade Fault; SCF, Surprise Creek Fault; KP, Kaipo Fault/Shear Zone. Note the undifferentiated designation of southwest Fiordland.

The two Eastern Province terranes immediately east of the Western Province are the Brook Street Terrane; a Permian Island arc, and the Permo-Triassic Murihiku Terrane; a fore-arc or back-arc basin assemblage (Landis et al., 1999). Accretion of the Brook Street Terrane to the Western Province is thought to have occurred in the late Triassic (Mortimer et al., 1999a, 1999b). Docking of the Brook Street Terrane and the Murihiku Terrane is difficult to constrain, although Landis et al. (1999) suggested some degree of spatial proximity by the Jurassic. Accretion of the more eastern terranes is even less well constrained (see Wandres

5 and Bradshaw, 2005 for discussion). Translation of Eastern Province terranes along the convergent paleo-Pacific Gondwana boundary during the Mesozoic is suspected as being important, as some provenance studies link them to either Antarctica (Amundsen Province, Marie Byrd Land) or alternatively Australia (New England Fold Belt) (Wandres and Bradshaw, 2005, and references therein). The difficulty arises in that the composition of the Eastern Province terranes is not specifically unique to any particular section of the paleo Gondwana margin (Wandres and Bradshaw, 2005).

2.1 Western Province Terranes 2.1.1 Northwest Nelson and Westland Buller Terrane metasediments have a continental affinity, and are thought to have been deposited at the edge of the Australian-Antarctic segment of Gondwana, while Takaka Terrane rocks are more lithologically diverse, and are thought to have formed adjacent to a Cambrian island arc (Cooper and Tulloch, 1992, Munker and Cooper, 1997). Amalgamation of the Buller and Takaka Terrane (Pre-Baton event) along the Anatoki Thrust is thought to have occurred between the Late Silurian to Earliest Devonian, inferred from thrusting and folding affecting both terranes with the upper limit constrained by lack of deformation in the Early Devonian Baton Formation of the Takaka Terrane (fig.2, Cooper and Tulloch 1992). The Buller Terrane is thought to be the most extensive with five exploration wells in the Great South Basin indicating that Buller Terrane sediments form much of the Campbell Plateau (Bradshaw et al., 1997). Stratigraphy of the Buller and Takaka terranes have been extensively summarised by Cooper and Tulloch (1992). These authors describe the Buller Terrane as being composed of a relatively uniform suite of quartz-rich clastics and black shales of Late Cambrian to Early Silurian in age. The oldest unit is the Greenland Group (latest Cambrian-Early Ordovician), which is composed of a continentally derived turbidite sequence. The Takaka Terrane incorporates volcanics, carbonates, and coarse-grained clastics, which are notably sparse in the Buller Terrane. The Devil River Volcanics and Health Creek Beds are the oldest (Middle Cambrian) formations in the Takaka Terrane and characterise a suite of typical island arc volcanics dominated by andesitic breccia flows, tuffs, and coarse and fine-grained clastics. On the slopes of the arc and in basins, sequences of siliceous siltstone and limestone interrupted by turbiditic deposits and debris-flow conglomerates accumulated forming the Tasman Formation (late Middle Cambrian). Carbonates are represented by thick sequences of Ordovician Mount Arthur Marble.

6

Figure 2. Terrane sequences in Northwest Nelson as represented by (Cooper and Tulloch, 1992)). Only 5 out of 11 tectonic slices are represented. Suturing in the early Devonian is indicated. Asterisk indicate age diagnostic fossils. Key to tectonic events 1 = Greenland; 2 = Haupiri; 3 = Pre-Baton; 4 = Post-Baton/Post-Reefton. Key to formations, Buller Terrane: Gg = Greenland Group (undifferentiated); Gr = Roaring Lion and Webb Formations; Ga = Aorangi Mine Formation; Gl = Slaty Creek, Leslie, Douglas, and Peel Formations; Rg = Reefton Group; Kg = Karamea Suite Granite. Takaka Terrane: Hd = Devil River Volcanics; Ht = Tasman Formation; Hl = Lockett Conglomerate; Ha = Anatoki Formation; Mb = Mytton Beds (informal); Ps = Summit Limestone; Pb = Baldy Formation; Mo = Owen Formation; Ma1 = Mount Arthur Marble 1; Mw = Wangapeka Formation; Ma2 = Mount Arthur Marble 2; Eh = Hailes Quartzite; Bb = Baton Formation; Rc = Riwaka Complex.

Southwest Fiordland lithologies from Cape Providence and Preservation Inlet have been correlated with lithologies from the eastern Buller Terrane in Northwest Nelson through correlations in Ordovician fauna (Cooper, 1979, Park, 1922, Benson and Keble, 1935). Central, northern and eastern Fiordland are more enigmatic due to large parts being obscured by intrusive rocks and higher metamorphic grades, mostly between amphibolite and granulite facies; whereas Northwest Nelson is seldom above greenschist facies. Ward (1986, 1984) proposed that the continuation of the Anatoki Thrust occurs in southwest Fiordland, where he maps it as the Old Quarry Fault, south of Dusky Sound. Its continuation south of Dusky Sound has remained uncertain, although House et al. (2005) position its continuation passing through Last Cove based on results of a thermobarometry study. This might prove slightly speculative as Qmap Fiordland has spent much time trying to locate the fault in outcrop, but

7 so far have been unsuccessful (Ian Turnbull, pers. comm). At present detrital zircon and geochemical studies are been undertaken in south Fiordland to isolate the hypothesised presence of a Buller-Takaka Terrane boundary (Chloe Simpson, work in progress). Results so far are inconclusive (Simpson, pers. comm), and it might be that much of Fiordland is Buller Terrane, or the Takaka Terrane lithologies present in southern Fiordland are too similar to the Buller Terrane to discriminate easily. Some recent reconstructions removing the total deformation resulting in the formation of the plate boundary place Fiordland directly east of Northwest Nelson (fig.3). This could indicate that the Takaka Terrane exists as a slice within the Buller Terrane, which may be inherently similar to relationships seen in the originally contiguous Lachlan Fold Belt (Collins, 1998).

Figure 3. Suggested reconstruction of the New Zealand continent at 65 Ma before deformation associated with the formation of the Pacific-Australian plate margin. Note relative position of Fiordland with respect to Northwest Nelson (Beuth et al., 2006).

Both Westland and Fiordland have been overprinted by Cretaceous extensional deformation characterised by half grabens and core complexes, which culminated in the rifting of New Zealand from Gondwana and the formation of the Tasman Sea. The Cretaceous Paparoa Metamorphic Core Complex (PMCC) in north Westland is an example, and was first recognised in New Zealand by Tulloch and Kimbrough (1989). Tulloch and Kimbrough (1989) describe the PMCC as consisting of a core of high-grade paragneiss, and Cretaceous and Paleozoic plutons separated to the northeast and southwest from overlying Paleozoic metasediments, and Cretaceous and Paleozoic plutons by shallowly dipping, highly deformed

8 shear-zones. Half-grabens occur on the flanks, which are infilled with sedimentary and volcanic rocks of the Pororari Group. Much of western and central Fiordland is also considered to be an elongate metamorphic core complex and shows many similarities to the PMCC (Gibson et al., 1988).

2.1.2 Fiordland Park (1922) was first to recognise a direct correlation between graptolitic carbonaceous shales of Cape Providence, southwest Fiordland with those of the Aorere Group (later included in the Buller Terrane), Northwest Nelson. The lithologies present in the area about Preservation Inlet and Cape Providence include greywackes, slates and quartzites, which Wood (1960) referred to as the Preservation Formation. He noted that the Preservation Formation increased in metamorphic grade to the east into chlorite and biotite zone schists. Within higher grade amphibolite facies rocks to the north, Ward (1986, 1984) mapped quartzite, graphitic pelites, and quartz-rich turbidites with minor calc-silicates and marble as a continuation of the Buller Terrane into Fiordland. This he named the Fanny Terrane. This is now recognised as the northern continuation of the Preservation Formation, and is referred to as the Fanny Bay Group (Turnbull et al., 2005, see map in back pocket). To the east, Ward (1986) inferred a west dipping fault structure, which he termed the Old Quarry Fault. The Old Quarry Fault is thought to separate the Fanny Bay Group from a collection of metamorphosed conglomerates, mafic volcanics, marble, and quartzofeldspathic sandstones and siltstones, which Ward (1986) correlated with the False Edgecumbe and Middle Stream Formations of the Takaka Terrane. Ward (1986) referred to these as the Goodyear Terrane, which has now been replaced by the Edgecumbe Group (Turnbull et al., 2005, see map in back pocket). In an extensive study of metamorphism in Paleozoic metasediments, Ireland and Gibson (1998) isolated two Paleozoic regional metamorphic events. The first (M1) is a low-P (3 – 6 kbar)/high-T (630 – 680 oC) event. The second (M2) is a more typical Barrovian medium-P (5 – 9 kbar)/ medium-T (580 – 780oC) event, which is inferred to be the result of a period of crustal thickening. Dating of monazite suggests M1 occurred at c.360 Ma, and was followed by M2 at c.330 Ma. These authors thought M1 was most likely related to thermal effects of Karamea Suite rocks and equivalents in Fiordland. Western Fiordland is dominated by the Cretaceous Western Fiordland Orthogneiss (WFO), which intrudes earlier Paleozoic plutons and metasediments (Mattinson et al., 1986). North of Dusky Sound much of the WFO is metamorphosed to amphibolite facies and in parts is overprinted by granulite facies. The WFO is thought to have intruded at mid crustal depths at

9 pressures of < 8 kbar between 126 and 116 Ma, where magma accumulated to a total thickness of c.10 km (Klepeis et al., 2003a). At some point, probably before intrusion had ended, it suffered a brief major uploading event, which saw the deeper sections experience pressures between 14 – 16 kbar (Clarke et al., 2000). Not long after the loading event, extension initiated, which saw a period of rapid unroofing of western Fiordland between 105 and 90 Ma (Klepeis et al., 2003a). This extension and unroofing saw the development of a metamorphic core complex (Gibson et al., 1988). In Doubtful Sound a cover sequence of Paleozoic plutonic and metasedimentary rocks is recognised. Separating the two is the Doubtful Sound shear zone; originally interpreted by Oliver (1980) as a thrust fault, but now considered a low angle normal fault (Gibson et al., 1988). At Breaksea Sound and Wet Jacket Arm, north of Dusky Sound, clear intrusive relationships between WFO and Tuhuan Paleozoic cover sequences are preserved (Milan et al., 2005). As pointed out by Ian Turnbull (pers. comm), the existence of intrusive relationships with Paleozoic cover, combined with the observation that where inferred detachment faults are observed, orientations are inconsistent with the existence of a simple detachment fault. It is thus likely that core complex development structures are more complicated than a single main low angle detachment fault. Metamorphic grade south of the Dusky Sound is seen to decrease progressively southward, with the loss of kyanite indicating decreasing pressure (Ian Turnbull, pers. comm). Further south at Preservation Inlet no significant regional metamorphic overprint can be discerned, as is evidenced by the preservation of Ordovician fauna (Benson and Keble, 1935). Adjacent to plutons, Preservation Formation often shows contact metamorphic aureoles. From the late Cretaceous to Present, northern and central Fiordland have experienced greater uplift and exhumation than in southwest Fiordland. The structural relationship between Fiordland basement north and south of Dusky Sound has been treated as enigmatic based largely on a perceived sharp transition in metamorphic grade across Dusky Sound. Many workers have tended to infer a fault running down Dusky Sound, although its continuation to the east has not yet been identified (eg, House et al., 2005). Emplacement pressures obtained for the WFO from Wet Jacket Arm (c.11-15 kbar) at its southern extent by Milan et al. (2005) are broadly consistent with pressures determined from Doubtful Sound by Gibson et al. (1988). Between Wet Jacket Arm and Preservation Inlet, a distance of c.30 km, a pressure difference of c.9 kbar (equivalent to c.27 km based on average crustal densities) is estimated based on a minimum emplacement pressure for the WFO of 11

10 kbar and field and petrologic evidence that Preservation Inlet experienced pressures no greater than 3 kbar. Using simple geometry, and presuming Fiordland has experienced Cenozoic tilting of 40° as evidenced by southerly dips in late Cretaceous-Cenozoic sediments of c.3040° (Turnbull, per.com), a lateral distance of 36 km is calculated to account for the pressure difference of 9 kbar. This result agrees with the present horizontal separation between Wet Jacket Arm and Preservation Inlet, thus suggesting that simple tilting of a rigid crustal block may explain the observed difference in metamorphic grade.

2.2 New Zealand Intrusive Rocks Intrusive rocks of the Western Province were referred to first as the Tuhua Intrusive Group (Bell and Fraser, 1906). Although Bell and Fraser (1906) recognised older granites and metamorphic rocks, it was not until detailed mapping by Nathan (1966, 1975, 1978) that subdivision of intrusive rocks truly began, although Reed (1958) had previously noted three belts trending north-south in Northwest Nelson that had distinct chemistry. These belts were named the Paparoa Granite, Karamea Granite, and Separation Point Granite (Grindley, 1961) with the later two been renamed the Karamea Batholith and Separation Point Batholith (SPB) (Grindley, 1971).

Figure 4. Simplified geological map of Northwest Nelson and Westland showing distribution of Paleozoic and Mesozoic intrusive suites. [taken from Muir et al (1996b) but in essence that proposed by Tulloch (1988) except the Rangitoto Batholith southwest of the Hohonu is no longer distinguished].

11 Tulloch (1988) applied the nomenclature of batholith and suite by the American Geological Institute (AGI) for describing geographical and genetic relationships respectively for New Zealand plutonic rocks, which he suggested would clear up some confusion, and bring interpretation of New Zealand intrusive rocks more into line with current thinking. To clarify the AGI nomenclature; a batholith is a contiguous mass of plutonic rocks with an area exceeding 100 km2 with no connotations as to age or chemistry. In contrast, a suite comprises groups of plutons with a characteristic age, mineralogy and chemistry, which may or may not be contiguous. On this basis, Tulloch (1988) recognised five distinct batholiths separated from each other by lower Paleozoic country rock, while again emphasising the subdivision of the Karamea, Separation Point, and Rahu Suites that he had previously made (Tulloch, 1983). Of the two already named batholiths, the Paparoa was recognised as been separated from the Karamea, while the Hohonu and Rangitoto were added in Westland (refer to fig.4). The division of the Rangitoto Batholith has since been dropped after Waight et al.’s (1997, 1998a, 1998b) extensive work on the Hohonu Batholith.

2.2.1 Median Batholith [incorporating the Median Tectonic Zone (MTZ)] A more in depth study of the MTL by Bradshaw (1993) highlighted the inconsistency of placement of the MTL, with various authors placing the MTL on different sides of the dominantly plutonic Drumduan Terrane. Bradshaw (1993), Coombs (1985), and Kimbrough et al. (1993) suggested the MTL be widened and renamed the Median Tectonic Zone (MTZ) to include the enigmatic Drumduan Terrane and other plutons in Fiordland situated between the Eastern and Western Province. Kimbrough et al. (1993, 1994) emphasised the tectonised nature of MTZ rocks and inferred them to be a disseminated island arc traceable along the eastern side of Northwest Nelson and continuing along the eastern side of the originally contiguous Fiordland. Mortimer et al. (1999a, 1999b) more recently emphasised the intrusive relationships of most of the rocks of the MTZ on Stewart Island and in the Longwood Range, and proposed that most of the tectonism, especially apparent in the Northwest Nelson and eastern Fiordland sections of the MTZ, is Cenozoic in origin. This is possible, as it is thought that c.800 km of Cenozoic deformation occurred across the Australian-Pacific plate margin, of which only about half has been accommodated on the Alpine Fault (Molnar et al., 1999, Sutherland et al., 2000). Some of this deformation was almost definitely accommodated on structures within eastern Fiordland referred to collectively as the Moonlight Fault System (Lebrun et al., 2003). Cenozoic movements on faults within MTZ basement in the Taranaki Basin are known to be reactivated on Mesozoic structures, many of which are thought to be related to rifting of New Zealand from Gondwana (Muir et al., 2000). The preferred

12 interpretation of the MTZ by Mortimer et al. (1999a, 1999b) is that it is part of a continental arc batholith (which they refer to as the Median Batholith) which formed along the Gondwana margin from 230 – 125 Ma after accretion of the Permian Brook Street Terrane to the Western Province. In this model the MTZ is considered autochthonous to the Western Province. In the proposed Median Batholith, Mortimer et al. (1999b) highlighted that greater than 90% of rocks in the MTZ are plutonic with a combined area of c.10200 km2 and age range from Carboniferous to Early Cretaceous. This dominance of plutonic rocks led Mortimer et al. (1999b) to propose the term Median Batholith, which has had a mixed reception and led to widespread confusion, although most would accept the term with a more limited spatial and chronologic range. The confusion is partly due to the retention of the term Median in the name Median Batholith as this term has associations with the Median Tectonic Zone disseminated island arc hypothesis. It has been suggested that the Median Suite would be better referred to as the Darren Suite (Tulloch pers. comm.). The concept of the Median Batholith has now evolved to include most rocks of the Darran Suite and some from the Separation Point Suite. Section 2.2.3 contains a more detailed discussion of the Darran and Separation Point suites.

2.2.2 Paleozoic Suites Muir et al. (1994) found the Paleozoic components of the Karamea Batholith to be composed of plutons having a bi-modal age distribution, with the first phase intruded in the Devonian (c.380 Ma), and the second in the Early Carboniferous (c.330 Ma). The Devonian plutons incorporate the extensive O’Sullivan Granite with an age 377.7 ± 4.1 Ma (SHRIMP U-Pb), while Carboniferous plutons include the Windy Point Granite in the Paparoa Batholith segment with an age of 328.6 ± 4.4 Ma (Muir et al., 1994). These authors also noted that the Early Carboniferous Cape Foulwind and Windy Point Granite contrast to the S-type Devonian phases by having distinct A-type characteristics. Tulloch et al. (2003) has proposed subdividing the Paleozoic plutons of the Western Province into five distinct suites; two Stypes named the Karamea and Ridge suites, two I-types named the Paringa and Tobin suites, and the A-type Foulwind Suite. An overview of the proposed subdivision of Paleozoic suites by Tulloch et al. (2003) follows. 2.2.2.1 Karamea and Ridge suites The Karamea and Ridge are referred to as S-type suites by Tulloch et al. (2003), although as Muir et al. (1996b) point out this is an over simplification as the bulk of the Karamea Suite may only contains a 20 – 30 % metasedimentary component. No analogues of the extreme S-

13 type compositions described by Chappell and White (1992) in the Lachlan Fold Belt have been identified in New Zealand. The Karamea Suite is better represented in Northwest Nelson-Westland where most plutons intrude the Buller Terrane (fig.4). Karamea Suite plutons are characterized by biotite and two mica granites to tonalites, with an age range of 382-369 Ma. Type examples of the Karamea Suite include the O’Sullivans and Dunphy plutons of the Karamea Batholith. Younger Ridge Suite plutons have an age range of 353-342 Ma and are characterised by biotite granodiorite to tonalite. Unlike the Karamea Suite the Ridge Suite does not appear restricted to the Buller Terrane and is better represented in southern Fiordland and Stewart Island, where the Ridge Pluton is the type example. In Fiordland the Widgeon Granitoid Gneiss and Seaforth Pluton are also correlated with the Ridge Suite. The Ridge Suite can be distinguished from the Karamea Suite by having higher Al and Sr for equivalent SiO2 values, lower Rb and higher Ba, and slightly higher Na2O and lower K2O. Isotopically the Ridge Suite has more restricted and primitive Nd and Sr than the Karamea Suite. 2.2.2.2 Paringa and Tobin suites Two Paleozoic I-type suites are recognised. The Paringa Suite occurs as sporadic plutons within Northwest Nelson-Westland and Fiordland. Paringa plutons are hornblende-biotite gabbro to granite, and range in age from 369-359 Ma. Younger Tobin Suite plutons have ages from 349-340 Ma and occur in the Karamea Batholith, Central Fiordland and Stewart Island. The type example of the Paringa Suite is the Paringa Granodiorite in South Westland, while other plutons correlated include the Riwaka Complex, Northwest Nelson and the Pembroke Gneiss, Milford Gneiss, and Dolphin Igneous Complex, Fiordland. Although Paleozoic intrusive rocks are dominantly confined to the Buller Terrane, the Riwaka Igneous Complex (layered ultramafic-gabbroic rocks and pyroxene bearing diorite-monzodiorite) intrudes lower Paleozoic metasediments of the Takaka Terrane in Northwest Nelson. The type example of the Tobin Suite is the Tobin Diorite in the south of the Karamea Batholith. Also correlated with the Tobin Suite are the Mount Flanagan and Domett Granites in the north of the Karamea Batholith and the Lake Roxburgh Tonalite and Pomona Island Diorite in Central Fiordland. The Ruggedy Pluton on Stewart Island is also correlated. The Paringa and Tobin suites are geochemically more similar than different, although the Paringa has higher Sr at low to intermediate SiO2 values and higher Sr/Y ratios. Isotopically both suites are comparable, although both can be distinguished from the Ridge and Karamea Suites due to their more primitive composition.

14 2.2.2.3 Foulwind Suite The Carboniferous magmatism is represented by biotite (±muscovite) granites of the A-type Foulwind Suite with ages from 320-300 Ma. Geochemically the Foulwind Suite does not have strong A-type characteristics as defined by Eby (1992) and Whalen et al. (1987). The type example for the Foulwind Suite is the Foulwind Granite in North Westland. Other plutons correlated with the Foulwind Suite are the Pomona Island and George Sound Granites in Fiordland, and the Freds Camp Pluton on Stewart Island (Tulloch et al., 2003). The subtle Atype characteristics are evident in higher HFS element concentrations when compared to fractionated I-types of the Tobin Suite. The Foulwind Granite and Freds Camp Pluton plot in the within plate granite (WPG) field on the tectonic discrimination diagram of Pearce et al. (1984), although not all correlated plutons do. 2.2.2.4 Correlations of Karamea Batholith granitoids with Australia and Antarctica Cooper et al. (1982a) first proposed that the New Zealand terranes were a southeast extension of the Lachlan Fold Belt (LFB). Cooper & Tulloch (1992) went as far as correlating all major Paleozoic sedimentary, volcanic and tectonic events, based predominantly on evidence from Northwest Nelson, with those of the LFB. Regional magnetic data now supports New Zealand’s position as a southeast continuation of the LFB (Sutherland, 1999). Ion microprobe dating of Paleozoic granitoids of the Karamea Batholith by Muir et al. (1996b) showed a large amount of inherited zircon, with distinct populations at 390, 500-600, and 1000 Ma. They correlated the 390 Ma event to widespread plutonism in the Lachlan Fold Belt, while the 500-600 Ma (Ross-Delamerian age) and the 1000 Ma (Grenvillian age) age components are observed within LFB granitoids. Karamea Suite granites also share similar I-S type classifications, age, petrography, major and trace element geochemistry, and Sr and Nd isotopic compositions with granites of the LFB (Muir et al., 1996b). Karamea granites also show similar spatial correlations with Ordovician metasediments of the Buller Terrane that have proven faunal correlations with Ordovician sediments of the LFB. Although the bulk of Lachlan Fold Belt granites are older than Karamea plutons with ages from 440 to 380 Ma, an age progression down to c.360 Ma does occur within the central region of the LFB, north of Melbourne (Muir et al., 1996b). Paleozoic granitoids also occur within the Bowers and Robertson Bay Terranes of Northern Victoria Land (NVL), and Marie Byrd Land (MBL), Antarctica. The Admiralty Group aged between 390-360 Ma intrudes Ordovician greywackes of the Robertson Bay Group in NVL (Muir et al., 1996b), which have been correlated with the Greenland Group in the Buller

15 Terrane (Bradshaw et al., 1997). Also intruding Ordovician metasedimentary rocks of the Swanson Formation is the Ford Granodiorite, MBL, which consists of a series of plutons with ages between 380-350 Ma (Muir et al., 1996b). The Swanson Formation has also been correlated with the Greenland Group (Adams, 1986).

2.2.3 Mesozoic Suites 2.2.3.1 Darran (Median) Suite Darran Suite plutons are typical of continental arc calc-alkaline I-type rocks (main features summarised in Table.1) and include Triassic to Cretaceous plutons. Early studies suggested that Darran Suite rocks appeared confined to the MTZ (Kimbrough et al., 1994). The position of the MTZ between the Western Province and Brook Street Terrane meant relationships between the Darran Suite and these two terranes was enigmatic. Docking of the Brook Street Terrane with the Western Province is thought to have occurred between 245-230 Ma based on two lines of evidence (Mortimer et al., 1999a, 1999b). Firstly, plutons which intrude the Longwood Complex section of the Brook Street Terrane appear to be influenced by a continental source by 230 Ma, which is evidenced by a sharp transition in εNd values of > +7 before 230 Ma, to values less than +5 after 230 Ma (Mortimer et al., 1999a). Secondly, intra batholith shear zones were active while magma was been intruded at 220 Ma. These intrusive rocks within the Longwood Complex are referred to as the Holly Burn Intrusives (Mortimer et al., 1999b) and correlate with the late Triassic Mistake Diorite (224 Ma, Kimbrough et al., 1994), Slip Hill Diorite (229 Ma, Kimbrough et al., 1994) and Buller Diorites (228 Ma, Kimbrough et al., 1994) within the MTZ. It is likely that all these units can be included in the first phase of the Darran Suite. In addition, a second phase of plutonism within the MTZ ranging in age from 168 – 137 Ma was isolated by Kimbrough et al. (1993) and subsequently reinforced by Muir et al. (1998). Most second phase Darran Suite rocks identified to date occur within the MTZ, although increasingly, more are been identified in the Western Province. In the Northwest Nelson section of the MTZ, Kimbrough et al. (1994) included the Rotoroa Complex (155 Ma) and Drumduan Terrane (142 Ma). Here these authors noted that MTZ rocks were stitched to the Western Province by the Separation Point Batholith (SPB), while in Fiordland lithologies correlated with the SPB occur within the MTZ, effectively requiring the MTZ to be in close proximity to the Western Province my c.126 Ma. Continuing into the originally contiguous Fiordland, these authors inferred the Darran Complex (137 - 142 Ma) to be the southern continuation of the Rotoroa Complex, while also including the Halfway Peak Gabbro (146 Ma) and Lugar Burn Quartz Monzonite (141 Ma); all within the MTZ (fig.1). In Stewart

16 Island, diorite and granite within the Anglem Complex yielded dates between 141 and 138 Ma. Metasedimentary sequences of mostly Jurassic age occur in the MTZ (Full Formation, Drumduan Group, Largs Terrane, Loch Burn Formation, Teetotal Group, and Peterson Group) and are thought to be intruded by Darran Suite plutons (Kimbrough et al., 1994). Darran Suite rock with ties to the Western Province includes the Arthur River Complex (ARC) in northern Fiordland. It is inferred to be a younger phase of the Darran Suite rocks in the MTZ, although Gondwanan inheritance suggests it intruded through Western Province basement (Hollis et al., 2003). The ARC is situated between the Darran Complex and the WFO and has an emplacement age of 136-129 Ma immediately prior to the WFO (Hollis et al., 2003). Although gabbroic to dioritic in composition, these authors believe the ARC is equivalent to granites and monzonites (157-131 Ma) along the western boundary of the MTZ at Lake Manapouri. The existence of Carboniferous plutons within the MTZ was established by Kimbrough et al (1994) with two ages on the Lake Roxburgh Tonalite (337 and 343 Ma), Similar age units occur in the Western Province such as the Windy Point and Cape Foulwind granites with ages of c.330 Ma (Muir et al., 1994). Kimbrough et al. (1994) thought also that the Echinus Granite with an age of 310 Ma (Kimbrough et al., 1993), was also part of the MTZ; an interpretation also supported by Muir et al. (2000). Importantly Kimbrough et al. (1994) noted that inheritance in the Echinus Granite and Jurrassic to Cretaceous MTZ plutons required a continental basement most likely the Western Province and suggested a parautochthonous relationship between the MTZ and Western Province. Other coeval I-type rocks similar to Darran Suite lithologies intrude Western Province basement. These include the Lake Mike Granite (164 Ma, Davids, 1999) and Revolver Pluton (132 Ma, Mortimer et al., 1999b) in southwest Fiordland (see map in back pocket). In Northwest Nelson the Crow Granite (137 Ma, Muir et al., 1997) intrudes Buller Terrane metasediments, while to the north the Copperstain Creek Granodiorite (134 Ma, Brathwaite et al., 2004) intrudes Takaka Terrane metasediments. In Stewart Island the Deceit Pluton (145 Ma, Allibone and Tulloch, 2004) contains xenoliths of Paleozoic Pegasus Group schist and cuts the Paleozoic Knob Pluton south of the Escarpment Fault, which marks the northern limit of Paleozoic basement. This led Allibone and Tulloch (2004) to conclude that the Deceit Pluton intruded Western Province basement.

LoSY

HiSY

Darran

Separation

LoSY

PIC

Darran)

(Transitional

LoSY- HiSY

Rahu

Point

Classification

Suite

127

diorite,

biot

granite, 132-

105

120-

105

hbl gabbro

hbl-bt diorite, 125

Bt.

granodiorite

hbl,

monzodiorite;

2 px diorite- 127-

granite

170-

(Ma)

Age

Gabbro,

Rock Type (MALI)

Peacock

0-3 Kb

0-6 Kb

2-7 Kb

Kb

ca

ca

ca-ac

Mostly < 3 ca-ac

Pressure

Emplacement

1.1

1.25

2

1.25

360

Avg.

550

1050

400

Na2O/K2O Sr

At 65% SiO2

30

27

178

23

Sr/Y

7.6

4.4-

6.9

done

Not

0.7094

6-9

10.0

0.7058- 9.4-

0.7045

Px, plag

source

in

Minerals

Residual

plag)

gnt

(±

(±

Ma

at

plag)

130 (±gnt,

1998) 0 to +1 Amph,

plag)

110 gnt (Waight

Ma

at

-4 to -6 Amph,

(Muir

120 Ma

c. +3 at Amph,

δ18O εNd

0.7038- 5.8-

0.7038

0.733-

Avg. Sri

Table 1. Summary characteristics of representative New Zealand suites. (Data mostly after Tulloch and Kimbrough 2003 not including Karamea suite and εNd data).

17

18

2.2.3.2 Separation Point Suite (SPS) The Separation Point Suite (SPS) includes the predominantly felsic Separation Point Batholith (SPB) in Northwest Nelson, as well as an increasing number of correlative plutons in Fiordland and Stewart Island, and dioritic to gabbroic orthogneisses in western Fiordland, referred to as the Western Fiordland Orthogneiss (WFO). The SPB is a continental marginparallel intrusive belt directly inboard of the volumetrically equivalent Darran Suite (Tulloch and Kimbrough, 2003). The western margin of the SPB is in fault contact with the Takaka Terrane, while its eastern contact with the MTZ is largely obscured by tertiary sediments (Kimbrough et al., 1994). SPS correlatives have been identified intruding Darran Suite rocks in the Rotoroa Complex and eastern Fiordland (Kimbrough et al., 1994, and references therein). Correlative rocks have been found to continue along strike for at least 900 km (Tulloch and Kimbrough, 2003). On Stewart Island, SPS plutons mainly intrude Western Province basement south of a major Cretaceous northeast (outboard) dipping thrust fault, while earlier Darran Plutons mainly occur to the north of this fault structure (Allibone and Tulloch, 1997, Allibone and Tulloch, 2004). The SPB was intruded over a relatively short time interval from 119-116 Ma (Muir et al., 1994). Correlative SPS plutons in eastern Fiordland are slightly older with ages between 121 and 124 Ma (Muir et al., 1998), while on Stewart Island SPS plutons span an age range from 125 to 105 Ma (Allibone and Tulloch, 2004). The WFO is broadly coeval with an age range between 126 and 119 Ma (Hollis et al., 2003, Mattinson et al., 1986, Muir et al., 1998). SPS rocks are distinctive in showing no inheritance. High SiO2 SPS rocks have very high Na and as a result they have been compared to TTG suites and modern adakites (Muir et al., 1995). In detail SPS rocks are alkali-calcic, have high Na2O, Al2O3, and Sr (typically >500ppm and up to 1000ppm) with conversely low Y (≤5ppm) and strong depletions in the HREE (Muir et al., 1995). Although high Sr/Y ratio was recognised as typical of adakitic rocks, a lack of MORB isotopic characteristics meant another mechanism needed to be invoked. Primitive isotopic compositions with initial 87Sr/86Sr ratios of c.0.7038 and εNd values of c.+3 highlighted the lack of interaction with continental crust (Muir et al., 1998). Muir et al. (1995) proposed partial melting of a basaltic amphibolite source to explain these features, favouring newly underplated material derived from a source that separated from depleted mantle at 500-600 Ma. They refer to this source as Separation Point depleted mantle (SPDM). Waight et al. (1998b) and Muir et al. (1996b) used SPDM as

19 an end member for isotopic mixing in petrogenesis of the Hohonu and Karamea suites respectively. This seems a requirement for almost all suites intruding Western Province basement, which require a more mafic source with a more evolved Nd isotopic composition than the mantle at their time of generation. Elevated Na2O and Sr are attributed to complete breakdown of plagioclase, while depletion in Y is attributed to retention in residual garnet (Muir et al., 1995). Tulloch and Kimbrough (2003) suggested the term HiSY for SPS like compositions referring to an acronym for the characteristic enrichment in [S]trontium and depletion in [Y]ttrium displayed in these rocks. They suggested the term to differentiate between adakitic rocks, which are dominantly volcanic and formed by inferred partial melting of subducted oceanic crust (Defant and Drummond, 1990), and ‘SPS-like’ rocks that are dominantly plutonic, and formed by partial melting of an underplated amphibolite source (Atherton and Petford, 1993). In Fiordland HiSY chemistry occurs sporadically before the main SPS phase, such as felsic sheets dated at 136.8±1.9 Ma , which intrude the Darren Complex (Wandres et al., 1998), and the Lake Mike Granite with an age of 164 Ma (Davids, 1999). This indicates that the conditions that produce HiSY chemistry can occasionally arise during predominantly normal calc-alkaline arc magmatism and that these earlier HiSY-like rocks may not necessarily be correlatives of the SPS. The WFO is considered the lower crustal analogue of the SPB of Northwest Nelson based on correlated age and similar HiSY chemistry, although it is average significantly more mafic (Muir et al., 1995). The WFO rocks are composed of vertically stratified mafic to intermediate gneisses (Mattinson et al., 1986 and others). Early phases are gabbroic in composition with subordinate ultramafics, while later phases are dominated by diorite. Early work by McCulloch et al. (1987) established a Rb-Sr isochron age of 120 ± 15 Ma, and inferred a mantle like source based on low initial 87Sr/86Sr ratios of 0.703 ± 4 and high εNd values (-0.42.7). These values are similar to the SPB, although lower than SPS plutons of eastern Fiordland (εNd = +3, Muir et al., 1998). The WFO is now interpreted as a root zone of the SPB, which accumulated c.10 km of magma between 126-116 Ma, and was almost immediately subjected to granulite grade metamorphism (Clarke et al., 2000, and references therein). McCulloch et al. (1987) highlighted the restricted time duration of the high-grade metamorphic event by establishing that the WFO had cooled to c.3000C by c.100 Ma. Peak metamorphism probably occurred at c.120 Ma as evidenced by 120 Ma metamorphic overgrowth on zircons from the ARC (Hollis et al., 2003). Melt segregations within WFO that had experienced granulite grade metamorphism have similar chemistry to the SPB. This led

20 Stevenson et al. (2005) to suggest that the SPB was sourced directly from remelting of WFO. Klepeis et al. (2003a, 2003b) suggested that following accumulation of this more felsic melt, crustal extension allowed the magma to ascend via upper-middle crustal shears from 116-105 Ma. The exact relationship between the mafic WFO and felsic SPS is still unresolved, and while more felsic SPS magma maybe sourced by remelting of WFO; at the very least the two probably shared the same source. A possible explanation of the largely bimodal composition between the WFO and SPB is that both have a common intermediate parent through fractional crystallisation. The intermediate primary melt after reaching neutral buoyancy in the lower crust fractionated allowing the more felsic SPS magma to carry on its ascent. Although some WFO are obviously cumulate having large positive Eu anomalies, the respective negative anomalies in SPB rocks are largely absent, thus making this hypothesis unlikely. 2.2.3.3 Rahu Suite Rahu Suite plutons mostly occur in Western Province basement to the west of the MTZ. They are broadly coeval with SPS magmatism with which they are believed to be related. Rahu Suite rocks were first described by Tulloch (1979) from the Rahu Saddle area, northeast of Reefton where smaller plutons intrude Paleozoic Karamea Batholith rocks. Subsequent work (Tulloch, 1988, Graham and White, 1990, Tulloch and Kimbrough, 2003, Waight et al., 1997) appears to indicate that the distribution of the Rahu Suite in Northwest Nelson is largely westward (inboard) of the Darran and SPS magmatic belts (fig.4). Tulloch (1979) described the Rahu Suite as consisting of biotite granodiorite to biotite-muscovite granite and ambiguous with regards to I-S type classification. Graham and White (1990) regarded the Buckland Granite forming the bulk of the Paparoa batholith as having the most typical Rahu Suite chemistry and is considered the type example. Tulloch (1983) suggested that Rahu Suite plutons may be the result of contamination of SPS magmas by Western Province metasediments based on age similarity and intermediate Sri between SPS and Karamea Suite. Dating of zircons in the Buckland Granite using the U-Pb SHRIMP spot technique by Muir et al (1994) confirmed that Rahu Suite plutons contained considerable Gondwanan inheritance. Of the 32 spot analyses on 30 zircons, 13 gave ages older than 300 Ma. Removing these components, they suggest a crystallisation age of 109.6±1.7 Ma. Dating of plutons in the Hohonu Batholith by Waight et al. (1997), which had previously been included in the Rahu Suite (Tulloch, 1988), found it to be dominated by plutons ranging in age between 114 and 109 Ma. Waight et al. (1997, 1998b) recognised two distinct suites within the Hohonu, which

21 they name the Te Kinga and Deutgam Sub-Suites. The Te Kinga Suite is compared to the Buckland Granite and is inferred to have involved high fH2O during melting. Compositions of the Te Kinga Suite have HiSY characteristics similar the SPS, although their Sr/Y ratios are believed to be subdued due to contamination by Greenland Group. In contrast, the Deutgam Suite has distinctly more mafic I-type compositions, and is inferred by Waight et al. (1998b) to be the result of dehydration melting (low fH2O). Waight et al. (1998b) believes the Te Kinga and Deutgam Suites had a source very similar to the SPS, referred to by Muir et al. (1995) as the Separation Point Depleted Mantle (SPDM). Modelling of mixing between SPDM and Greenland Group suggests about c.20% Greenland Group contamination (Waight et al., 1998b). Supported by more recent data from Rahu plutons which showed elevated initial Sr and δ18O, Tulloch and Kimbrough (2003) concluded that they were most likely SPS magmas that had variably incorporated a component of Paleozoic metasedimentary crust. This is further supported by chondrite normalised REE plots which follow a HiSY pattern rather than LoSY. Although Rahu Suite plutons mostly return ages of between 120-110 Ma, older ages have been obtained from plutons showing similar chemistry and spatial distribution, such as the Rocky Creek Pluton, Rahu Saddle, with a conventional U-Pb monazite age of c.132 Ma (per. com. Andy Tulloch). Rahu Suite correlatives are now recognised on Stewart Island, which include Blackies (c.116 Ma) and Upper Kopeka plutons (Allibone and Tulloch, 2004). No Rahu Suite plutons have yet been recognised in Fiordland, although similarities between southwest Fiordland and Stewart Island make their existence likely. 2.2.3.4 French Creek Suite (post orogenic alkaline magmatism) The French Creek Suite was proposed by Waight (1995) for plutons and intrusives related to New Zealand-Australia rifting. Alkaline magmas related to the rifting of New Zealand from Antarctica and Australia occur along the continental boundary of the Western Province with reported ages of c.110-80 Ma (Tulloch and Kimbrough, 2003). Waight et al. (1997) describe the occurrence of A-type granite and coeval doleritic-lamprophyre dyke swarms inland from Greymouth in North Westland. The French Creek Granite Pluton gives a crystallisation U-Pb SHRIMP age of 81.7±1.8 Ma (Waight et al., 1997), possibly similar to pegmatite dykes occurring throughout the Alpine Schist in Westland. The Mataketake Range pegmatite gives a discordant conventional U-Pb monazite age of c.68 Ma inferred to represent a thermal event associated with Cretaceous extension (Batt et al., 1999).

22

Figure 5. Tulloch and Kimbrough’s (2003) model for formation of HiSY suites inboard of LoSY suites. A. Cordilleran arc with emplacement of a voluminous LoSY suite and resultant magmatic underplating. Underplating transgresses the dashed line representing the garnet stability field. B. Shallowing of slab displaces the underplated mafic root continent-wards of the subduction margin. C. Partial melting of the mafic underplate results in HiSY plutons emplaced inboard of the earlier LoSY magmatism.

2.2.4 Model for the formation of HiSY magma High aluminium granitoids (Al2O3>15% at 70 % SiO2) are characteristic of the Archean granite-gneiss terranes where they form the tonalite-trondhjemite-granodiorite (TTG) series (Smithies, 2000, Drummond and Defant, 1990). Additional geochemical characteristics include Yb < 1, Na2O/K2O > 1, La/Yb normally > 30, and Sr and Ba > 500, broadly similar to Mesozoic adakites (Smithies 2000). TTG compositions have been attributed to melting of a hydrated basaltic crust in the garnet stability field (Atherton and Petford, 1993), where adakites are thought to be partial melts of subducted oceanic crust (Defant and Drummond, 1990). Smithies (2000) showed volcanic adakitic rocks can be characterised by low to intermediate SiO2, high Mg# (Mg#~50), and high Ni and Cr. These features are attributed to interaction of the slab-derived melts with the mantle wedge through which the magma must

23 traverse. TTG rocks, however, show no such evidence of interaction with the mantle, but are noted to have the same Mg# and SiO2 values as Phanerozoic Na-rich granitoids (termed HiSY by Tulloch and Kimbrough 2003). Smithies (2000) invoke the same process for generation of both the TTG and Paleozoic Na-rich suites by partial melting of a hydrous basaltic lower crust that has been magmatically or tectonically thickened. However, during the Achaean ‘HiSYtype’ magmatism may not necessarily have been restricted to regions of active subduction due to higher mantle temperatures (Smithies 2000). Tulloch and Kimbrough (2003) favour a tectonic thickening event as a mechanism for generation of the SPS caused by displacement of the arc root to greater depths; possibly by slab shallowing (fig.5).

2.3 Southwest Fiordland 2.3.1 Previous work A ship’s surgeon made the first geological examination of southwest Fiordland when the H.M.S Acheron first surveyed the coastline from 1849 to 1851. These preliminary geological investigations were followed with work by Hector (1863), Hutton (1875), and Park (1922). McKay (1896) conducted a more detailed account of the Preservation Inlet area, specifically focusing on mineralisation following the establishment of a goldfield after alluvial gold was discovered in 1889. The most systematic previous geological study of the metasedimentary basement was carried out by Professor Benson and others (Geology Department, University of Otago) in 1932 whose four companion papers (Benson, 1934, Benson and Bartrum, 1935, Benson and Keble, 1935, Benson et al., 1934) on the Preservation and Chalky Inlet area provide the most detailed account to date of the metasedimentary basement. The main focus of these authors was to infill gaps in the Ordovician biostratigraphy to obtain a better correlation with the Castlemaine Series of Victoria, Australia. Mid Cretaceous and minor Eocene-Oligocene sediments from Seek Cove, Coal Island and along the coast east of Coal Island have been studied by Lindqvist (in prep), while mapping with emphasis on these units was published by Bishop (1986) in the Puysegur Sheet. Gulches Head was mapped by Badger (1973), where he recognised four different intrusive phases. The area from the head on Long Sound to just north of Dusky Sound was mapped by Ward (1984). Powell (1997) worked to the east of Ward where he subdivided, mapped and informally named several more plutons, although unfortunately none of this work was submitted. Qmap Fiordland recognising the value of work by Powell has made arrangements with Powell to use the thesis material, and have incorporated much of his mapping and suggested unit names into the draft of Qmap Fiordland (pers. comm Ian Turnbull). Much of this thesis work has been incorporated into the draft of Qmap Fiordland.

24

2.3.2 Paleozoic metasediments Southwest Fiordland basement is dominated by a continuous sequence of tightly folded Ordovician flysch sediments (Benson, 1934). Park (1922) described the sediments from Cape Providence as greywackes with shale horizons, which strike north-north-east to south-southwest, while dipping steeply to the west. Carbonaceous slates contain beds which incorporate Middle Ordovician graptolites, which as mentioned earlier are correlated with Aorangi Mine Formation, Buller Terrane (Park, 1922). Benson (1934) observed quartzite horizons, and noted that the pelitic black carbonaceous argillite layers, even at the lowest grades show the first signs of metamorphism with the development of spots. About 20 km east of Cape Providence at the Morning Star Mine site slates contain lower Ordovician graptolites. Park (1922) made the inference that the age interval between Cape Providence and the Morning Star Mine would not correlate with the depth of sediments observed (being some 20 km), and thus the strata must be folded in a series of tight truncated folds. The Ordovician sediments generally increase in metamorphic grade moving to the north and east from Cape Providence; the sediments forming higher grade schists from the middle of Long Sound northwards (Benson, 1934).

2.3.3 Plutonic rocks Intrusive rocks were described by Benson and Bartrum (1935), who although recognising different units did not formally subdivide them. The term Kakapo Granite was introduced by Wood (1960) as an all encompassing term for plutons intruding metasediments in Preservation and Cunaris Sound. Although recognising different phases, Wood thought they were probably Late Paleozoic and coeval with the ‘Pomona Granite’. Bishop (1986) chose the type section for Kakapo Granite as the coastline around Revolver Hill, and although again recognising different phases he thought all compositions from mafic to felsic were genetically related. The only real subdivision made by Bishop, was the Red Head member previously recognised by Badger (1973); a red fine grained quartz monzonite forming much of Gulches Head. Although acknowledging a mid Cretaceous Rb-Sr whole rock age obtained by Aronson (1965, 1968) from a granite boulder on the shores of Chalky Sound; Bishop concluded after consulting Andy Tulloch that the Kakapo Granite, at least in part, was most likely Paleozoic based on correlations made with granitoids from the Karamea Suite in Northwest Nelson. Powell and Kimbrough (1987) described the Kakapo Granite as an arcuate belt running from the south side of Dusky Sound down to Preservation Inlet. Conventional U-Pb zircon dating by Powell and Kimbrough (1987) made the first subdivision of Kakapo granite based on an Early Carboniferous U-Pb zircon age obtained from the north end of the belt and a mid

25 Cretaceous age obtained from the Kakapo Granite type section as defined by Bishop (1986). Discordance was noted in both calculated dates which they attributed to inheritance of Proterozoic Pb. On the basis of the younger Cretaceous age, the name Revolver Pluton was proposed by Powell and Kimbrough (1987). The first serious mapping away from the coastline was carried out by Ward (1984) who subdivided and mapped several plutons between Cunaris and Dusky Sound. The plutons mapped and informally named included the extensive Lake Mike Granite (lmg), Mount Solitary Granodiorite (msg), Seal Lake Gabbro (slg), Dolphin Igneous Complex (dsi), Staircase Tonalite (sto), Widgeon Granitoid Gneiss (wgg), and Lake Row Granite and Gabbro (these units are now incorporated in Qmap Fiordland, see draft in back cover). The Mount Solitary Granodiorite and Lake Row Granite and Gabbro have subsequently been dated by Davids (1999), which yielded the following SHRIMP U-Pb zircon ages of 371.1 ± 3.0 Ma, 334.9 ± 5.8 Ma, and 380-400 Ma respectively. Dating of the Lake Mike Granite was first attempted by Ward who obtained a slightly discordant conventional U-Pb zircon age from Needle Peak, from which he proposed the age lay in the range of 167-174 Ma. Davids obtained a better resolved age from the same sample site with a SHRIMP U-Pb zircon age of 163.7 ± 3.0 Ma (Needle Peak); although still within error of age obtained by Ward. A Paleozoic age of 382.7 ± 1.8 Ma was obtained from a different location near Lake Mike by Davids, which added confusion. The two contrasting ages obtained from different localities suggest two plutons compose what was originally mapped as a single unit. On the basis of textural differences, Davids concluded that the pluton was dominated by the Mesozoic component, while the Paleozoic age component was confined to the region about Lake Mike. Muir et al. (1998) dated granite from Solitary Peak (KAK1) in the southeast of the belt, which yielded a poorly defined Carboniferous age (SHRIMP U-Pb zircon). Their best estimate of the intrusive age was c.340 Ma, although the dating precision can only confirm a Carboniferous age. They chose to call the sample Kakapo Granite even though the type example for Kakapo Granite had been situated along the shore of Revolver Hill by Bishop (1986), which by then had been shown to be Cretaceous in age, and had subsequently been renamed the Revolver Pluton by Powell and Kimbrough (1987). Ewing (2003) looked to clarify the poorly resolved age on KAK 1, by attempting to redate the sample. After analysing 14 grains from the same heavy mineral separate, the results revealed no grains of c.340 Ma in age, and Ewing decided from the results that the age was most likely between 350 and 370 Ma; while also noting inheritance of c.380 and c.390 Ma. Two other samples of Kakapo Granite were also dated (KAK 2A, south of Mt Aiken and KAK 3B, west of the Fred Burn on the South Coast, see

26 map in back cover). KAK 2A gave similar results to KAK 1, while KAK 3B, although having some Gondwanan inheritance had a more simple age distribution between 375 and 350 Ma. Ewing pooled the data between the three samples, based on an assumption that all samples were the same age, and concluded a preferred age of c.356 Ma. In addition Powell (1997) mapped the Thundercleft Quartz Diorite (tqd) and Tine Peak Tonalite (tpt); both extensive units occurring east and north of Long Sound (see map in back cover). Powell recognised that both these units were likely to be older than the Lake Mike Granite based on their intrusive relationships with deformational phases. Qmap Fiordland has identified additional plutons to those already mapped by Ward and Powell. The names and mapping used in this thesis are largely borrowed from Qmap Fiordland draft map (Turnbull et al., 2005), which is acknowledged to be work in progress; and it is recognised that the final published version of Qmap Fiordland may include different unit names and interpretations than the ones used in this thesis. Some interpretations of unit ages suggested in the draft map are my own. New plutons identified by Turnbull et al. (2005) include a lithologically distinct granite and diorite pluton intruding the Revolver Pluton on the Treble Mountain Peninsula. These units are informally named the Treble Mountain Granite (tmg) and Trevaccoon Diorite (tdi). Qmap Fiordland has placed a lot of emphasis on defining the contact between the Revolver Pluton and Kakapo Granite (pers. comm. Ian Turnbull), the results of which show that the Revolver Pluton is a regionally extensive unit (see map in back cover). During work east of Long Sound, the Monk Granite (mg) has also been subdivided, and is thought also to be Cretaceous. It is now believed that the Paleozoic component of the Kakapo Granite is isolated and it has tentatively been renamed Big Pluton (bp) to dissociate it from the confusion which surrounds the name Kakapo Granite (Turnbull et al., 2005). South of Dusky Sound, the region which has had the least work is between Edwardson Sound and West Cape, which has resulted in the mapping of many new plutons (Turnbull et al., 2005). These include the Mount Evans Pluton (ep), Anchor Island Intrusives (aio), Newton River Granodiorite (nrg); all interpreted as Paleozoic based on unpublished dating or geological relationships (Allibone and Turnbull, in prep). New plutons thought to be Cretaceous include the Brothers Pluton (bro), Inaccessible Diorite (id, marked as undifferentiated on the draft map), Fannin Pluton (fap) and North Port Granite (npg). In addition many undifferentiated smaller mafic bodies have been mapped. Additionally, dating, geochemistry, and lithology of some rocks in this thesis recognise three new units. Dating in the north of the Revolver Pluton reveals a significantly younger phase whose boundaries were not mapped, which is here named the Upper Blacklock Granite (ubg).

27 In addition a coeval granodiorite occurs in two localities on the margins of Revolver Pluton and is here named the Long Scarp Granodiorite (lsg), although this unit was not mapped. A mafic body crossing Long Sound mapped by Turnbull et al. (2005) was dated and here named the Only Island Diorite, which is now incorporated in Qmap Fiordland. Other geochronology in southwest Fiordland includes a more precise TIMS U-Pb monazite age for the Revolver Pluton from Revolver Hill of c.132 Ma by Tulloch (age reported in Mortimer et al., 1999b). Unpublished Carboniferous ages have been obtained for the Evans Pluton, Widgeon Granitoid Gneiss, Tine Peak Tonalite, and Newton River Granodiorite by Tulloch (referred to in Tulloch et al., 2003 and by pers. comm.). These Carboniferous ages are similar to plutons dated by Muir et al. (1998) and Kimbrough et al. (1994) in eastern Fiordland, including the Kakapo Granite, Hauroko Granite, Poteriteri Pluton, Lake Roxburgh Tonalite, and Pomona Island Diorite.

2.3.4 Cretaceous and Tertiary sediments Wood (1960) first named a sequence of basal breccias and conglomerates overlain by sandstone, mudstone, and coal seams the Puysegur Group; based on its apparent geographical restriction to this region. Pollen contained within the coal seams gave mid Cretaceous ages. Wood also observed that the Puysegur Group lay on an irregular fresh surface of Paleozoic rocks. More detailed study by Lindqvist (1975) of the Puysegur Group suggests an overall thickness of 1 km, where on the south side of Coal Island it is in apparent fault contact with Ordovician metasediments that form the northern part of the island. The sediments are dominated by sandstones with mudstone with rare carbonaceous horizons, which are named the Windsor Formation. In Seek Cove, basal mid Cretaceous sediments of the Puysegur Group form rift related conglomerates, which have been observed to overlie weathered granite on Gulches Head (Lindqvist, in prep). Conglomerates containing volcanic and plutonic cobbles form the basal unit and are found at Seek Cove and Prices Beach (Lindqvist, in prep). The plutonic cobbles resemble plutonic rocks of the region, although volcanic sediments have not been observed. Unconformably overlying the Cretaceous sediments at Puysegur Point and eastwards along the south coast are remnants of an Eocene-Oligocene sequence that dips gently seaward. The sediments are composed of sandstone and mudstone with minor coal horizons in the Eocene (Lindqvist pers. comm). Coeval ages, rift related volcanism, and the same depositional settings for Puysegur Group and Pororari Group sediments lead Powell and

28 Kimbrough (1987) to suggest both are correlated to mid Cretaceous extension prior to the breakup of Gondwana.

29

Chapter Three

3 LITHOLOGICAL UNITS Mapping of southwest Fiordland as part of Qmap Fiordland is largely complete and a draft made available by Dr Ian Turnbull is enclosed in the map pocket and shows all units discussed in this chapter. Unit names may be subject to change in the final version and subsequently may differ from the ones used in this thesis. The following sections describe the Preservation Formation, which hosts many of the plutons examined in the Preservation Inlet region in addition to the Paleozoic and Cretaceous plutons examined as part of this study.

3.1 Preservation Formation (Fanny Bay Group) The Preservation Formation forms most of the Fanny Bay Group and extends up west side of Treble Mountain Peninsula, while forming much of the coastline in Preservation Inlet (see map in back cover). It extends as a belt at least as far north as Dusky Sound, where it grades into higher grade schist. In Preservation Inlet the Preservation Formation forms well bedded interlaid argillites and sandstones. The bedding is mostly steeply inclined, and on the north side of Coal Island, the dips suggest tightly folded vertical sequences in agreement with Benson (1934). Minor interbedded, highly carbonaceous shale horizons occur on the north shore of Coal Island and Powells Beach and are known to contain graptolites, which were documented by Benson (1934). Metasediments adjacent to larger intrusions have generally well developed hornfelsic textures. Recrystallisation of greywacke adjacent to Revolver Pluton in Brokenshore Bay includes a completely recrystalised assemblage of biotite, muscovite, albite and quartz (OU 75126). Abundant well formed cordierite was documented by Benson (1934) in argillites from around Preservation Inlet, although none has been identified in this study. Current mapping as part of Qmap suggests that cordierite disappears further to the north as the sediments progressively increase in grade (Turnbull, pers. comm). Sandstone collected from Powells Beach is composed of poorly sorted quartz with a significant matrix component composed of sericite and larger muscovite grains combined with finer grained material, which is inferred to be a clay mineral (OU 75166).

3.2 Intrusive Rocks Lithological descriptions of intrusive units in southwest Fiordland are limited to units observed during fieldwork and samples collected for petrographic and geochemical analysis. Although geochemistry of units outside the immediate field area are included for comparison,

30 field relationships are either assumed unknown or are accepted from work by IGNS. Geochronological and geochemical work undertaken in this thesis has established the existence of an Early Cretaceous plutonic province showing clear intrusive relationships in a geographically constrained region about Preservation Inlet and the lower reaches of Long Sound. For ease of discussion, these rocks when referred to as a group will be called the Preservation Intrusive Complex (PIC). The term PIC is not necessarily proposed as a lithostratigraphic term and does not strictly comply with the nomenclature of a complex as defined by the IUGS. The rocks included in the PIC are those examined in the field including the Revolver Pluton, Treble Mountain Granite, Upper Blacklock Granite, Long Scarp Granodiorite, Trevaccoon Diorite, and Only Island Diorite, as well as various undifferentiated dykes and stocks in the area. Units mapped by Qmap Fiordland but not visited in the field, which are thought to be Cretaceous based on lithologic and geochemical relationships, are also included. These include the Monk Granite, Northport Granite, and Brothers Pluton. The Lake Mike Granite is not included as it is correlated with the early phase of the Darran Suite based on its age. Other units in the region that could be correlatives of the PIC include the Supper Cove Orthogneiss (128 Ma, Tulloch and Kimbrough, 2003), and Fannin Pluton. Although unit names have not yet been formalised in the literature, when referred to in the text their abbreviations are capitalised to make them stand out.

3.2.1 Paleozoic Intrusive Rocks 3.2.1.1 Tine Peak Tonalite (tpt) The Tine Peak Tonalite (Tpt) was first mapped and informally named by Powell (1997). The unit covers an area of c.50 km2 and is elongate in a north-south direction where it extends from the north end of the Cameron Mountains to about 3 km south of Lake Monk in the south of the Cameron Mountains. Allibone and Turnbull (pers. comm) during preliminary work for Qmap Fiordland have identified lithologies from tonalite to granite and have informally proposed the name Houseroof Pluton to reflect this variation. I will however use the name Tine Peak Tonalite, as in outcrop massive Tpt appears to be dominated by tonalite lithology, and as yet the name Houseroof Pluton has not been formalised. Granitic dykes were observed to crosscut Tpt in places, although are most likely Cretaceous in origin. The Tpt was examined from Rugged Mount in the north to as far south as spot height 1217 m. In outcrop the Tpt is a complex unit consisting of predominantly biotite-muscovite garnet tonalite intruded as stocks and dykes. Lit-par-lit intrusions (plate. 2D) of granite also invade the extensive metasediments that compose the lower Cameron Mountains, which are now metamorphosed from upper greenschist to amphibolite facies. The lit par lit intrusions in thin

31 section show medium grained biotite schist invaded by unfoliated biotite granite, which may or may not be related to the more typical tonalitic stocks and dykes. The lack of deformation of the granite implies intrusion occurred after peak deformation. The granite is possibly sourced from deeper metasediments adjoining the Tpt to the east, which according to observations made by Turnbull (pers. comm) are extensively migmatised. Calc-silicate rafts and a raft of marble 50m thick on the ridge above 1304 m have been mapped during Qmap fieldwork, as well as rafts of gabbroic and dioritic rock in the north of the pluton (pers. comm Ian Turnbull). A

C

B

DD

E

Plate 2. (A) Tine Peak Tonalite OU 75181; (B) Tine Peak Tonalite OU 75189; (C) Tine Peak Tonalite OU 75184; (D) Lit par lit intrusion of metasediment by biotite granite; (E) Big Pluton P 70589.

32

D

E

Plate 3. Photomicrographs of Tine Peak Tonalite samples. (A) Mafic clot of biotite, muscovite, and garnet in plain light, and (B), under cross polars. (C) Abundant zircon occurs both within biotite and quartz (OU 75184). Note the presence of muscovite which is absent from the less evolved sample OU 75189 (D, plain and E, cross polars). OU 75189 is also distinct from OU 75184 in having more euhedral garnet and more abundant magnetite.

In hand specimen the Tpt is a medium to coarse grained tonalite comprising approximately equal amounts of grey quartz and cream plagioclase with interstitial clots of biotite, muscovite, magnetite, and red garnet (plate. 3A and D). Accessory phases occurring in the mafic clots include abundant apatite and zircon, where zircon occurs in biotite and quartz (plate. 3C). Quartz and plagioclase form anhedral grains up to c.1 cm. More mafic tonalite examined below the summit of 1258 m (OU 75189, plate. 3B) lacks muscovite, but still

33 contains biotite and garnet, which suggests that muscovite-bearing tonalite is a more evolved facies. Biotite is straw to brown in colour and dominates muscovite by a ratio of approximately 5:1 in the more widespread evolved lithologies (plate. 3B). Garnet in the more evolved rocks occurs as poorly formed grains up to 3 mm within the clots of mica, and more occasionally as isolated grains. In the less evolved rocks the garnet is more euhedral (plate. 3D and E). Occasionally, garnet is mantled by plagioclase, which separates the garnet from quartz (plate. 3B); although this relationship is not consistent, suggesting the relationship is coincidental. 3.2.1.2 Big Pluton (bp) Kakapo Granite has now been informally renamed Big Pluton (Bp) by Qmap Fiordland, which I agree is necessary for reasons discussed in section 2.3.3. The Bp was not examined in the field, although a study of the geochemistry and thin sections supplied by Dr Ian Turnbull was undertaken. From this study it became apparent that eastern samples were consistently geochemically distinct from western samples. For ease of discussion I will refer to more eastern samples as Big Pluton east (BpE) and more western samples as Big Pluton west (BpW). Work in this thesis also suggests that mapped Bp contains Cretaceous lithologies which can be correlated with the Long Scarp Granodiorite. Dating of samples has until now been on samples correlated with BpW (refer to discussion in section 2.3.3). A sample of BpE (P 70791, collected during Qmap Fiordland fieldwork and supplied by Dr Ian Turnbull) was selected for dating to confirm whether it was a marginal facies of the BpW, or potentially related to geochemically similar Cretaceous I-type plutons. The dating yielded a well resolved age of 366.5 ± 2.2 Ma, which is seemingly older than the age suggested by Ewing (2003) of c.356 Ma and Muir et al. (1998) of c.340 Ma for BpW, although clearly BpE is not correlated with Cretaceous plutons. Due to the lack of resolution on the ages obtained by previous authors it is difficult to confidently separate the two phases on age alone and thus both BpE and BpW will be treated as different phases of the same pluton. One sample of BpW was obtained from Dr I.M Turnbull for petrographic examination (P 70589, plate. 2E). The rock is a medium grained biotite granodiorite, although this sample is apparently not as coarse-grained as the majority of the Bp (Turnbull, pers. com). Mineralogically the sample is composed of anhedral plagioclase and quartz of equal size, and smaller grains of anhedral microcline. Red-brown biotite clots and strings occur interstitially between the feldspar and quartz. Subordinate amounts of muscovite occur with the biotite, combined with accessory apatite and zircon. Opaque phases are rare in contrast to Cretaceous

34 rocks observed. A lower magnetic susceptibility, qualitatively determined, likely confirms a lack of magnetite in BpW, which is consistent with the more S-type geochemical character of BpW and more reduced composition. The lack of pink orthoclase means BpW lacks the pink colour of Revolver Pluton (Rp, refer to section 3.2.2.1), which again suggests a more reduced composition. The Red-brown colour of the biotite distinguishes it from the straw brown biotite of Rp. Mapping and petrographic analysis suggests BpE, which composes about 20% of Bp (including P 70791 dated in this study), is lithologically distinct in having darker biotite and containing titanite, clinozoisite, and allanite (Allibone, pers.comm). In contrast, BpW contains minor accessory muscovite and garnet. Based on the existence of transitional samples (P 70778) between east and west phases that contain biotite-garnet-muscovite-titaniteclinozoisite, Allibone (per.com) prefers the interpretation that BpW and BpE represent a compositionally zoned pluton. A

B

C

D

Plate 4. Revolver Pluton samples, (A) OU 75107; (B) OU 75122; and (C) OU 75143; (D) Treble Mountain Granite.

35

3.2.2 Preservation Intrusive Complex (PIC) 3.2.2.1 Revolver Pluton (rp) The Revolver Pluton (Rp) is massive porphyritic biotite granite covering an area of c.200 km2. It forms a belt between Isthmus Sound and the Cameron Mountains, and is continuous from lower Long Sound in the north, to the South Coast (Turnbull et al., 2004; see map in back cover). Strong magnetic anomalies in Preservation Inlet indicate that Rp or equivalents probably underlie the Ordovician metasediment at shallow depths, and probably continue offshore (Oliver and Coggon, 1979). Suspected shear structures are observed cutting the Rp in NW-SE trends, with resultant grain size reduction, giving rise to a more equigranular rock (see section 7.3). This is most evident in the disappearance of striking megacrystic orthoclase. Along the east side of Isthmus Sound, Rp is characterised by early crystallising euhedral white plagioclase and grey quartz of approximately equal grain size. Large euhedral light pink orthoclase forms crystals up to 5 cm (OU 75122, plate. 4B). Minor straw brown biotite occurs with magnetite, being a late crystallising phase located in the interstices between quartz and plagioclase. Biotite has a Fe/Mg ratio of 1.1 and TiO2 contents of c.3 wt % (microprobe analysis, OU 75107). Massive Rp along the western shore of Long Sound contains similar porphyritic mineralogy, although the orthoclase appears smaller than samples from Isthmus Sound. A notable difference is the crystallizing sequence with biotite forming inclusions in plagioclase indicating co-crystallization (OU 75143, plate. 4C). This is possibly a function of increased water pressure, which has acted to suppress the crystallisation of plagioclase and promote earlier crystallization of biotite. Alkali feldspar megacrysts frequently display simple twins and a micro-perthitic texture with little zonation (Or88-89, microprobe analysis OU 75107; plate. 5A). Smaller grains of interstitial microcline from OU 75107 show more compositional zonation (Or86-95). In OU 75143 the alkali feldspar appears to be late crystallising, and encases euhedral plagioclase and biotite as optically continuous poikilitic crystals (plate. 5F and G). Alkali feldspar from Revolver Bay is especially notable for its deep pink colouration compared to Isthmus Sound where it is lighter in colour. The pink colouration exhibited in some alkali feldspar is attributed to the exsolution of Fe3+ incorporated during crystal growth. Incorporation of Fe3+ also suggests the magma was relatively oxidised, which is supported by the presence of magnetite.

36

Plate 5. Photomicrographs from Rp (OU 75107). (A) Perthitic orthoclase megacryst; (B) Biotite (Bt) clots showing extensive chlorite replacement interstitial between microcline (Mc) and quartz (Qtz); (C) Resorbed plagioclase megacryst; (D)

37 Microcline megacryst. Photomicrographs from Rp (OU 75143) (E) Zoned plagioclase phenocryst; note poikilitic perthitic orthoclase enclosing a euhedral plagioclase crystal in bottom left; (F) Plagioclase phenocryst surrounded by late crystallising perthitic orthoclase and later quartz; G Poikilitic perthitic orthoclase encasing earlier plagioclase.

Plagioclase in OU 75122 close to the contact with Treble Mountain Granite lacks zoning, occurs as megacrysts up to 12 mm, and has a relatively uniform composition of An34 (Michel Levy method). In contrast, plagioclase from OU 75107; a less evolved sample shows limited normal zonation (An15-21, microprobe analysis). OU 75107 also contains plagioclase with resorbed highly altered cores, which are often mantled by fresher euhedral overgrowths (plate. 5C). Many grains have complex growth patterns frequently overgrowing earlier crystals. OU 75143 from Revolver Bay contains a fine grained unfoliated mafic inclusion about 3 cm in length, with disseminated pyrite proximal to the enclave. Smaller mafic enclaves were also observed in Rp along Narrow Bend in Long Sound. Indications are that OU 75143 is an oxidised rock, which is not consistent with the presence of sulphides crystallising in the magma. The sulphides present in OU 75143 are maybe related to the mafic inclusions, which are possibly forming a localised reduced environment. Mafic enclaves could represent assimilated material. The lack of muscovite in Rp is consistent with a weakly peraluminous composition, which is typical of I-type granites. 3.2.2.2

Treble Mountain Granite (tmg)

Treble Mountain Granite (Tmg) forms most of the Treble Mountain Range including Treble Mountain itself and outcrops along the western shoreline of Isthmus Sound. It covers an area of c.15 km2 (see map in back cover). In hand specimen Tmg is dominated by pink medium grained equigranular biotite granite (plate. 4D), although along the western shore of Isthmus Sound it is locally a coarse grained porphyritic granite similar to Rp. Coarse grained granite (OU 75100) is composed of large megacrysts of plagioclase from 5-8 mm, pink perthitic orthoclase megacrysts greater than 10 mm in a fine grained mosaic of microcline, plagioclase, and quartz (OU 75100, plate 4.D). Plagioclase has little zonation (An23-29, microprobe analysis) and often shows advanced sericitization of the cores (plate. 6A). Perthitic orthoclase sometimes shows zonation through to microcline on the rims, with the microcline showing much less exsolution (plate. 6A and B). Biotite is straw brown in colour where it is not altered to chlorite and occurs as clots with accessory magnetite, zircon, and apatite. Microprobe analysis of biotite gives Fe/Mg ratios between 0.81-0.96, and TiO2 contents of 2.4 to 3.3 wt %. Medium grained granite is composed of grains typically between 1 and 5 mm of anhedral quartz and microcline with subhedral plagioclase. Euhedral and subhedral titanite occur with magnetite in mafic clots associated with the original biotite. Euhedral titanite, mostly hosted

38 within orthoclase is particularly abundant in OU 75162, and is inferred to be magmatic in origin (plate. 6C). Brittle deformation and hydrothermal brecciation (OU 75108) becomes more pervasive in the west of the pluton along the Treble Mountain Range. Biotite is completely pseudomorphed by often well formed chlorite. Hydrothermal brecciation fractures are infilled by epidote and zeolites. Plagioclase is generally altered to sericite and epidote, been often more extensively altered in the cores. Such an alteration assemblage is similar to the deeper propylitic zone of porphyry systems.

Plate 6. Photomicrographs of Treble Mountain Granite. (A cross polars, and B plane light) Zoned plagioclase megacryst mantled by microcline on right with single zoned orthoclase crystal at the top left. Note the core of the orthoclase has advanced perthitic exsolution while the rim of microcline shows little exsolution. Straw brown biotite shows some hydrothermal alteration to chlorite (OU 75100). (C). Euhedral magmatic titanite and subhedral plagioclase hosted in late forming orthoclase (OU 75162). (D). Microcline megacryst surrounded by a mosaic of finer grained microcline, plagioclase, quartz, and biotite. Note also the presence of myrmekite.

Cataclastic zones orientated roughly NW-SE were observed to cut the Treble Mountain range in at least three places. At the south end of Treble Mountain (B45: 2021089/5446306) a cataclastic zone oriented NW-SE contains fragments of hydrothermal quartz (no mineralisation observed), diorite, and metasediment. This shear zone is of interest as it is

39 approximately along strike from a similar structure, which contains intrusive diorite as well as containing the Tarawera lode. The western shore of Isthmus Sound lacks the intense brittle deformation of samples from the Treble Mountain Range and are slightly less altered with some remnant fresh straw brown biotite (OU 75100). An intrusive contact occurs on the western side of Isthmus Sound with the Revolver Pluton. The contact is sharp and well defined with little evidence of alteration to either unit. Mafic enclaves and megacrystic plagioclase xenocrysts are contained in Tmg from the contact (OU 75164). The plagioclase xenocrysts are inferred to be ‘mined’ from the Rp during intrusion. 3.2.2.3 Long Scarp Granodiorite (lsg) Biotite granodiorite outcrops on the ridge line overlooking Long Sound directly east of Last Cove (OU 75175, plate. 9A). The granodiorite lacks the pink orthoclase, which is distinctive in both the Revolver Pluton and adjoining Upper Blacklock Granite (Ubg) and thus represents a different mapable unit based on mineralogy, which is here named the Long Scarp Granodiorite (Lsg). Dating of OU 75175 collected from the ridgeline suggests that both Rp and Lsg are coeval. No contact was observed between Lsg and the adjacent Ubg and no mineralogical or textural gradation was observed between the two. This is evidenced by a sample of Ubg (OU 75176) collected only 50 m distant from OU 75175, which is mineralogically and geochemically similar to another Ubg sample collected 400 m distant (OU 75170) (see plate 9D). This suggests that the relationship between Lsg and Ubg is intrusive, which is supported by the significantly younger age of c.125 Ma obtained for Ubg. The lithological extent of Lsg was not defined during fieldwork due to weather constraints. However, another sample of granodiorite collected during Qmap fieldwork (P 70583), c.2.75 km SE of OU 75175 is almost geochemically and mineralogically identical. The almost identical chemistry of OU 75175 and P 70583, combined with their close proximity makes it probable that they are the same unit. The Lsg is dominated by anhedral plagioclase, quartz, and microcline, with strings of strawbrown biotite containing abundant magnetite, apatite, titanite, clinozoisite, and accessory zircon (plate. 7A and B). Quartz occurs as mosaics up to 5mm and has probably been deformed, as a weak alignment is apparent in OU 75175 (plate. 9A). Plagioclase occurs as grains up to 7 mm, which show evidence of complex zonation as well as intergrowth of fragments at different crystallographic orientations. Microcline is typically less than 1 mm, although occasionally encases plagioclase fragments in a poikilitic fashion indicating its later

40 crystallisation (plate. 7C). Magnetite is distinctive in forming skeletal grains up to 2 mm, and its abundance gives the rock a high magnetic susceptibility (qualitatively determined). Plagioclase often has dusty, but not pervasive sericitic alteration, while biotite exhibits little alteration. The lack of pleochroism of euhedral epidote group minerals encased in biotite, suggests it is most likely clinozoisite (plate. 7D). Clinozoisite has been found to be stable at pressures above 20 kbar, while its breakdown has been found to be important in melting tonalite at higher pressures where hornblende and biotite are not stable (Douce, 2005). P70583 has the same major mineral components as OU 75175, although it is a finer grained rock (plate. 9B). Biotite in P 70583 is also more red-brown, as opposed to straw brown biotite in OU 75175. The red-brown biotite possibly reflects slightly more oxidising conditions in P70583. Deformation of quartz and brecciation of feldspar indicate deformation during crystallisation, which may be magmatic in origin.

Plate 7. Photomicrographs of Long Scarp Granodiorite (OU 75175). (A) Interstitial mafic clot of biotite and magnetite in plagioclase and quartz. (B) Enlargement of (A) showing abundant zircon and apatite. (C) Plagioclase encased in late crystallising microcline. (D) Clinozoisite encased in biotite.

41 3.2.2.4 Trevaccoon Diorite (tdi) Trevaccoon Diorite (Tdi) was observed at the north end of the Treble Mountain Range, and along the west shore of Long Sound between Jane and Last Coves, covering an area of c.6 km2 (see map in back cover). Lithologically Tdi ranges from gabbro to diorite, with more gabbroic compositions along the shore of Long Sound and dioritic compositions on the Treble Mountain Range. The gabbroic rock (OU 75154; plate. 9C) is dominated by plagioclase, with subequal proportions of biotite and hornblende. Biotite, hornblende, magnetite and rare magmatic epidote form clots interstitial between plagioclase (plate. 8A and B). Often biotite surrounds hornblende cores, although the sharp grain boundaries of the biotite against the surrounding minerals and excellent crystal form suggests it was a late crystallising phase, rather than the result of subsolidus recrystallisation. Diorite from the Treble Mountain Range (OU 75111) shows a weak foliation indicated by alignment of plagioclase laths and shows a dominance of biotite over green hornblende, which makes up about 25% of the rock. A

B

C

D

Plate 8. Photomicrographs from Tdi (OU 75154). (A) Sericitization of plagioclase cores. (B) Biotite replacing hornblende. Mafic cluster of biotite, hornblende, and magmatic epidote (C) under plain light, and (D) under cross polars).

42 A

B

C

D

E

FF

Plate 9 Long Scarp Granodiorite samples, (A) OU 75175, and (B) P 70583; (C) Trevaccoon Diorite OU 75154; (D) Upper Blacklock Granite OU 75176. Only Island Diorite samples, (E) Dioritic dyke OU 75168, and (F) Massive Oid OU 75171.

Rare epidote has well defined crystal faces in contact with biotite and partially resorbed cores where in contact with plagioclase (plate. 8C and D). The euhedral nature of the epidote against biotite suggests these two minerals were in equilibrium. Schmidt and Poli (2004) note that epidote crystallisation is typically before or contemporaneous with biotite. The embayment of epidote in contact with plagioclase is also characteristic of magmatic epidote indicating that epidote was not stable during the final crystallisation of the magma. This is an important observation, as rocks containing magmatic epidote frequently yield lower pressures using independent geobarometers than what has been experimentally determined for

43 magmatic epidote (Tulloch and Challis, 2000). Schmidt and Poli (2004) suggest that evidence of disequilibrium during final crystallisation is probably indicative of magma ascent subsequent to the crystallisation of the epidote and its slow re-equilibrium with the melt. The dominance of hydrous mafic minerals, especially the presence of hornblende is indicative of an I-type character. The contact between Tdi and Rp in Long Sound is ambiguous as to sequence of intrusion (B45: 2026992/5449898), although the presence of a two dioritic dykes inferred to be Tdi crosscutting Rp is in agreement with a younger age for Tdi. One of these dykes cuts Rp c.2 km south of Last Cove (Turnbull, pers. comm.). The second, a 3 m wide gabbroic dyke (OU 75163) cutting Rp on the western shore of Isthmus Sound, is correlated with Tdi based on almost identical chemistry (see section 10.1.7) to massive Tdi from the western shore of Long Sound (OU 75154). No contact relationships between Tdi and Tmg were observed or could be inferred.

C

D

Plate 10. Photomicrographs of Upper Blacklock Granite (A), Mosaic quartz adjacent to a microcline megacryst in OU 75170. (B), Quartz mosaic surrounding seriticised plagioclase and biotite altered to chlorite. Only Island Diorite OU 75171 (C), intergrown euhedral plagioclase with euhedral titanite (Ttn) often enclosed in poikilitic hornblende. (D), Large plate of poikilitic hornblende.

44 3.2.2.5 Upper Blacklock Granite (ubg) Upper Blacklock Granite (Ubg) is identified as a separate lithology to Rp due to a significantly younger age (refer to section 4.4.4) and subtle differences in geochemistry (refer to section 5.4.2). Samples were collected from around the most easterly tarn (B45: 2034400/5449700) on the tops separating Blacklock Stream and Long Sound. The northeast contact is inferred where an obscured transition to Long Scarp Granodiorite occurs as discussed in section 3.2.2.3. The southwest contact was not observed but Ubg is presumed to intrude Rp based on its younger age. A lineament running N-S just west of 1021 m may reflect the contact between Ubg and Rp, although this would only be confirmed by more fieldwork. Geochemical correlation (section 5.4.2) of the several samples suggests a pluton of at least c.4 km2. The Ubg (OU 75170 and OU 75176. plate. 9D) is a medium to coarse grained granite composed of grey quartz up to 10 mm; commonly as mosaics (plate. 10A and B), white plagioclase up to 10 mm and pink orthoclase up to 20 mm (plate. 10A). Isolated black biotite, commonly showing complete replacement by chlorite (plate. 10B), occurs interstitially between the feldspars and quartz as grains generally no larger than 3 mm. Orthoclase in Ubg is light pink as compared to the dark pink to red orthoclase found in the Rp from Long Sound (OU 75143), which gives Ubg a more leucocratic colouring. 3.2.2.6 Only Island Diorite (oid) During Qmap fieldwork many undifferentiated mafic bodies of varying sizes were mapped (see map in back cover). It was of interest to compare these undifferentiated mafic bodies with the Tdi to see if they were related. As a result a larger body covering an area of approximately 5 km2, which forms Only Island in Long Sound and outcrops to just south of spot height 1020 m in the head of Blacklock Stream was examined. The mafic body is here informally named the Only Island Diorite (Oid), while the type locality is here suggested to be hornblende gabbro at spot height 1020m in the head of Blacklock Stream where a sample dated in this study yielded a U-Pb zircon age of 122.1 ± 0.9 Ma. Another possibility would be Only Island, although this locality was not visited as part of fieldwork for this thesis. Only Island is composed of diorite (Turnbull and Allibone, pers. comm), although lithological variation seen in the head of Blacklock Stream would suggest that the gabbro at 1020 m is probably the same unit as the diorite forming Only Island. In outcrop Oid is a massive medium to fine grained hornblende diorite to gabbro, suggesting that it was possibly intruded at hypabyssal crustal levels. Cumulate gabbro (OU 75171) from

45 the head of Blacklock Stream is representative of this unit (plate. 9F). Euhedral, variably zoned plagioclase laths of c.1 mm form c.50% of the rock. The intergrown nature of plagioclase laths is indicative of gravity settling (plate. 10C). Plagioclase is encased by large poikilitic clots (up to c.10mm) of hornblende (plate. 10D) containing lesser amounts of yellow-brown biotite and abundant accessory euhedral titanite, magnetite and epidote and rarer zircon. More evolved dykes such as OU 75169 are much finer grained quartz-bearing diorites that contain biotite in approximately equal proportion to hornblende. One intrusive relationship was observed, where a fine grained flow banded biotite-hornblende diorite dyke (OU 75169, plate. 9E), inferred to be a more evolved late stage dyke of Oid, crosscuts massive Ubg at GR. 2034493/5449454. The younger age suggested for Oid by this observation is consistent with dating of both units. 3.2.2.7 Minor Intrusives and Dyke Rocks Gabbroic to dioritic rocks intrude metasediments in Brokenshore Bay (OU 75131, B45:2020832/5443011), Cuttle Cove; referred to herein as the Cuttle Cove Gabbro (OU 75140, B45:2019868/5441747), and Single Tree Island (OU 75138, B45:2019942/5441361). In Brokenshore Bay the diorite is difficult to distinguish from the host argillite metasediment due to grainsize coarsening of the argillite from contact metamorphism. Intrusive margins are irregular in nature making the orientation of the intrusion difficult to distinguish. An intrusive contact exists between the dioritic intrusion and a medium grained granitic intrusion that forms an irregular contact surface. Inclusions of Rp within the medium grained granitic rock indicate it was post Rp, while enclaves of mafic material in the medium grained granitic rock suggest the dioritic intrusion was post Rp, but before the medium grained granitic rock. A gabbroic stock intrudes metasediment in Cuttle Cove where it outcrops along the foreshore for c.50m. Gabbro intruding metasediment on Single Tree Island is probably part of the same intrusion, where diorite forms a complicated partial assimilation texture with host metasediment, which contains abundant sulphide.

46

47

Chapter Four

4 GEOCHRONOLOGY This chapter presents the results from U-Pb zircon LA-ICP-MS dating of Paleozoic and Mesozoic units in southwest Fiordland. In total nine ages were obtained from eight units. All units dated gave well resolved ages with acceptable statistical errors. A new age from the east of Big Pluton gives the first well resolved age on this unit. Two new ages for the Tine Peak Tonalite are within error of each other and confirm a Paleozoic age. Five new Cretaceous ages were obtained from the previously undated Long Scarp Granodiorite, Treble Mountain Granite, Trevaccoon Diorite, Upper Blacklock Granite, and Only Island Diorite. In addition dating of Revolver Pluton gave an age in agreement with a previously reported age in Mortimer et al. (1999b). The LA-ICP-MS spot dating method has allowed inherited populations to be isolated in both Paleozoic and Mesozoic units, which gives important clues to their petrogenesis.

4.1 ELA-ICP-MS methodology Zircon separation was conducted in the mineral separation laboratory, Department of Geology, University of Otago. The samples were first crushed using intermittent bursts on a tungsten-carbide mill and sieved between bursts to isolate material < 250 µm. Particular attention was paid to cleaning the mill between samples. The finest fraction was removed by mixing the sample with water and using a water venturi to remove the suspended fines. Complete drying of the samples was achieved using a drying oven at 700C for at least 24 hrs. Heavy liquids LST (relative density 2.8) and methlene iodide (relative density 3.3) were used in sequence to isolate the heavy mineral fraction. Magnetite was then separated from the samples using a hand magnet covered with a plastic bag which allowed zircons to be picked by hand from the heavy mineral fraction in all case except Treble Mountain Granite where additional separation was required using a Frantz separator, at GNS, Dunedin. Zircons from each rock sample were mounted on double adhesive tape on a glass base. Strips were cut carefully in the backing and folded back to reveal an elongate window of adhesive tape. The zircons were placed individually on the exposed tape using a dissecting probe with a sharpened steel tip under a binocular microscope under 4 x power. Once approximately 50-60 grains were mounted the backing was carefully folded back and taped down to avoid contaminating samples. TEMORA standard zircons and chips of NIST 610 glass were mounted in a similar fashion. A 25mm nylon cylinder was placed around the zircons once the

48 backing had been carefully removed and into which epoxy was poured. The mount was then carefully ground on 600 grit wet and dry sandpaper removing c.20 µm of material to expose the zircon cores. Once exposed, the mounts were then polished on a mechanised lap with aluminium powder by Mr. Brent Pooley. Backscatter and cathodoluminescence images were taken on the JEOL JXA-8600 microprobe in the Department of Geology, University of Otago under direction from Dr. Lorraine Paterson. Zircons were also photographed under planepolarised transmitted light to help avoid inclusions and cracks within the grains during target selection. Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) followed procedures first described by Ballard et al. (2001) and more recently and extensively by Bryan et al. (2004) using facilities in the Research School of Earth Sciences at the Australian National University. Laser ablation was conducted with a pulsed LambdaPhysik LPX 120I UV ArF excimer laser operated at a constant energy of 70 mJ, at 5 Hz using a spot diameter of 32 µm. Ablated material was carried by a mixed He-Ar gas from a custom-designed sample cell and flow homogeniser to an Agilent 7500 ICP-MS. Data for 18 mass peaks were collected in time-resolved mode with one point per peak. Because of a relatively high background, 235

204

Pb was not included. Integration times were 40 ms for

U and 238U, 5 ms for

29

La,

140

208

Ce,

Sm,

153

147

Pb,

Dy, 175Lu, and 177Hf, and 20 ms for 49Ti for a total scan time of 396 ms. Background data

Y,

139

Pb,

163

P,

89

207

Th,

Zr, 10 ms for

31

Pb,

Hg

232

Si and

91

206

204

Eu,

were acquired for 20 s followed by 40 s with the laser on giving about 100 mass scans and a penetration depth of approximately 20 µm. A purge time of 30-60 s is allowed between each spot analysis to permit a return to background signal levels. Background laser off signals were subtracted and outliers (defined at > 3 standard deviations) were excluded. Custom software was used to reduce the raw ICP-MS data offline. Correction factors were applied to each spot analysis to correct simultaneously for instrumental mass bias and ablation-related elemental fractionation. After triggering, it takes several mass scans for the ion counts to reach a smooth signal, so routinely the first 10 scans were excluded. Ablation related inter-element fractionation of Pb, Th, and U were corrected by reference to standard zircon TEMORA and NIST610 silicate glass. Measured 207Pb/206Pb, 206Pb/238U, and 208

Pb/232Th ratios in TEMORA and

232

Th/238U in NIST610 were averaged over the course of

each analytical session and used to calculate correction factors based on accepted values (Pearce et al., 1997, Black et al., 2003).

49 Ages older than 800 Ma were calculated from 207Pb*/206Pb* ratios whereas younger ages were based on 206Pb*/238U ratios (where * indicates radiogenic Pb, the total Pb minus common Pb). Common Pb corrections were applied based on the difference between measured and expected

208

Pb/206Pb or

207

Pb/206Pb ratios for the measured

232

Th/238U value, according to

methods described by Compston et al. (1984). Internal errors for individual analyses were calculated from the observed variation in the corrected isotope ratios for each mass scan over the data interval selected for age calculation. Analyses with resolvable age zonation or disturbance exhibit internal MSWD (mean squared weighted deviates, calculated using counting statistics based errors for each mass scan) values in excess of 3 and, in many cases, were rejected from further consideration.

Concordance was calculated on the basis of

agreement between 207Pb*/235U and 206Pb*/238U ages. Pooled

206

Pb*/238U ages and associated errors include a consideration of within session

analytical error, external uncertainties in the analysis, and age calibration of the zircon standard, as outlined by Stern and Amelin (2003). Within-session error determines these and was estimated using the standard deviation of

206

Pb/238U and

207

Pb/206Pb measurements of

NIST610 silicate glass. External errors incorporated only the uncertainty in the analysis of TEMORA, here taken as the standard error in the

206

Pb/238U measurements, as the reported

error in the age is negligible (Black et al., 2003). The typical within session errors were c.1% for

206

Pb/238U and c.2% for 207Pb/206Pb based on

15-20 analyses of TEMORA with each session. As a measure of the reproducibility of the LA-ICP-MS method, 85-521, an in-house reference zircon was also measured. The resulting (208Pb-corrected) 206Pb*/238U age was 42.9 ± 0.3 Ma (N=15 of 16, MSWD = 1.25) which falls within error of the long term average result of 42.3 ± 0.4 Ma reported in Kesler et al. (2005) and the independent SHRIMP determination given by Ballard (2001).

4.2 Paleozoic geochronology 4.2.1 Tine Peak Tonalite Zircons from two Tpt samples were dated. The first sample, OU 75184, has only moderate enrichment of Zr (288 ppm). The second sample, OU 75193 has extreme enrichment in Zr (565 ppm) and was dated subsequent to OU75184, because it was suspected to contain abundant inheritance, which would have provided clues to the genesis. The two samples are lithologically similar, both being coarse grained two-mica garnet tonalites.

50 4.2.1.1 Tine Peak Tonalite (OU 75184) Sample OU 75184 is from a tributary on the west side of Lake Monk. Of 34 grains mounted, 26 analyses were made on 24 grains. Data are listed in appendix 10.4.2 and isotopic ratios uncorrected for common-Pb are plotted on a Tera-Wasserburg diagram in figure 7. One analysis with a younger age of c.330 Ma was excluded due to probable Pb-loss while the core of the same grain gave a Grenvillian age which was also not included in the age interval. A 208

Pb-based correction was applied, although common-Pb accounts for no more than 0.7 % of

total 206Pb. It is apparent from the probability density plot in Figure 7 that the analyses have a bi-modal age distribution. The younger population includes 14 analyses and gives an age of 347.6 ± 2.9 (2σ) Ma with a MSWD of 1.84. The older population forms a statistically viable population of 10 analyses, which gives pooled 206*Pb/238U age of 365.8 ± 3.5 (2σ) Ma with a MSWD of 1.73.

Figure 6 Cathodoluminescence images of selected Tine Peak Tonalite zircons from sample OU75174.

51 0.08

Tine Peak Tonalite (OU 75184) 347.6 ± 2.9 Ma MSWD 1.84 (2σ) 14 of 25

3

0.07

2

0

0.06

207

Pb/206Pb

N 1

400 390 380 370 360 350 340 330 320 310 300

date (Ma)

410

390

370

350

330

0.05

0.04 15

16

17 238

18

U/

206

20

Pb

0.08

Tine Peak Tonalite (OU75193) 352.3 ± 2.3 Ma MSWD 1.76 (2σ) 33 of 39

8

0.07

19

7 6 5

N

3 2 1

0.06

207

Pb/206Pb

4

0 400 390 380 370 360 350 340 330 320 310 300

date (Ma)

410

390

370

350

330

0.05

0.04 15

16

17 238

18

U/

206

19

20

Pb

Figure 7. Common-Pb uncorrected Tera Wasserburg concordia diagrams for Tine Peak Tonalite samples OU 75184 and OU 75193. Red elipses are the primary magmatic population, blue are older inherited population, and black are analyses excluded due to Pb-loss and high MSWD’s. (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those included for age calculation and unfilled columns being excluded analyses. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red guassian curve is the selected primary population, and the blue curve an older inherited population.

52

4.2.1.2 Tine Peak Tonalite (OU 75193) Sample OU 75193 was collected from c.0.5 km west of Rugged Mount and 3 km north of the first Tpt sample dated (OU 75184). Of the 45 grains mounted, 39 analyses were made on 36 grains. Data are listed in appendix 10.4.3 and isotopic ratios uncorrected for common-Pb are plotted on a Tera-Wasserburg diagram in figure 7. Core and rim analyses were made on 4 grains that were suspected of having inherited cores. In contrast to OU 75184, a single statistically resolvable population is exhibited by OU 75193 albeit with larger amounts of common-Pb. Of the 39 remaining analyses, 33 form a statistically viable population after 208

Pb correction for common-Pb. The pooled

206*

Pb/238U age of 352.3 ± 2.3 (2σ) Ma has an

MSWD of 1.76. Of the 6 excluded analyses, 3 are excluded on the basis of inferred Pb-loss, while 2 are excluded because of high internal MSWD (> 10). One grain with an age of 372 ± 4 Ma falls outside the main population and probably represents subtle inheritance. The age resolution from LA-ICP-MS for Paleozoic zircons would not allow a confident separation of OU 75184 and OU 75193 based on their calculated ages and thus the two ages of 348 ± 3 Ma and 352 ± 2 Ma are considered coeval.

4.2.2 Big Pluton (P 70791) Sample P70791 is from near Mt Aiken (GR B46: 2047529/5439055), c.1.5 km SW of Solitary Peak where Muir et al. (1998) previously attempted to date Big Pluton (sample dated was KAK1, referred to as Kakapo Granite at the time of publication). The zircon population is dominated by elongate prisms between 100 and 150 μm with aspect ratios from 3:1 to 5:1. Most grains display concentric magmatic zonation, with occasional complexity in the cores. Of 67 mounted grains, 50 analyses were made on 43 grains. Data are listed in appendix 10.4.3 and isotopic ratios uncorrected for common-Pb are plotted on a Tera-Wasserburg diagram in figure 8. A

208

Pb based correction was applied, although common Pb accounts for no more

than 1% of total 206Pb. One zircon gave a Grenvillian

207

Pb/206Pb age of 1810 ± 12 Ma; although its rim gave a

magmatic age of c.365 Ma and another yielded an age of 411 Ma. Of the 48 analyses within the selected age interval, 33 form a statistically viable population with a pooled

206*

Pb/238U

age of 366.5 ± 2.2 (2σ) Ma, with a MSWD of 1.80. Rim analyses of 6 zircons give slightly younger ages than their respective cores. The average age for the cores is c.370 Ma; while the average age for the rims is c.358 Ma. As the rim ages show scatter between 347 and 378 Ma, it is likely that Pb-loss has occurred from the rims. On the basis of this hypothesis all analyses were pooled for age calculation and the calculated age of 367 Ma is probably close to the true

53 crystallisation age. Of the analyses not included in the main population; 3 are excluded on the basis of Pb-loss, 5 have high internal MSWD, and the remaining seven are inherited with a pooled age of 386.4 ± 2.4 (2σ) Ma and MSWD of 0.40. 0.08

Big Pluton (P 70791) 367.6 ± 2.2 Ma MSWD 1.77 (2σ) 33 of 48

8

0.07

7 6 5 4

N

2 1 0 400 390 380 370 360 350 340 330 320 310 300

date (Ma)

0.06

207

Pb/

206

Pb

3

410

390

370

350

330

0.05

0.04 15

16

17 238

18

U/

206

19

20

Pb

Figure 8. Common-Pb uncorrected data for Big Pluton (P 70791) plotted on a Tera Wasserburg concordia diagram. Red ellipses are the primary magmatic population, blue are an older inherited population, and black are grains excluded due to Pbloss and high MSWD. (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those analyses included for age calculation and unfilled columns being excluded analyses. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red Gaussian curve is the selected primary population.

4.2.3 Discussion Table 2 presents a summary of dating results for Tine Peak Tonalite and Big Pluton. A lack of significant Gondwanan inheritance is notable in the two Tine Tonalite samples and Big Pluton sample, which suggests early Paleozoic metasediments that comprise the crustal superstructure in the region are not involved in their genesis. This contrasts with the abundance inheritance in the Karamea Batholith (Muir et al., 1994). This would suggest that the elevated Zr in Tine Peak Tonalite is not due to inherited zircon and is of magmatic origin. Zr saturation in a melt has been found to be most dependent on temperature, which implies that Tine Peak Tonalite is a high temperature melt. The coeval ages of 348 ± 3 Ma and 352 ± 2 Ma obtained for the two Tine Peak Tonalite samples combine to give an age of 350 ± 2 Ma for the unit. This age is agreement with an unpublished TIMS U-Pb zircon age of c.348 Ma for Tine Peak Tonalite obtained by Tulloch (pers. comm). A well resolved inheritance

54 component with an age of c.366 Ma forms c.40 % of analyses in OU 75184. The abundance of c.366 Ma zircons is unlikely to be accidental in origin and most likely originated from the source of Tine Peak Tonalite. As an age of c.366 Ma is significantly younger than most metasedimentary source components of the Western Province, it likely the inherited zircons in Tine Peak Tonalite come from reworking of earlier plutons. The 368 Ma age obtained for the Big Pluton sample confirms the existence of plutons coeval with the inheritance seen in Tine Peak Tonalite, although zircon trace element chemistry would appear to suggest that it was not Big Pluton itself that was the source of the inheritance (see section 5.3.2.4). Table 2. Summary table of zircon U-Pb age results for Tine Peak Tonalite and Big Pluton. Uncertainties are 2σ, based on number of analyses in brackets with population MSWD indicated.

Sample

384-390 Ma

372 Ma

366 Ma

350 Ma

Others

Tpt OU75184

365.8 ± 3.5

347.6 ± 2.9

1 Pb-loss

(25 in interval)

(10)

(14)

MSWD

1.73

1.84

(# analyses)

Tpt OU75173

352.3 ± 2.2

4 Pb-loss

(39 in interval)

(31)

MSWD

1.76

3 high MSWD

Bp P70791 (48 in interval) MSWD

372 (1)

386.4 ± 2.4

366.5 ± 2.2

3 Pb-loss

(7)

(33)

0.40

1.80

5 high MSWD

The sample of Big Pluton east dated in this study gives a well resolved age of 366.5 ± 2.2 Ma. Dating of BpW by Ewing (2003) and Muir et al. (1998) yielded significantly younger ages of c.356 and c.340Ma respectively. The age of c.340 Ma is mostly speculative as the analyses are spread from c.320 – 390 Ma with no distinct single population, and as Muir et al. highlights, an age of c.370 Ma is just as likely. An older age for BpW is indicated by Ewing’s (2003) data, which also included a reanalysis of Muir et al’s (1998) sample (KAK1). The majority of individual analyses range in age from c.350 - 380 Ma. The pooled ages of all grains for two samples give ages of 363 ± 9 Ma and 366 ± 8 Ma, while a third sample gives an age of 355 ± 9 Ma. Although the age is not well resolved for BpW, there is a hint that the true age is between 360 and 370 Ma. The average of the means would suggest that BpW is younger than BpE, although with the spread of data and high errors involved no certainty can be attached to this conclusion.

55 A dearth of published zircon ages in southwest Fiordland makes correlation of units in the region difficult. Big Pluton has an age similar to that obtained for the Mt. Solitary Granodiorite of 371 ± 3 Ma by Davids (1999), which suggests that more plutons of this age exist in southwest Fiordland. Unpublished ages for Evans Pluton, Widgeon Granitoid Gneiss, and Newton River Granodiorite by Tulloch (pers. comm.) suggest that they are all c.340 Ma. It is apparent that a period of widespread plutonism occurred in southwest Fiordland in the Carboniferous.

0.08

5 1

2

N

110

115

120

125

130

135

140

145

150

155

160

0

0.06

date (Ma)

207

Pb/206Pb

3

4

0.07

6

Long Scarp Granodiorite (OU 75175) 133.1 ± 0.8 Ma MSWD 1.40 (2σ) 20 of 32

0.05

150

140

130

120

0.04 40

44

48 238

52

56

U/206Pb

Figure 9. Common-Pb uncorrected data for Long Scarp Granodiorite (OU 75175) plotted on a Tera Wasserburg concordia diagram. Red elipses are the primary magmatic population, blue are an older inherited population, purple are two older inherited grains, and black are grains excluded due to probable Pb-loss. (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those analyses included for age calculation and unfilled columns being excluded analyses. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red Guassian curve is the selected primary population, and the blue curve an older inherited population.

4.3 Mesozoic Geochronology 4.3.1 Long Scarp Granodiorite (OU 75175) Sample OU 75175 is a biotite granodiorite from the ridgeline north of Blacklock Stream. Zircons were typically about 100 μm in length with varying morphology from elongate prisms with aspect ratios of 3:1 to more stubby prisms with aspect ratios of about 2:1 (Fig.10). Growth zoning is commonly concentric, although occasional grains show more complexity. Thirty two analyses from 32 grains were taken from a total of 97 mounted grains. Data are listed in appendix 10.4.4 and isotopic ratios uncorrected for common-Pb are plotted on a

56 Tera-Wasserburg diagram in figure 9. A

208

Pb based correction was applied, although

common Pb accounts for no more than 1% of 206Pb in most grains (up to 2.7 %).

Figure 10 Cathodoluminescence images of selected Long Scarp Granodiorite zircons. Primary are those zircons that yielded analyses that are included in the calculated age. Inherited are older analyses interpreted as inheritance.

The largest single statistically resolvable population incorporates 20 analyses and gives a pooled

206*

Pb/238U age of 133.1 ± 0.8 (2σ) Ma with an MSWD of 1.40. Seven analyses are

excluded on the basis of probable Pb-loss, which give apparent ages down to 121 Ma. Five older grains are interpreted as inherited; of which three give an age of c.138.5 ± 1.5 Ma, one gives a concordant age of c.142 Ma, while the last gives a discordant age of c.149 Ma.

4.3.2 Revolver Pluton (OU 75143) Sample OU 75143 is a coarse grained biotite granite with distinctive pink orthoclase megacrysts from the eastern shore of Revolver Bay. Zircons typically display concentric magmatic zonation with aspect ratios from 3:1 to 2:1 (fig.11). Most grains are between 100 and 150 μm. Twenty eight analyses were taken from 28 grains from a total of 43 mounted. Data are listed in appendix 10.4.7 and isotopic ratios uncorrected for common-Pb are plotted on a Tera-Wasserburg diagram in figure 12. A common Pb accounts for less than 1% of total

208

206

Pb based correction was applied, although

Pb in most grains (up to 1.3%).

57

Figure 11. Cathodoluminescence images of representative primary magmatic zircon from Revolver Pluton with visible ablation pits and grains from a secondary inherited population. Note position of inherited ablation holes appears to coincide with a visible core. 0.08

5

Revolver Pluton (OU 75143) 132.4 ± 1.0 Ma MSWD 1.69 (2σ) 17 of 28 3

4

0.07

2 1 0 110

115

120

125

130

135

140

145

150

155

160

0.06

date (Ma)

207

Pb/206Pb

N

0.05

150

140

130

120

0.04 40

44

48 238

U/

52 206

56

Pb

Figure 12. Common-Pb uncorrected data for Revolver Pluton (OU 75143) plotted on a Tera Wasserburg concordia diagram. Red ellipses are analyses included used in calculation of the magmatic age, blue are an older inherited population, and black are analyses excluded due to inferred Pb-loss. (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those considered for age calculation and unfilled columns being excluded analyses. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red Guassian curve is the selected primary population, and the blue curve an older inherited population.

58 Three analyses are excluded as they are inferred to have suffered Pb-loss. The youngest single statistically viable population incorporates 17 analyses, which give a pooled 206*Pb/238U age of 132.4 ± 1.0 (2σ) Ma and MSWD of 1.69 (fig.12). An older population includes eight of the remaining analyses, which give an age of 138.0 ± 1.3 (2σ) Ma with an MSWD of 1.25. High resolution cathodoluminescence images of grains and ablation pits in figure 11 confirm the existence of inherited cores.

Figure 13. Cathodoluminescence images of Treble Mountain Granite zircons analysed for dating. Primary zircons are grains from which analyses were obtained that were used to calculate the magmatic age. Inherited zircons give ages older than the magmatic population.

4.3.3 Treble Mountain Granite (OU 75108) Sample OU 75108 is hydrothermally altered-medium grained biotite granite from the summit of Treble Mountain. Grains are of consistent morphology being squat prismatic crystals seldom larger than 100 µm (fig.13). Grains mostly display concentric growth zonation with low-U rims. Of the 67 grains mounted, 29 analyses were made on 29 grains. Data are listed in appendix 10.4.8 and isotopic ratios uncorrected for common-Pb are plotted on a TeraWasserburg diagram in figure 14. A 208Pb-based correction was applied, although common-Pb accounts for no more than 1% of total

206

Pb (maximum 2.4 %) Four of the analyses are

excluded due to probable Pb-loss. The youngest statistically viable population includes 20 analyses, which give a pooled 206*Pb/238U age of 130.4 ± 0.9 (2σ) Ma with an MSWD of 1.64. The remaining 5 analyses are considered inherited. Three of the inherited analyses give a pooled age of 134.2 ± 0.2 Ma. Of the two remaining analyses, one gives a discordant age of c.138 Ma, and the other a concordant age of c.140 Ma. On the basis of these results, further

59 LA-ICP-MS analyses were made on zircons from Tmg that exhibited petrographic and cathodoluminescence image evidence for older cores. These results are presented in section 4.4.2.

Treble Mountain Granite (OU 75108) 130.4 ± 0.9 Ma MSWD 1.64 (2σ) 20 of 29 6

0.08

0

1

2

N

160

155

150

145 140

0.06

135

130 125

120

115

110

date (Ma)

207

Pb/206Pb

3

4

5

0.07

0.05

150

140

130

120

0.04 40

44

48 238

U/

52 206

56

Pb

Figure 14. Common-Pb uncorrected data for Treble Mountain Granite (OU 75143) plotted on a Tera Wasserburg concordia diagram. Red ellipses are the primary magmatic population, blue and purple are groupings of older inherited analyses, and black are analyses excluded due to probable Pb-loss. (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those considered for age calculation and unfilled columns being excluded analyses. Black are ages from a core study showing subtle inheritance. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red Gaussian curve is the selected primary population, and the blue curve an older inherited population.

4.3.4 Trevaccoon Diorite (OU 75154) Sample OU 75154 is a fine grained hornblende-biotite gabbro from the western shore of Long Sound, about 3 km south of Trevaccoon Head. Zircon morphology is highly variable (fig.15), although elongate prisms were the most likely to give resolvable magmatic ages. Elongate prisms have laminar zoning, while broader grains have no systematic zonation and maybe reflect higher U contents. Grains are noted for their anomalously high U contents (mostly between 700 and 8500 ppm), which accounts for their dark cathodoluminescence. Thorium concentrations are extremely high (up to 30000 ppm) so Th/U ratios vary from 1 to 7, which compare with more typical ratios of less than 1 in Rp, Tmg, Lsg, and Ubg. High Th/U ratios reflect late-stage crystallisation of zircon from a fractionated magma as Zr is much more soluble in a mafic magma than in felsic magma. As U and Th are highly incompatible elements, their concentrations in the last melt to crystallise will be strongly amplified. The

60 high U and Th concentrations and high Th/U ratios in Tdi zircon thus support a magmatic origin of the grains distinct from other more felsic units of the PIC.

Figure 15. Cathodoluminescence images showing representative zircons from the Trevaccoon Diorite from which primary magmatic ages were obtained. 0.08

Trevaccoon Diorite (OU 75154) 128.4 ± 0.9 Ma MSWD 1.48 (2σ) note additional 18 of 27

7

datum at 0.155

0.07

6 5 4

N

2 1 0 160 155 150 145 140 135 130 125 120 115 110

date (Ma)

0.06

207

Pb/206Pb

3

0.05

150

140

130

120

0.04 40

44

48

238

52

56

U/206Pb

Figure 16. Common-Pb uncorrected data for Trevaccoon Diorite (OU 75154) plotted on a Tera Wasserburg concordia diagram. Red ellipses are the primary magmatic population, blue is an older inherited analysis, and black are analyses excluded due to Pb-loss and high internal MSWD (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those considered for age calculation and unfilled columns being excluded analyses. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red Gaussian curve is the selected primary population.

61 Of the 67 grains mounted, 28 analyses were made on 28 grains. The exclusion of one Paleozoic analysis of 499 ± 6 Ma left 27 analyses in the considered interval. Data are listed in appendix 10.4.10 and isotopic ratios uncorrected for common-Pb are plotted on a TeraWasserburg diagram in figure 16. As a result of the high Th concentrations the 208Pb method of Compston et al. (1984) for common Pb correction gave erroneous results as the assumption that most

208

daughter of

Pb is from common-Pb no longer applies due to

232

208

Pb also being the radiogenic

Th. For this reason, the analyses of the zircons are corrected for common Pb

using the analogous

207

Pb method. It is suggested that 1000 ppm Th probably represents a

reasonable cutoff for using the preferred 208Pb common Pb correction. Three grains were excluded on the basis of high internal MSWD (> 6). Five more analyses were excluded on the basis of inferred Pb-loss. The largest single statistically viable population incorporates 18 grains, which gives a pooled

206

*Pb/238U age of 128.3 ± 1.0 (2σ)

Ma with a MSWD of 1.48. One grain with an age of 136 ± 2 was considered outside the main population and was excluded.

4.3.5 Upper Blacklock Granite (OU 75170) Sample OU 75170 is from the outlet of a large tarn on the south side of the ridge separating Blacklock Stream from Long Sound. Grain morphology is variable with many grains been poorly prismatic (fig.18). Grains often show evidence of overgrown cores, which is confirmed by a separate dating study, discussed later in section 4.4.2. Zonation is mostly concentric, although cores are often darker and more complex. Rims are typically lighter suggesting the last phase of growth was poor in U. Of 98 mounted grains, 31 analyses were made on 31 grains. The youngest 4 analyses are excluded on the basis of inferred Pb-loss. Data are listed in appendix 10.4.8 and isotopic ratios uncorrected for common-Pb are plotted on a Tera-Wasserburg diagram in figure 17. A 208

Pb based common-Pb correction was applied, although common Pb accounts for no more

than 1% of total

206

Pb in most analyses (up to 4.9 %). The overall distribution curve for all

analyses (fig.17) indicates the presence of two distinct populations. The youngest and preferred magmatic population includes 12 analyses, which give a pooled

206

*Pb/238U age of

126.1 ± 1.0 (2σ) Ma with a MSWD of 1.44. The older population includes 11 analyses, which give a pooled age of 132.3 ± 1.4 (2σ) Ma with a MSWD of 1.72. Four older analyses include three concordant grains with ages of c.137, 142, and 144 Ma and 1 discordant grain with an age of c.138 Ma.

62

Figure 17. Cathodoluminescence images of selected Upper Blacklock Granite zircons with visible ablation pits. 0.10

Upper Blacklock Granite (OU 75170) 126.1 ± 1.0 Ma MSWD 1.44 (2σ) 12 of 31

0.09

6 5 4 3 N 2 1

0.07

0 160 155 150 145 140 135 130 125 120 115 110

date (Ma)

207

Pb/206Pb

0.08

0.06

0.05

150

140

130

120

0.04 40

44

48 238

U/

52 206

56

Pb

Figure 18. Common-Pb uncorrected data for Upper Blacklock Granite (OU 75170) plotted on a Tera Wasserburg concordia diagram. Red ellipses are analyses included in the primary magmatic population, blue are an older inherited population, purple are older inherited grains, and black are analyses excluded due to inferred Pb-loss. (Inset) Probability density plots of

63 208

Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those included for age calculation and unfilled columns being excluded analyses. Black columns are analyses from the core study. (Top) Probability density curves with the thin red line showing distribution of all analyses, while the heavier red Gaussian curve is the selected primary population. Blue and purple curves are an older inherited population.

4.3.6 Only Island Diorite (OU 75171) Sample OU 75171 is a medium-grained hornblende gabbro collected from the ridgeline north of the head of Blacklock Stream. One mineral separation yielded only 9 zircons. Zircons are anhedral (fig.19), which is interpreted to result from late-stage crystallisation. A

208

Pb based

correction for common-Pb was applied to 9 analyses from 9 grains. Common-Pb accounts for up to 3% of total 206Pb. Data are listed in appendix 10.4.11 and isotopic ratios uncorrected for common-Pb are plotted on a Tera-Wasserburg diagram in figure 20. Although the number of analyses is small it is apparent from the age distribution histogram and density plot in figure 20 that two distinct populations exist. The preferred age is based on the 6 youngest grains, which form a statistically viable population, which give a pooled

206

*Pb/238U age of 122.1 ±

0.9 (2σ) Ma with a MSWD of 0.54. The 3 remaining grains also form a statistically viable population and give a pooled age of 127.4 ± 0.7 (2σ) Ma with a MSWD of 0.11. Although the MSWD is low for the two populations, it is within the expected range for a small number of data. The presence of a hornblende-bearing dioritic dyke in Ubg, discussed and interpreted in section 3.2.2.6 evolved Oid, supports the younger age based on the larger population of six analyses.

Figure 19. Cathodoluminescence images of selected Only Island Diorite zircons.

64 0.09

Only Island Diorite (OU 75171) 122.1 ± 0.9 Ma MSWD 0.54 (2σ) 6 of 9

0.08

3

2

N

0.07

0 150 146 142 138 134 130 126 122 118 114 110

date (Ma)

0.06

207

Pb/

206

Pb

1

0.05

150

140

130

120

0.04

0.03 40

44

48 238

U/

52 206

56

Pb

Figure 20. Common-Pb uncorrected data for Only Island Diorite (OU 75171) plotted on a Tera Wasserburg concordia diagram. Red ellipses are the primary magmatic population; while blue are an older inherited population. (Inset) Probability density plots of 208Pb corrected analyses. (Bottom) Histogram of individual analyses with red filled columns being those included in age calculation and unfilled columns being excluded analyses. (Top) Probability density curves with the thin red line showing distribution of all grains, while the heavier red Gaussian curve is the selected primary population, and the blue curve an older inherited population.

4.4 Discussion 4.4.1 PIC Inheritance Initial dating revealed the existence of substantial inherited zircon populations in Rp and Ubg and more subtle inheritance in Lsg, Tmg, and Tdi. Revolver Pluton contains an inherited population of 138.0 ± 1.3 Ma (32% of all analyses, dark grey boxes, Table.3). More than half of the total zircons in Ubg are inherited, with an older population which gives an age of 132.3 ± 1.4 Ma. Coeval inheritance is reflected to a greater or lesser extent in Lsg, Tmg, Tdi, being only absent in Oid. More subtle zircon inheritance of c.142 Ma and c.148 Ma occurs in Lsg and Ubg. Paleozoic zircon inheritance in PIC rocks is noticeably absent, with only one c. 500 Ma grain identified in Tdi, which may have been incorporated during emplacement or could have been an artifact of sample preparation.

4.4.2 Zircon core study No systematic zircon core-rim analyses were undertaken in the initial dating of Rp, Lsg, Tmg and Ubg, which made it difficult to identify clear correlations between ages interpreted as inherited and position of spot in relation to any potential overgrown core. Little evidence was

65 found in the low resolution cathodoluminescence (CL) images of the zircons for a correlation between older ages and grains where an inherited core could be observed. However, some zircons did show evidence of cores with magmatic overgrowth. Zircons from Tmg and Ubg showed the best evidence of resorbed cores and were selected for further study. These cores were dated in a second LA-ICP-MS session to see if core ages correlated with petrographic evidence for inferred inheritance. After the core analyses were done, high resolution CL images were taken of grains in Tmg and Ubg (figs. 21 and 22). A much clearer correlation between inherited ages and evidence of overgrown cores is apparent compared to the lower resolution CL images taken prior to ablation. The ages obtained from the core study are summarised in Table 3 (light grey rows) and are shown as black filled columns on the age distribution histograms in figures 14 and 18. 4.4.2.1 Upper Blacklock Granite Of the eight cores targeted in Ubg all yielded ages coeval with previously identified inheritance, which increased the c.132 Ma population from 11 to 14 analyses, the c.137 Ma from 2 to 3 analyses, and the c.142 Ma population from 2 to 4 analyses, and added one c.147 Ma age analysis. Figure 21 shows high resolution CL images of Ubg zircons after ablation and reveals a clear correlation between inherited ages and overgrown cores. One grain where a partially resorbed core can be identified gives an age of 147 Ma for the core, while the rim has an age of 123 Ma within the magmatic age population interval. It seems unequivocal that Ubg contains inherited cores with distinct inheritance of c.132, c.137, c.142, and c.147 Ma. 4.4.2.2 Treble Mountain Granite Tmg cores contain similar inheritance to Ubg. Of the six resolvable Tmg analyses, 1 analysis gives an age of 128 Ma; three analyses give a pooled age of 134 ± 0.5 Ma, while two older analyses give ages of c.142 and c.147 Ma. Again the high resolution CL images show that inherited ages come from spots positioned in overgrown cores (fig.22). Magmatic overgrowth is most evident in one grain where a resorbed core gives an inherited age of 142 Ma, while the rim has an age of 132 Ma, which falls in the age interval of analyses used to calculate the magmatic age.

142 (1)

147 (1)

0.11

127.4 ± 0.7

130 (1)

MSWD

133 ± 2 (3)

1.72

(11)

132.3 ± 1.4

1.48

(18)

128.4 ± 0.9

128 (1)

(3)

137 (1)

138 (1)

137 (1)

134 ± 0.5 (3)

1.64

(20)

130.4 ± 0.9

127-130 Ma

(9 in interval)

Oid

Ubg Core (8)

142 (2)

144 (1)

(31 in interval)

MSWD

142 (1)

Ubg

discord

(1)

(27 in interval)

MSWD

136

Tdi

Tmg Core (6)

MSWD

(29 in interval) (3)

134.2 ± 0.2

138 (1) 140 (1)

Tmg

1.69

1.25

(17)

132.4 ± 1.0

MSWD

138.0 ± 1.3

1.40

(20)

133.1 ± 0.8

132-135 Ma

(8)

147 (1)

concord

(3)

138.5 ± 1.5

136-140 Ma

(28 in interval)

Rp

MSWD

discord

(1)

(32 in interval)

(1)

142

149

Lsg

(# analyses)

142 Ma

146-149 Ma

Sample

1.44

(12)

126.1 ± 1.0

125 Ma

0.54

(6)

122.1 ± 0.9

122 Ma

4 Pb-loss

1 Paleozoic

3 high MSWD

5 Pb-loss

4 Pb-loss

3 Pb-loss

7 Pb-loss

Others

Table 3. Ages are based on pooled 206Pb/238U dates corrected for common Pb using the 208Pb method. All errors are 2 of the mean and include a contribution from measurement uncertainty of the TEMORA zircon standard for the analytical session. MSWD for populations of 3 or more grains is reported below number of constituent grains.

66

67

Figure 21. Cathodoluminescence images of selected ablated grains from the core study of Ubg. Grains were selected where an overgrown core was thought to be visible in CL after the initial dating run, thus not all grains have a corresponding rim analysis.

Figure 22. Cathodoluminescence images of selected ablated grains from the core study of Tmg. Grains were selected where an overgrown core was thought to be visible in CL after the initial dating run, thus not all grains have a corresponding rim analysis.

68

4.4.3 Zircon thermometry and survival of inherited zircon Zircon saturation thermometry was first discussed in detail by Watson and Harrison (1983). If the inherited zircons described here were from the source rock it would normally be assumed that magma would not be saturated in zircon and that any inherited grains would be dissolved in the magma. King et al. (2001) pointed out that zircon saturation temperatures may not give magmatic temperatures for the following reasons: (1) granite compositions probably do not correspond to their melt composition as they might include significant quantities of crystals either restitic or cumulate; (2) incorporation of inherited zircon; and (3) some melts might not reach zircon saturation. The presence or absence of pre-magmatic zircon was addressed by Chappell et al. (1998) who divided I-type granites into high (> c.800oC) and low-temperature (c.750oC) sub-classes. These workers hypothesised that high-temperature granites lack inherited zircon because magmatic temperatures were high enough to dissolve all of the source components including zircon, where the converse is true for low-temperature granites. This view was supported by Miller et al (2003) who showed that 23 intrusions that contained abundant inheritance gave an average zircon saturation temperature TZr [defined as TZr = 12900/(2.95 + 0.85M + ln(496 000/ZrMELT)) where M = Na + K + 2.Ca/(Al.Si) all in cation fractions] of 772oC, compared to 34 intrusions with little apparent inheritance that yielded an average TZr of 831oC. Modelling of zircon solubility by Watson (1996) found dissolution rates to be most dependant on temperature; but also dependant on the degree of undersaturation, crystal radius, and heating rate. Whether a zircon survives in magma is thus dependant on the initial size of the zircon, the intensity and duration of the magmatic event, and the volume of the local melt reservoir in which the zircon interacts. Miller et al. (2003) recognised two possibilities for presence of older zircon. Firstly, that they were entrained through late stage contamination processes; or secondly, that they are inherited from a deeper level from contributing source material. The first possibility is often invoked to explain the occurrence of inherited zircon cores in rocks where the source compositions should have been undersaturated in zircon. A classic case is the Boggy Plain Pluton of southeastern Australia, which is concentrically zoned from gabbro through to an aplitic core. The initial temperature of the parental melt was estimated by Wyborn et al. (2001) to be 1100oC, well above zircon saturation which is indicated by incompatible behaviour of Zr to c.66 wt % SiO2. Hoskin et al. (2000), however, found that every zircon contained an inherited core in a sample containing only 63 wt % SiO2. The common occurrence of zircon encased in biotite in PIC granites could potentially provide a mechanism where inherited zircons could survive due to lack of contact with the

69 undersaturated magma. This may provide a mechanism for allowing zircon to be incorporated from the source. Importantly though, if dissolution of zircon was inhibited so too would nucleation of new magmatic overgrowth. Thus, as we see magmatic overgrowth of inherited cores in PIC zircons it would be difficult to argue that they were protected from the melt by being encased in biotite. Zircon saturation temperatures of PIC rocks taken from the least evolved samples give temperatures of 812 oC for Tmg, 825 oC for Ubg, c.780 oC for Rp, and 862 oC for Lsg (fig.24). According to the classification of Chappell et al. (1998), Tmg, Ubg, and Lsg are high temperature I-type granites, while Rp is somewhat transitional. The incorporation of abundant inheritance in PIC rocks is enigmatic since according to Chappell et al. (1998) and Miller et al. (2003) there should be little inherited zircon, especially in the case of Lsg, as with only 63 wt % SiO2 it should still be undersaturated in Zr. One explanation could be that incorporation of inherited zircon in PIC rocks gives erroneously high TZr temperatures as warned by King et al. (2001). Rahu Suite granitoids may provide evidence of artificially elevated TZr as they are widely accepted as been contaminated Separation Point magmas that assimilated varying amounts of metasedimentary crust. As a result Rahu Suite rocks are often characterized by abundant zircon inheritance (Muir et al., 1997). As can be seen in figure 24, Rahu Suite rocks on average tend to give higher TZr than Separation Point rocks. Separation Point Granitoid Separation Point Mafic Western Fiordland Orthogneiss Rahu Suite

900 850

T zr

800

Darran Granitoid Darran Mafic

750

Treble Mtn Granite

700

Revolver Pluton Upper Blacklock Granite Trevaccoon Diorite

650 600 40

50

60

SiO2

70

80

Long Scarp Granodiorite Only Island Diorite

Figure 23. Zircon saturation temperature (TZr) plotted against SiO2 for representative Mesozoic plutonic suites in NZ and selected samples of PIC rocks. TZr calculated using formula from Miller et al. (2003).

Another explanation could be that inherited zircon was incorporated after the magma had become saturated in zircon; that is, after the magma starts crystallising zircon. Indeed once the

70 magma starts to crystallise the latent heat of crystallisation may provide additional heat to assimilate portions of the surrounding rock through which it intrudes through widely accepted assimilation + fractional crystallisation (AFC) processes. In this case the inheritance may not directly reflect the source. If PIC rocks incorporated the zircon from wall rocks, it must have been at deeper crustal depths than their present position in the crust because the country rock are Ordovician metasediment and Paleozoic plutons. The only older Mesozoic pluton known in the area is the Lake Mike Granite with an age of 163.7 ± 3.0 Ma, which is older than the observed inheritance in the PIC rocks. An argument against incorporation of zircon through assimilation would be the abundance of inherited zircon observed in PIC rocks; particularly in Rp and Ubg. If the percentage of inherited zircon reflects the percentage of assimilated material (ie. zircon contents of initial magma and contaminant are equal), Rp would be approximately 32 % assimilated material by volume. To assimilate this volume of material would require more heat than that released by the volume of one pluton crystallising and would seemingly require extra heat input from an external source. The most likely source of inherited zircon in PIC rocks is from a source composed of Darran Suite plutons or alternatively assimilation of Darran Suite plutons at mid-crustal levels. The age range of inheritance not only requires that a heterogeneous range of Darran Suite plutons was involved in the source, but in the case of Tmg and Ubg, earlier PIC age rocks must have been reworked. The coeval timing and close spatial relationship of more mafic and more felsic rocks in the PIC, such as Rp and Lsg, and Tmg and Tdi suggest some genetic linkage as does their similar initial isotopic composition (section 5.4.7). It is possible that the mafic correlative could have provided the heat necessary for driving melting of the PIC granite source. Further dscussion of this hypothesis occurs in the following chapters. Incomplete dissolution of zircon in PIC granites possibly indicates one or more of the following: they were originally low temperature melts; the duration and intensity of the magmatic event was too short for complete dissolution; the original size of the zircons was too large for complete dissolution. As Watson (1996) highlights; dissolution of zircon is dependant on more factors than temperature alone, which is necessitated in such examples as the Boggy Plain Pluton.

4.4.4 Summary A clearly resolvable magmatic population can be identified in each of the PIC units dated. Pbloss is evident in all samples except Oid and so the distribution of analyses to younger ages introduces the possibility that the calculated ages could suffer slight negative shifts. However, it is unlikely that a significant shift in age has resulted as Pb-loss would be unlikely to pool the individual ages into statistically viable populations. Older analyses outside the selected

71 magmatic population are interpreted as subtle inheritance, which is supported by two key observations. Firstly, the inherited ages often pool in age groups, which appear to be shared between multiple units. Secondly, grains in Tmg and Ubg where a core could be identified give mostly inherited ages, while the occasional rims of the same grains give magmatic ages. The resolved intrusive ages for analyzed PIC rocks are presented in table 3 in red shaded boxes. Dating highlights that PIC rocks were intruded over a relatively short time interval of c.10 Myr. Rp and Lsg give ages of 132.4 ± 1.0 Ma, and 133.1 ± 0.8 Ma respectively, which are within analytical error of each other. The age of 130.4 ± 0.9 Ma for Tmg is within error of Tdi with an age of 128.4 ± 0.9 Ma. The well resolved 126.1 ± 1.0 Ma age for Ubg makes a direct relationship with the lithologically similar Rp improbable, which supports the distinction of these two units based on their subtle differences in chemistry discussed in the following chapter. Previous work by Muir et al. (1998) found Mesozoic plutonism fell into three distinct phases with the first phase represented by an early burst of Darran Suite plutonism in the late Triassic. A second phase of Darren Suite plutonism occurred in the Mid-Jurassic to the EarlyCretaceous (168-137 Ma), which reinforced earlier dating by Kimbrough et al. (1994). Crosscutting Separation Point plutons in eastern Fiordland had a more limited age range than equivalents in Northwest Nelson with ages from 124 to 121 Ma. The gap between the end of Darran Suite plutonism and the onset of Separation Point plutonism is apparent in the summary of the Median Batholith by Mortimer et al. (1999b). Also evident from dating so far is an apparent lull in Darran Suite plutonism between 146 and 138 Ma. However, more recent work indicates that plutons with ages between 137 and 126 Ma are present and of significant volume (summarised in table 4). Tulloch and Kimbrough (2003) report an age of 128 ±1 Ma on the Supper Cove Orthogneiss about 30 km to the north of the PIC, which they correlate with the Darran Suite. Plutons within this apparent transition period also occur in Northwest Nelson-Westland with the 133.5 ± 0.2 Ma Copperstain Creek granodiorite porphyry (Brathwaite et al., 2004) and the Crow Granite with an age of 137 ± 3 (Muir et al., 1997). Both of these lack Paleozoic inheritance like PIC granites and the Crow Granite has a similar distribution of SHRIMP zircon ages as found in this study. An unpublished age of c.132 Ma has also been obtained for the Rocky Creek Granite in the Victoria Range (pers. comm. Dr Andy Tulloch).

72

Table 4. Age summary of Jurassic to Cretaceous plutonism in New Zealand showing comparison of PIC ages with dated plutons from Fiordland, Stewart Island and Northwest Nelson-Westland.

73

Figure 24. Map showing distribution of known established suites in the South Island of New Zealand. R

Recent dating of plutons from Stewart Island (Allibone and Tulloch, 2004) has found multiple plutons with ages similar to the PIC. These include the c.132-130 Ma North Arm Pluton, c.130 Ma Richards Point Porphyry, c.128 Ma Smoky Pluton, and c.125 Tarpaulin Pluton.

74 Zircons in orthogneisses of the Arthur River Complex in northern Fiordland have been dated by SHRIMP giving an age range for the ARC of 136-129 Ma (Hollis et al., 2003). Figure 24 shows the geographical distribution of known New Zealand Suites in the South Island. LoSY plutons with ages from 137-125 Ma were distinguished on the map to establish their distribution. It is apparent that plutons within this age range are a volumetrically significant component of the magmatic record in the South Island. It is also noted that they seem to be better represented in Fiordland and Stewart Island, although this may be biased due to a less than comprehensive geochronological database at present. Preservation Intrusive Complex dating shows these plutons probably represent a unique suite of rocks spanning the transition period between Darren Suite plutonism and Separation Point plutonism; something hinted at in their geochemistry, which is discussed in the following chapter.

75

Chapter Five

5 IGNEOUS GEOCHEMISTRY 5.1 Introduction This chapter presents geochemical data for Paleozoic and Mesozoic igneous rocks in southwest Fiordland. These will be discussed in separate sections as there is unlikely to be any genetic link over such a wide time gap. Localities of all samples included in this analysis are marked on the map in the back cover, while the geochemical data are presented in the appendices along with source references. When considering the compositions of plutonic suites covering a broad SiO2 range it must be remembered that relationships between mafic and felsic rocks cannot always be explained by one process alone. Often multiple processes may be involved to produce in-suite chemical variation which as highlighted by Chappell (1996) may include: magma mingling and/or mixing; assimilation or contamination; fractional crystallisation, restite separation, hydrothermal alteration, and variation inherited from the source rocks.

5.2 Analytical techniques 5.2.1

Mineral analysis (electron microprobe)

Silicate and sulphide minerals mounted on polished thin sections and briquettes were analysed at the Department of Geology, University of Otago, using a JEOL JXA8600 electron microprobe (EMP) with direction from Dr. Lorraine Paterson. A ‘reference map’ of individual grain locations was prepared using transmitted and reflective light microscopy. Polished thin sections and briquettes were then carbon coated using an Edwards Auto306 carbon coater and tested for conductivity before anaylsis. Quantitative wavelength dispersive spectrometry (WDS) was used for silicate and sulphide analyses. The operating conditions for silicates were 15 kV with a current of 2.0 x 10-8 mA and beam diameter of 20 μm. Conditions for sulphides were 25 kV with a current of 2.0 x 10-8 mA and beam diameter of 5 μm. CITZAF techniques using PAP corrections as defined by John Armstrong (Caltech) were applied to the raw data to obtain oxide proportions. Standards used for silicates were the following: albite (Amelia) for Na2O, adularia (St Gothard) for K2O, rutile (synthetic) for TiO2, corundum (syntetic) for Al2O3, Hematite (Funahozawa) for FeO, wollastonite (synthetic) for CaO, quartz (synthetic) for SiO2, periclase (synthetic) for MgO, and manganese oxide (synthetic) for MnO. Sulphide standards used were Fe metal for Fe, ZnS for Zn, AsS for As, Cu metal for

76 Cu, PbS for Pb, CdS for Cd, and ZnS for S. In addition to quantitative analyses of sulphides, backscatter images were made and identified grains were analysed by EDS to identify potential new phases.

5.2.2

Major and trace element chemistry

5.2.2.1 XRF Major and trace elements (V, Cr, Ni) were determined at the Department of Geology, University of Otago by standard wavelength dispersive XRF techniques using lithium borate fusion discs and pressed powder discs. Rock samples were carefully trimmed of alteration using a rock saw to give samples sizes of approximately 500 g. These samples were further cleaned using a water cooled stone grinding wheel to remove potential contamination from the saw blade. Samples were then dried at 60oC before been reduced to smaller size fractions using hydraulic rock splitter. Rock chips were then ground using a TEMA tungsten carbide mill for 1-2 minutes. The mill head was cleaned thoroughly between samples by first cleaning with hot water, drying with compressed air, cleaning with alcohol, running clean quartz sand, and then repeating the first cleaning steps to avoid contamination. Rock powder was then split into representative sub-samples. Major element fusion discs were prepared with 0.64000 ± 0.00005 g of sample, 6.8000 ± 0.0001 g of lithium borate flux, and about 1.000 g of ammonium nitrate. Trace element fusion discs were prepared using 5 g of sample bined with 5 ml of Moviol binding solution. XRF analyses were performed by Damien Walls using a Phillips PW2400 XRF spectrometer. Loss on ignition (LOI) was calculated to determine H2O content by heating 2.0 g of rock powder to 1100oC for 1.5 hours and then reweighing the sample. 5.2.2.2 LA-ICP-MS Trace element (Ga, Rb, Pb, Sr, Y, Ba, Zr, Cu, Zr, Nb, Hf, Ta, Th, U) and REE (La, Ce, Pr, Nd, Sm, Eu, Gd, Dy, Ho, Er, Tm, Yb, Lu) data was obtained by LA-ICP-MS at the Research School of Earth Sciences, Australian National University under the direction of Dr. Charlotte Allen. The samples analysed were the same glass discs that were used for major element analysis, although they were mounted in the following manner. The discs were broken from which the glass shards were mounted on double sided sticky tape being careful to map sample locations. A 25 mm nylon cylinder was placed around the glass shards and filled with epoxy and cured to create briquettes. The briquettes were then ground with wet and dry sandpaper before being finished by polishing on a mechanised polishing machine using aluminium powder. A ‘reference map’ of these briquettes was carefully made so samples could be easily identified optically in the sample chamber. Each sample was analysed for 40 seconds over

77 which period each element was analysed every 40 ms. Raw data from the ICP-MS was reduced offline using custom software and laser-off backgrounds were subtracted and concentration determined by comparison with the results of NIST 610 glass standard which was analysed in series with the samples and using concentrations recommended by Pearce et al. (1997).

5.2.3 Oxygen isotopes 5.2.3.1 Oxygen (laser assisted fluorination) Oxygen isotope values of silicate samples were obtained Dr Kevin Faure (Institute of Geological Sciences, Lower Hutt, New Zealand) using the laser ablation, BrF fluorination method of Sharp (1990). Values of samples were normalised to the garnet standard UWG-2 (Valley et al., 1995) assuming a δ18O value of 5.8 (VSMOW). An internal quartz standard ("Tubby") with an accepted δ18O value of 11.1 was also run for data quality assurance. Replicate measurements of quartz (n=3) and garnet (n=5) usually differed by less than 0.1 ‰ and always less than 0.2 ‰. 5.2.3.2 Oxygen (conventional) Whole rock analyses were obtained by Prof. Chris Harris (University of Capetown, South Africa). Two NBS28 standards in the same run gave raw values of 9.83 and 9.61 per mil based on the nominal value of the CO2 reference gas. The data obtained was normalised to a δ18O value of 9.64 for NBS. Yields all seemed appropriate for the rock types.

5.3 Paleozoic geochemistry Cretaceous plutonism was the main focus of this study. However, the Tine Peak Tonalite (Tpt) was examined on the periphery of the Cretaceous units. Due to its unique chemistry; a more detailed examination was carried out. My own Tpt analyses are supplemented by three supplied by Qmap Fiordland. A brief geochemical examination of Big Pluton (Bp) is also included to complement the dating of this unit and to establish broader understanding of Paleozoic plutonism in Fiordland. Geochemistry from other southwest Fiordland plutons of confirmed and likely Paleozoic age (Tulloch et al., 2003, Turnbull et al., 2005) are also included for comparison. These plutons include Evans Pluton (Ep), Newton River Granodiorite (Nrg), and Widgeon Granitoid Gneiss (Wgg, included in the Ridge Suite by Tulloch et al., 2003). Indian Island Granite (Iig) geochemistry is also included based on intrusive relationships which suggest a Paleozoic age (Allibone and Turnbull, 2005, In preparation). This interpretation appears correct based on geochemistry. The geochemical

78 analyses of these units were supplied by Drs Ian Turnbull and Andrew Allibone of GNS in advance of their listing on the Petlab database.

5.3.1 Previous work on NZ Paleozoic suites Muir et al. (1996b) in a geochemical study of the Karamea Batholith found that it displayed a continuum of I-type through to mildly S-type compositions and noted that extreme S-type compositions were largely absent in New Zealand when compared with the Chappell and White (1992, 1974) criteria for S-types of the LFB. The lack of more extreme S-type compositions in New Zealand compared to the Lachlan Fold Belt is probably due to a lack of appropriate pelitic sediments in the mid to lower crust of the Western Province.

5

Karamea Suite Ridge Suite

4.5

Tobin Suite

Peraluminous

Foulwind Suite Newton River Granodiorite

3.5

mol Al/(Na+K)

Metaluminous

Paringa Suite

4

Evans Pluton Big Pluton west

3

Big Pluton east Tine Peak Tonalite

2.5

Indian Island Granite

2 1.5 1

Peralkaline

0.5 0 0

0.2

0.4

0.6

0.8

1

1.2

1.4

1.6

ASI Figure 25. Aluminium Saturation Index of Paleozoic southwest Fiordland plutonic rocks. Representative compositions are also plotted from the New Zealand Paleozoic plutonic suites (refer to appendices for sources). ASI is calculated by the following formula: molar Al/(2(Ca-1.67-P)+Na+K). Note formula contains a correction to the formula in Frost (2001) originally defined by Zen (1986).

Tulloch et al. (2003) reported the results of further study of Paleozoic plutonic rocks from throughout the South Island. These authors suggested subdividing a younger S-type Ridge Suite (353-342 Ma) from the older S-type Karamea Suite (382-369 Ma) and recognised two distinct I-Type suites; the Paringa Suite (369-359 Ma), which is slightly younger than the Karamea Suite, and the Tobin Suite (349-340 Ma), which is approximately coeval with the Ridge Suite. Included in their type examples of the Ridge Suite is the c.340 Ma Widgeon Granitoid Gneiss in southwest Fiordland. The S-type Karamea and Ridge suites are uniformly more peraluminous (ASI > 1.1, fig.25) and limited to more felsic compositions (SiO2 > 65 wt %, fig.26) (total alkalis vs. silica, Le Maitre et al., 1989) as compared to the I-type Paringa

79 and Tobin suites. Ridge and Karamea suite compositions are also more potassic and less sodic than the Paringa and Tobin suites reflecting their greater S-type affinities (fig.27). In detail, the younger Ridge Suite exhibits lower K2O and Rb and higher Na2O and Sr than the Karamea Suite and hence a more transitional I-S type character (figs. 27 and 28).

20 Karamea Suite Tobin Suite Foulwind Suite Evans Pluton Big Pluton east Indian Island Granite

18

Na2O+ K2O wt %

16 14

Ridge Suite Paringa Suite Newton River Granodiorite Big Pluton west Tine Peak Tonalite

12

alkaline rocks

10 8 6

G

4

Gabbro

2

Granite

Diorite

sub-alkaline/(tholeitic/calc-alkaline)

0 30

40

50

60

SiO2 wt %

70

80

Figure 26. TAS diagram (total alkalis versus silica, Le Maitre et al., 1989) showing SiO2 versus Na2O + K2O for classification of plutonic rock types from Cox et al. (1979) and dividing line separating tholeiitic/sub-alkaline from alkaline (Irene and Baragar 1971). Southwest Fiordland rocks are plotted with representative compositions from Paleozoic plutonic suites from New Zealand. Refer to appendices for included samples.

Figure 27. Na2O versus K2O for plutonic Paleozoic granitoids of southwest Fiordland and representatives of Paleozoic New Zealand plutonic suites. I-S type granite dividing line taken from Chappell & White (2001).

80

Figure 28. Sr versus Rb plot showing Carboniferous units from southwest Fiordland plotted with representative compositions of Paleozoic Suites as suggested by Tulloch (2003).

Figure 29. Discrimination diagram for A-, I- and S-type plutonic rocks after Whalen et al. (1987).

In addition, a younger (320-300 Ma) suite of plutons was recognised, which Tulloch et al. (2003) names the Foulwind Suite. Although the Foulwind Suite is referred to as having AType affinities (Tulloch et al., 2003), most plutons lack the distinguishing geochemical characteristics noted by Eby (1992) and Whalen et al. (1987). This is apparent when the compositions are plotted in the discrimination diagram of Whalen et al. (1987) for I-, S-, and A-type granites (fig.29). Foulwind Suite samples do not exhibit the high Ga/Al ratios typical of A-type granites and largely overlap the compositions of the other NZ Paleozoic plutonic suites. The later timing of the Foulwind Suite relative to earlier I- and S-type suites is a

81 common occurrence for A-type granites in other plutonic provinces where they occur at the end of conventional arc magmatism (Creaser et al., 1991). Limited previous dating of granites in southwest Fiordland by zircon U-Pb methods by Muir et al. (1998) and Tulloch et al. (2003) suggest that plutons coeval with the Ridge and Tobin Suite may be well represented.

5.3.2 Geochemistry of Paleozoic granitoids in southwest Fiordland 5.3.2.1 Major elements Relative to other Paleozoic granitoids in New Zealand, Tpt is characterised by its extremely low total alkalis (fig.26), which is reflected in its calcic classification on a modified alkali lime index (MALI) diagram of Frost et al. (2001)(fig.30). Furthermore, depletion in K2O relative to Na2O gives correspondingly high Na2O/K2O ratios (1.5 - 3.7)(fig.27). The samples are weakly to moderately peraluminous with ASI between 1.05 and 1.2 (where ASI was defined by Zen (1986) as molar Al2O3/(CaO+K2O+Na2O)), which is similar to other Paleozoic granitoids (fig.25), although they are distinctive in having high molar Al/(Na + K). I have used the modified ASI of Frost et al. (2001) which takes into account the presence of apatite. However it is noted that the formula reported by these authors is incorrect and should be ASI = Al/[2(Ca – 1.67P) +Na+ K] rather than Al/(Ca – 1.67P + Na +K). Chappell and White (1992) originally defined S-type granites from the Lachlan Fold Belt (LFB) as having aluminium saturation indices (ASI) > 1.1, and I-types as having ASI < 1.1. On a normative An-Ab-Or granite classification diagram (Barker, 1979) three samples of Tpt plot as granodiorite, while four plot in the tonalite field (fig.32). In addition, Tpt is further characterised by high Fe2O3, CaO, TiO2, (fig.31). Harker variation trends are generally smooth but with distinctive enrichment in CaO, Fe2O3 (total), TiO2, and P2O5 and depletion in K2O and to a lesser extent Na2O (fig.27). The most mafic Tpt sample is enriched in Al2O3, CaO, although not Na2O (fig.31), which may indicate accumulation of plagioclase. Table 5. Summary of defining geochemical characteristics of Tine Peak Tonalite, Big Pluton east, and Big Pluton west.

Unit

ASI

Peacock (MALI)

Zr

Sr

Y

Sr/Y

(avg)

(avg)

(avg)

(avg)

Na2O/K2O

Tine Peak Tonalite

1.03 - 1.17

calcic

1.4 – 3.7

475

400

30

20

Big Pluton East

1.08 - 1.15

calc-alkali

1.1 – 1.7

120

640

15

47

Big Pluton West

1.07 – 1.21

calc-alkali – calcic

0.5 – 1.3

170

230

29

8

82 Analysed samples of Big Pluton can be divided into two distinct geochemical groups referred to as Big Pluton west (BpW) and Big Pluton east (BpE). Big Pluton west includes samples that can be geochemically correlated with KAK1; the sample for which Muir et al. (1998) reported an age of c.340 Ma and referred to as Kakapo Granite (refer to discussion in section 2.4.3). Samples here correlated with KAK1 include P70580, P70589, P70613, P70778, P70785, and P70798 (Qmap Fiordland samples, reported in Petlab). Big Pluton east outcrops in a north-south linear belt to the east of Solitary Peak and includes samples: P70776, P70777, P70783, P70791, and P70799 (Qmap Fiordland samples, reported in Petlab). Sample P70791 was dated in this study and yielded a U-Pb zircon age of 368 ± 2 Ma (see section 4.1.2). The major element distinction between BpE and BpW is most apparent in Al2O3 and Na2O, which are significantly more enriched in BpE (fig.31). Accordingly BpE has higher Na2O/K2O ratios (table.5). On a MALI (Frost et al., 2001), BpE is classified as calc-alkali, while BpW is calcalkali to calcic (fig.30). On a normative An-Ab-Or granite classification diagram BpE is limited in composition being confined to the granodiorite field, while BpW ranges from granodiorite to monzogranite and granite in composition (fig.32). Major element compositions of BpE are noted for being very similar to those of Newton River Granodiorite and Indian Island Granite. 15.00

alkalic

10.00

Na2O+K2O-CaO

Karamea Suite Ridge Suite Paringa Suite Tobin Suite Foulwind Suite Newton River Granite Evans Pluton Big Pluton west Tine Peak Tonalite Indian Island Granite Big Pluton east

5.00

alkalicalcic

0.00

calcic

-5.00

calc-alkali

-10.00

-15.00 45

50

55

60

65

70

75

80

SiO2 Figure 30. Paleozoic southwest Fiordland rocks and representative samples from New Zealand Paleozoic plutonic suites plotted on a modified alkali lime index (MALI) diagram (after Frost et al. 2001). Figure 31. (next page) Major element Paleozoic plutonic Harker diagrams with data plotted for southwest Fiordland Paleozoic rocks. Representative samples from other New Zealand suites are plotted for comparison (See appendices for list of data and sources).

83 6 20

5.5 5

18

4.5

14

12

Na2O

Al2O3

4 16 Foulwind Suite

Tine Peak Tonalite

Karamea Suite

Newton River Granodiorite

Ridge Suite

Indian Island Granite

Paringa Suite

Big Pluton east

Tobin Suite

Big Pluton west

3.5 3 2.5 2 1.5

Evans Pluton

1

10 45

50

55

60

65

70

75

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80

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10 9

6

8 5

7

K2O

CaO

6 5

4 3

4 2

3 2

1

1 0

0 45

55

65

75

45

85

50

7

12

6

10

5

MgO

Fe2O3

8

6

4

4 3 2

2

1 0

0 45

50

55

60

65

70

75

45

80

50

55

60

2.0

0.6

0.5 1.5

TiO2

P2O5

0.4

0.3

1.0

0.2 0.5

0.1

0.0

0.0

45

50

55

60

SiO2

65

70

75

80

45

50

55

60

SiO2

84

An Tpt Nrg Big Pluton west Ep Wgg Indian Island Granite Big Pluton east Tpt this study

tonalite

granodiorite monzogranite

trondhjemite granite

Ab

Or

Figure 32. Paleozoic southwest Fiordland rocks plotted on a normative Ab-Or-An diagram. Granitoid discrimination fields are after Barker (1979).

Tine Peak Tonalite Indian Island Granite Big Pluton west Big Pluton east Evans Pluton Widgeon Granitoid Gneiss Newton River Granodiorite

Figure 33. Diagram showing H2O-saturated cotectics and minimums in the Qtz-Ab-Or-H2O system at 0.1-1000 MPa. Normative compositions for Paleozoic granitoids of this study corrected for normative An according to Blundy and Cashman (2001). Compositions not conforming to the requirement of having less than 20 wt % normative anorthite, and 2 wt % normative corundum are highlighted with a red dot.

85 Normative compositions corrected according to Blundy and Cashman (2001) are plotted in the H2O-saturated Ab-Or-Qtz (haplogranite) phase diagram in figure 33. The data for BpE appears to fall along a trend of plagioclase fractionation which intercepts the quartz + alkali feldspar cotectic at a pressure of c. 500 MPa. In contrast, BpW data trend along the 200 MPa quartz + alkali feldspar cotectic. This could indicate that BpW magmas equilibrated with quartz + alkali feldspar residues at a lower pressure than BpE. In both cases final emplacement pressures may have been lower. 5.3.2.2 Trace elements Samples of Tpt are distinguished from other NZ Paleozoic granitoids by enrichment in Zr and Sr (figs. 34). Samples plotted on Harker diagrams show no enrichment or depletion in either Rb or Ba with increasing SiO2, while Sr shows a general decreasing trend (fig.33). A scattered decreasing trend is also shown in Zr, although three samples exhibit extreme enrichment in Zr relative to the other samples which are themselves high (fig.34). Large ion lithophile (LIL) elements Pb and Rb are significantly depleted in Tpt relative to average I- and S-type compositions and representative Karamea Suite samples (figs.34 and 35). Two distinct ranges of Sr are apparent in Tpt, one with Sr = 200-300 and the other with Sr = 500-600. However, in other respects there seems to be no systematic differences in these samples. Complete rare earth element (REE) data were measured for the five Tpt samples collected in this study. The data presented on a chondrite normalised REE diagram in figure 36. The most mafic Tpt sample (OU 75189, 58 wt % SiO2) has a positive Eu anomaly [(Eu/Eu*)N = 1.25 where (Eu/Eu*)N = EuN/(SmN.GdN)½ and N subscript indicates chondrite normalised values) fig.36] which as previously noted is consistent with plagioclase accumulation.. The four remaining samples all have SiO2 > 73 wt %. Samples OU 75181, OU 75182, and OU 75184 have slight to moderate negative Eu anomalies [(Eu/Eu*)N = 0.86 - 0.66]. Sample OU 75193 has a deep Eu anomaly [(Eu/Eu*)N = 0.21] and is highly enriched in the trivalent REE as well as Ga and Y relative to other Tpt samples (figs. 34 and 36). These characteristics are similar to those of typical A-type granites and this sample plots well into the A-type field on a 10000.Ga/Al versus Zr diagram of Whalen et al. (1987) (fig.29). As with other Tpt samples, OU 75193 differs from typical A-type chemistry (Eby, 1992, Creaser et al., 1991, Whalen et al., 1987) in having lower Rb, Nb, K2O, and higher CaO, and Al2O3. In contrast to other Tpt samples, OU 75193 is also depleted in total Fe2O3 and concentrations of high field strength elements (HFSE) Ti, Nb, and Ta when compared to the other samples of Tpt.

86

3500

Foulwind Suite

Tine Peak Tonalite

Karamea Suite

Newton River Granodiorite

Ridge Suite

Indian Island Granite

Paringa Suite

Big Pluton east

Tobin Suite

Big Pluton west

3000 2500 2000

Ba

Evans Pluton

1500 1000 500 0

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45

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Sr

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SiO2

65

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75

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45

50

55

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SiO2

65

70

75

80

Figure 34. Trace element Harker diagrams with data plotted from southwest Fiordland Paleozoic granitoids and representative samples from other New Zealand Paleozoic plutonic suites.

Trace element compositions of analysed samples of Big Pluton appear to mirror the major elements by clustering into two geochemically distinct groups. This is most apparent for Sr and Y where BpE consistently has lower Y and higher Sr than BpW. Accordingly, this is reflected in the significantly higher Sr/Y ratio (23-62) for BpE compared to BpW (Sr/Y = 4-

87 9). Samples of BpE are also distinct in having lower Zr for equivalent SiO2 compared to BpW (fig. 34).

Figure 35. Primitive mantle normalised Spider diagram with Tine Peak Tonalite (grey shaded field) and Big Pluton plotted with representative Paleozoic rocks from the S-Type Karamea Suite and I-type Tobin Suite. Also plotted are the average compositions of S- and I-type granites from the Lachlan Fold Belt (Chappell and White, 1992), and average A-types from around the world (Whalen et al., 1987, REE from Collins et al., 1982). Normalising values from Sun and McDonough

Figure 36. Chondrite normalised REE plot for Tine Peak Tonalite and Big Pluton west (Sample KAK 1, Muir et al., 1998) and representative samples from the I-type Tobin Suite (Lake Roxburgh Tonalite NF 11, Muir et al., 1998) and two S-type

88 granites from the Karamea Suite representing two extremes of crustal contamination (Whale Ck Granite RNZ 169 and Dunphry Granite RNZ 194, Muir et al., 1996b). Chondrite values from Sun and McDonough (1989).

Figure 37. Zircon saturation temperature (TZr)(Miller et al., 2003) versus Rb/Sr ratio of southwest Fiordland Paleozoic rocks and representative samples from New Zealand Paleozoic suites.

5.3.2.3 Zircon saturation temperatures (TZr) Zircon saturation temperatures (TZr) of Miller et al. (2003) define possible differentiation trends for Tpt and BpW when plotted against Rb/Sr ratio as a fractionation index (fig.37). It must be remembered that TZr cannot necessarily be interpreted as the crystallisation temperature (see discussion 4.5). Samples with the lowest Rb/Sr ratios within Tpt and BpW have the highest TZr, as would be expected from a cooling and fractionating magma. Again, the distinction between BpE and BpW is apparent with BpW having higher Rb/Sr ratios for the samples with highest TZr. If BpW was a more evolved phase of BpE, as suggested by its higher Rb/Sr, it would be expected to have lower TZr as well. The cause of the higher Rb/Sr ratio in BpW is due solely to lower Sr relative to BpE, as the Rb concentrations in both are relatively similar (fig.28). This argues against a relationship through simple fractionation and suggests that differences between BpW and BpE are likely to be due to distinct magma source regions or conditions of melting. 5.3.2.4 Zircon REE chemistry Partitioning of trace elements between zircon and melt is principally controlled by similarity of size and charge of the ions substituting for Zr4+ in the zircon crystal lattice. As a

89 consequence, Hf, U, Th, the trivalent HREE and Ce (IV) are strongly enriched in zircon relative to the LREE and Eu (II). In this study, positive Ce anomalies in zircon are quantified using an approach modified from that described by Ballard et al. (2002). A subset of the REE (La, Ce, Sm, Eu, Dy, Lu) were analysed simultaneously during LA-ICP-MS analysis for U-Pb dating. The zircon REE data were normalised to values for the bulk continental crust from Taylor and McLennan (1985). Crust normalised Ce/Ce’ and Eu/Eu’ values were calculated by dividing measured Ce and Eu concentrations by interpolated (Eu’) and extrapolated (Ce’) values determined by linear regression of the Sm, Dy and Lu versus an lattice-strain (ionic radius) parameter according to Ballard et al. (2002). This method minimises the effects of contamination by LREE-rich inclusions during analysis. 10000

10000

Tpt OU 75184

1000

zircon Ce/Ce'

zircon Ce/Ce'

1000 100 10 1 0.1 0.00

100 10 1

0.20

0.40

0.60

0.80

1.00

zircon Eu/Eu' 10000

zircon Ce/Ce'

Tpt OU 75193

0.1 0.00

0.20

0.40

0.60

zircon Eu/Eu'

0.80

1.00

Bp P70791

1000

100

10

1 0.00

0.20

0.40

0.60

0.80

1.00

zircon Eu/Eu' Figure 38. Crust-normalised Eu/Eu’ versus Ce/Ce’ anomalies from individual zircons used in the dating of Tine Peak Tonalite (OU 75184, OU 75193), and Big Pluton (P 70791). See text for definition and method of calculation.

Zircons in Tpt samples (OU 75184 and OU 75193) show anomalously low Eu/Eu’ and Ce/Ce’ (Ballard et al., 2002)(Fig.38). More typical values for zircons from high SiO2 rocks are

90 Eu/Eu’~ 0.4 and Ce/Ce’~ 100, as seen in zircons from Bp (P70791) and the PIC (fig.51). Low Ce/Ce’ anomalies in zircon have been suggested by Ballard et al. (2002) to indicate a reduced magma, while low Eu/Eu’ may also reflect a reduced magma or residual plagioclase in the source, extensive fractionation of plagioclase, or a source with an inherited negative Eu anomaly. 5.3.2.5 Oxygen isotopes Oxygen isotopes were determined for quartz and feldspar mineral separates and a whole rock powder of sample Tpt (OU 75184). The results are listed in table 6. A review of whole rock values obtained from many large sized plutons shows that δ18O varies from -2 to +16, although most are restricted to values of +5.5 to +11 (Taylor, 1988). Realistic δ18O values for magmatic quartz are no lower than +6, and for feldspar +5. The difference in δ18O values between quartz and feldspar can be used in conjunction with the equilibrium oxygen isotope fractionation factors of Chacko et al. (2001) to calculate a temperature. The resulting temperature of c.300oC is less than the estimated solidus temperature of ≥ 650oC for a granitic magma. This indicates that some sub-solidus exchange of oxygen isotopes occurred during closed-system cooling (Gilletti, 1986) or hydrothermal alteration (Taylor and Sheppard, 1986). Because quartz is most resistant to such effects, a magma δ18O value can be estimated from the δ18O quartz value by subtracting an empirical “correction factor” of 1‰ (Faure and Brathwaite, 2006). The δ18O obtained in this way is 12.9 and agrees with the whole rock δ18O, indicating a closed-system sub-solidus history for oxygen isotopes. Such a high δ18O value requires involvement of sedimentary materials in the source of Tpt. Table 6. Oxygen isotopic compositions of quartz, feldspar, and whole rock for Tine Peak Tonalite.

Sample Tine Peak Tonalite (OU 75184)

δ18O qtza

δ18O fspa

δ18O wrb

(‰, VSMOW)

(‰, VSMOW)

(‰, VSMOW)

13.9

11.4

12.3

δ18Omagmac

12.9

a Analyses are on quartz and feldspar, which were hand picked grains after crushing the rock sample to 250 μm. Isotopic analysis was done by Dr. K. Faure (Stable Isotope Facility, IGNS, New Zealand). b

Samples processed from whole rock powders by Prof C. Harris (Geological Sciences, University of Cape Town).

c

δ18Omagma estimated by subtracting 1.0 from δ18OQtz (Faure and Brathwaite, 2006) for Δ18Oqtz-melt closed system cooling effects.

5.3.2.6 Neodymium and strontium isotopes Neodymium isotopes and Sr isotopes were measured by TIMS on whole rock powders (crushed at Department of Geology, University of Otago) from the same two Tpt samples that were dated (OU 75184 and OU75193) by Dr M. Norman of PRISE at the ANU. Initial isotopic compositions were calculated at the U-Pb zircon age using Sm/Nd measured by LA-

91 ICP-MS and listed in table.7. Recalculated values give initial εNd values of -1.9 and -4.6 and 87

Sr/86Sr values of 0.707537 and 0.707136.

Table 7. Presented Sr- and Nd- isotopic data for Tine Peak Tonalite showing calculated compositions. 87Rb/86Sr and 147 Sm/144Nd ratios calculated from trace element results reported in section 10.3.7. Ages from table.2.

Sample Tine Peak Tonalite (OU75184) Tine Peak Tonalite (OU75193)

Age (Ma)

87

147

87

350

0.748

0.711242 0.70754

0.1007

0.512321

-1.9

350

0.239

0.708333 0.70714

0.1025

0.512188

-4.57

86

Sr/ Srm

87

86

Sr/ Sri

Sm/144Nd

εNd

Rb/86Sr

143

Nd/

144

Ndm

(initial)

5.3.3 Discussion and Petrogenesis Compositional variation within Tpt appears haphazard in some respects, particularly with regard to Zr (fig.34), although smooth trends in the major elements CaO and MgO (fig. 31) suggest the variation is related to crystal-liquid fractionation and not the result of unrelated batches of magma. It is proposed that the best explanation for the internal chemical variation within Tpt is fractionation of cumulate minerals, which trapped varying amounts of interstitial melt. The positive Eu anomaly in the lowest SiO2 sample confirms that some samples have a cumulate plagioclase component, while varying negative Eu anomalies in the higher SiO2 samples suggest some degree of feldspar fractionation. Although complete REE data for the intermediate SiO2 samples was not available, it is suspected that they would have Eu anomalies intermediate between those of the high and low SiO2 samples. If crystal fractionation is invoked as the cause of Tpt variation, the parental magma composition should be of only slightly lower SiO2 than the majority of the samples probably in the range of 67-72 wt % SiO2. The A-type granite characteristics of Tpt (fig. 29, particularly OU 75193) are here ascribed to high temperature, low fH2O conditions of melting of a source of intermediate composition. The depletion of K2O, Rb and Pb in Tpt argue for a source lacking in either orthoclase or biotite, as these are the most common mineral reservoirs for these elements. Such a source could be broadly tonalitic in composition. Rutter and Wyllie (1988) found orthoclase and biotite to be among the early phases to break down during dehydration melting of a garnet-bearing tonalite at 10 kbar. Above the orthoclase-out and biotite-out reactions, the residual assemblage was composed of plagioclase + quartz + hornblende + orthopyroxene + garnet + magnetite + titanite. High concentration of CaO, Fe2O3, and TiO2 in Tpt with no concomitant enrichment in Al2O3 could result from the breakdown of a hornblende. This seems likely given that hornblende would be the most likely H2O-bearing residual mineral capable of fluxing

92 significant volumes of melt, as seems required given the size of Tpt. Dehydration melting experiments have established that breakdown of hornblende occurs between 900 and 1000oC (Beard and Lofgren, 1991, Rutter and Wylie, 1988), which agrees with the high TZr of many of the Tpt samples. A potential residual source lacking orthoclase and biotite was considered and rejected by Collins et al. (1982) for A-type granites, who concluded that ‘typical’ A-type granites required a more fertile source given their enrichment in alkalis and Rb. However, such a residual source may be consistent with the Tpt compositions, which show severe depletion in alkalis. Collins et al. (1982) suggested that such a residual source could retain accessory phases such as apatite, titanite, and zircon, after yielding a near minimum temperature I-type melt, although the same could be true of a metasedimentary source yielding a S-type melt. The breakdown of these accessory phases would be expected at the higher melting temperatures required to melt a source deficient in biotite and orthoclase. Thus, high P2O5, Zr, Y, LREE in Tpt argues that high temperature melting resulted in the dissolution of apatite, zircon, and monazite. Fluorine released from amphibole may have aided HFSE mobility in the magma due to its potential complexing capabilities (Collins et al., 1982). In addition, high F may have aided in-situ fractionation by further reducing viscosity in an already high temperature magma, as proposed by Clemens et al. (1987) for A-type magmas. As saturation in Zr appears to have occurred only at high SiO2 contents, and as zircon was likely lost to the more mafic compositions through crystal fractionation, it is possible that the Zr concentrations in the parental melt were initially higher and that the magmatic temperatures were considerable higher than that suggested by TZr (840-940oC). As genesis of Tpt most likely involved the breakdown of hornblende, the true magmatic temperatures may have been closer to 950oC. Dissolution modelling of zircon by Watson (1996) indicates that survival of zircon above 850oC would be unlikely, which is supported by an absence of inherited zircon in high temperature I-type (Chappell et al., 2004) and A-types granites (King et al., 2001). This suggests that inheritance identified in OU 75184 resulted from AFC processes at a relatively high crustal level and occurred after the magma temperature had reduced. However, zircon survival in high temperature magmas has been documented elsewhere (Hoskin et al., 2000) so, although its survival is enigmatic, incorporation of zircon from the source cannot be eliminated as a possibility. The overall tendency of incompatible elements such as K2O and Rb not to increase with fractionation in Tpt suggests that some melt was incorporated amongst cumulate crystals and since these elements were in low concentrations in the parent melt any increase with

93 fractionation would be more subtle. K2O, Rb, and Ba all decrease sharply at higher SiO2 values. This is attributed to late saturation in H2O, and hence late crystallisation of biotite. This supports an initially low but not negligible H2O content in the primary Tpt magma. Samples plotted within a haplogranite phase diagram using the method of Blundy and Cashman (2001), define a trend of decreasing pressure up to the 100 MPa cotectic and displaced towards Ab-Qtz, which may reflect equilibration at 100 MPa and emplacement at 4 km. However, it is cautioned that many Tpt samples have normative corundum between 2 – 3 wt %, which is higher than the 2 wt % limit for use of this approach. Nonetheless, a lack of deformation fabrics in Tpt is broadly consistent with such a low emplacement pressure, which may be further evidence of initially low magma H2O contents permitting ascent higher in the crust before complete crystallisation. Calculated magma δ18O value of 12.9 and initial εNd of -4.6 to -1.9 and initial

87

Sr/86Sr =

0.7073 ± 0.0002 for Tpt argue for a source dominated by metasediments, most likely Paleozoic Greenland Group greywacke, rather than amphibolite or tonalite. It must be pointed out that the Western Province is dominated by Greenland Group greywacke and that it has been considered the sole sedimentary source component in petrogenetic studies of Western Province granitoids dominated by a sedimentary source (Muir et al., 1996b, Muir et al., 1995, Waight et al., 1998b). This seems justified as granitoids considered to have a sedimentary component show strong enrichment in Rb, an element highly enriched in Greenland Group greywacke. It is also supported by isotopic mixing models presented by Muir et al. (1996b, 1995) for Karamea Suite, and Waight et al. (1998b) for Rahu Suite that are consistent with involvement of a Greenland Group component. Greenland Group greywackes have εNd ranging from -4 to -10 (age corrected to 350 Ma, data from Waight et al., 1998b), which are slightly more evolved on average than Tpt. However, it is possible that A-type magmatism was itself driven by basaltic magmatism resulting from decompression melting of the asthenosphere which accompanied crustal thinning. Such mafic magmatism is the most likely method of heat transfer to supracrustal material needed for A-type melt generation. Some degree of mixing between these mafic melts and the supracrustal melts they help generate would be expected to occur. The calculated magma δ18O of Tpt are consistent with a less weathered Greenland Group greywacke source, although Paleozoic metasediments in southwest Fiordland cover a large δ18O range from 10 to 19 (Simpson, in prep), which makes proportions difficult to constrain and allows for mixing with a mafic magma component. Conversely, if a model involving mixing of a more evolved Greenland Group greywacke and basement igneous source was invoked, inheritance in Tpt (OU 75184) coeval with the I-type

94 BpE could provide evidence for its assimilation into Tpt magmas, although differences in zircon trace element systematics appears to argue against such a proposal. Thus, the most viable source of Tpt from the evidence discussed is a less fertile Greenland Group greywacke that had already yielded a melt, here proposed to be the older transitional I-/S- type granitoids of the Karamea Suite. The Black Giants Anorthosite (349 ± 5 Ma, Gibson and Ireland, 1999) in central Fiordland is coeval with the Tpt. Gibson and Ireland (1999) state that the tectonic environment in which the Blacks Giant Anorthosite was intruded was likely to be continental crust undergoing extension. They hypothesised thermal weakening followed the previous magmatism and lowpressure, high-temperature sillimanite grade metamorphism dated at 360 Ma. Anorthosite magmas are thought to be extremely high temperature (Winter, 2001), so the coeval Black Giants Anorthosite relationship to the high temperatures recorded by the Tpt is unlikely to be coincidental. It is most probable that the Tpt and Black Giants Anorthosite are both related to an episode of crustal thinning accompanied by elevated geotherms following a period of convergent calc-alkaline arc magmatism. The base of the crust in such an arc may have been less fertile due to previous melt generation and required high temperatures to melt. The Blacks Giant Anorthosite requires that anomalously high mantle heat flow was operative at this time. The crustal residue after generation of the Tpt of would have been even more refractory in nature. An interesting question to consider is the nature of the residue left in the lower crust after later stage high-temperature melting of a source that had already yielded earlier S-type melts. The enrichment of CaO and lack of a corresponding increase in Al2O3 in Tpt would leave an already residual source depleted in CaO, which would make the stabilization of plagioclase difficult, thus making garnet the most likely restitic phase to accommodate the Al2O3 . It is likely then, that the Western Province metasediments, at least during the Paleozoic, and possibly to this day are underlain by a highly refractory granulite lower crust. If this lower crust had become garnet-granulite and suffered delamination after its formation, then it would have presumably resulted in whole scale melting of the overlying crust as hot mantle come into contact with more fertile crust; something not seen in the plutonic record of the Western Province. The BpE, Nrg, and Iig have similar chemistry and a narrow compositional range. They all show characteristics typical of I-type granites, including Na2O/K2O > 1, high Sr, and low P2O5 in contrast to BpW and Ep. A lack of Gondwanan zircon inheritance in BpE also

95 suggests that Paleozoic metasediments were not a dominant source component. Depletion in Na2O and Sr in BpW and enrichment of Al2O3, Na2O, and Sr in BpE, are consistent with the effects of varying fH2O during melting; attributable to the effect of H2O on the stability of plagioclase and amphibole in the source. Higher fH2O during melting in BpE would have stabilised amphibole +/- garnet in the source, while the albitic component of plagioclase was partitioned into the melt, which is consistent with the low Y and high Al2O3, Na2O, and Sr in BpE. Lower calculated TZr in BpE may reflect lower magmatic temperatures, which further support higher fH2O during melting. The contrasting chemistry of BpE and BpW appear to suggest that they were derived from similar sources, but involved melting at different fH2O. Geochronology is inconclusive with respect resolving an age difference between BpW and BpE due to a lack of resolution on dating of BpW (Ewing, 2003, Muir et al., 1998), although the data permit BpW to be slightly younger.

5.4 Mesozoic Geochemistry Mesozoic chemistry includes an analysis of PIC data (as defined in section 3.2) from rocks from my own fieldwork, and also includes analyses of PIC rocks and Lake Mike Granite analyses made available by Qmap Fiordland. Refer to appendices for a complete listing of data sources. Sample locations are plotted on the map in the back cover.

5.4.1 Previous work on NZ Mesozoic suites Mesozoic plutons in New Zealand are presently correlated with the following three main suites: Separation Point Suite (SPS), Darran Suite, and Rahu Suite. The SPS is geochemically similar to Na-rich Phanerozoic granitoids, Archean trondhjemites [part of the TonaliteTrondhjemite-Granite (TTG) association], and modern adakites (Drummond and Defant, 1990). According to Smithies (2000) the defining criteria for ‘TTG like chemistry’ is Al2O3 > 15 wt %, Na2O/K2O > 1 and Sr and Ba > 400 at c.70 wt% SiO2. By these criteria SPS tend to be strongly ‘TTG like’ with Al2O3 > 16 wt %, Na2O/K2O > 1.5 and Sr and Ba > 700 (figs. 39, 43 and 45). SPS have extreme Sr/Y ratios (>100), which are considerably greater than minimum values of 40 normally used to distinguish ‘TTG like’ or ‘adakitic’ chemistry (eg. Atherton and Petford, 1993). Tulloch and Kimbrough (2003) have suggested the term HiSY (high Sr/Y) for granitoids with these characteristics based on their probable genesis from melting of an amphibolite underplate in continental arc environments (Atherton and Petford, 1993) in order to differentiate them from true adakites, which are thought to be melts from subducted oceanic crust (Defant and Drummond, 1990).

96 8

Darran granitoid Darran mafic

Na-Series

7

Transitional Darran Rahu

6

K-Series

Separation Point granitoid Separation Point mafic

Na2O

5

WFO Treble Mtn Granite

4

Revolver Pluton Long Scarp Granodiorite

I-type

3

Upper Blacklock Granite North Port Granite

S-type

2

Trevaccoon Diorite Cuttle Cove Gabbro

High-K-Series

1

Lake Mike Granite Brothers Pluton

0 0

2

4 K2O

6

8

Monk Granite Only Island Gabbro

Figure 39. Na2O versus K2O for Mesozoic plutonic rocks of southwest Fiordland plotted with representative analyses from New Zealand Mesozoic suites.

20 18

Na2O + K2O wt %

16 14 12

Darran Suite

Transitional Darran Suite

Rahu Suite

Separation Point Suite

WFO

Treble Mountain Granite

Revolver Pluton

Long Scarp Granodiorite

Upper Blacklock Granite

North Port Granite

Trevaccoon Diorite

Cuttle Cove Gabbro

Lake Mike Granite

Brothers Pluton

Monk Granite

Only Island Gabbro

10

alkaline

8

G

6

Q

granite

4 sub-alkaline (tholeitic/calc-alkaline)

2 0 35

45

55

65

SiO2 wt %

75

Figure 40. TAS diagram (total alkalis versus silica) of Cox et al. (1979) showing SiO2 versus Na2O+K2O. Fields for the classification of plutonic rock are adapted from Wilson (1989) The curved dividing line subdivides tholeiitic/sub-alkaline from alkaline compositions and is after Irvine and Baragar (1971). Southwest Fiordland rocks are plotted with representative compositions from the Rahu, Separation Point, Karamea, and Darran Suites (Refer to appendices for included samples).

The Rahu Suite is coeval with SPS and is thought to represent SPS magma contaminated by Greenland Group greywacke (Waight et al., 1998b). Lower Na2O, higher Fe2O3, and highly enriched Rb in Rahu Suite rocks are consistent with this hypothesis. Darran Suite rocks are more typical of I-type continental arc compositions with Al2O3 between 14 and 15 wt %,

97 Na2O of c.4 wt %, and Sr mostly < 400 ppm (at c.70 wt% SiO2). K2O, Ba and Rb are more variable than in either the SPS or Rahu suites. Mesozoic analyses from southwest Fiordland, and representative analyses from New Zealand suites are plotted on a TAS diagram in figure 40. It is apparent that the analysed samples are dominated by compositions with SiO2 between 65 and 75 wt %. Although mafic units are numerous in southwest Fiordland, their relative volumes are small compared to the felsic units, which has resulted in fewer analyses, and even less emphasis on determining their petrogenesis. Some of these mafic units were deliberately targeted in this study to test for genetic links with the more felsic units.

5.4.2 PIC geochemistry 5.4.2.1 Major elements The PIC granites span a SiO2 range between 69 and 75 wt % (fig.40). Harker diagrams for major elements show negative trends with SiO2 for CaO, Al2O3, Fe2O3 (total), and TiO2, (fig.42). Na2O shows smooth trends for most PIC granites, although only slight depletion or enrichment is evident with increasing SiO2. No such trends are seen in K2O, which shows considerable scatter, especially in Rp and Tmg. Granite samples are mostly weakly peraluminous with aluminium saturation indices (ASI, see Paleozoic section for definition) between 1.0 and 1.2 (fig.41). All PIC granites are K-series rocks (fig.39), and in this respect are similar to granitoids from the Rahu and Darran suite. In contrast, SPS granitoids, WFO, and mafic Darran Suite samples fall mostly within the Na-series, as does the Tdi, Oid, Npg and Lsg. The Lsg is also similar to SPS in other aspects of major element chemistry including its granodioritic to trondhjemitic classification on a normalised Ab-Or-An granitoid classification diagram of Barker (1979)(fig.42). In addition the Lsg is also high in Na2O (c.6 wt %), Al2O3 (c.19 wt %), and low in CaO, MgO, Fe2O3t, and TiO2 (fig.43). The Rp, Ubg, Tmg, and Monk Granite plot in the field of granite sensu-stricto, and have trends towards increasing An content (fig.42). The Lmg plots in the granite field also, but samples contrast in having a trend towards increasing Ab content. The Tmg contrasts to Rp and Ubg in having lower SiO2 (Avg. Tmg, 70 wt %; Rp, 72 wt %; Ubg, 73 wt %) and Na2O (Avg. Tmg, 3.6 wt %; Rp, 4.2 wt %; Ubg, 4.1 wt %), which is reflected in its slighter lower average Na2O/K2O of 0.89; compared with 1.05 for Rp and 1.03 for Ubg. This is most likely magmatic as samples that are no more altered than either of Rp and Ubg still show depletion in Na2O.

98 5

mol Al/(Na+K)

4

3

Darran Mafic

Darran granitoid

Transitional Darran

Rahu

Separation Point granitoid

Separation Pt mafic

WFO

Lake Mike Granite

Treble Mtn Granite

Revolver Pluton

Upper Blacklock Granite

North Port Granite

Long Scarp Granodiorite

Trevaccoon Diorite

Cuttle Cove Diorite

Monk Granite

Brothers Pluton

Only Island Diorite

Peraluminous

2 Metaluminous 1

Peralkaline

0 0

0.2

0.4

0.6

0.8 ASI

1

1.2

1.4

1.6

Figure 41. Aluminium Saturation Index of Mesozoic southwest Fiordland plutonic rocks. Representative compositions of Rahu, Separation Point, and Darran Suite rocks are plotted for comparison. ASI is calculated by the following formula: Al/(2(Ca-1.67-P)+Na+K). Note formula contains a correction to the formula in Frost (2001). Correction made by Palin, J.M (2004).

An Treble Mtn Granite Revolver Pluton Upper Blacklock Granite Long Scarp Granodiorite Brother Pluton Lake Mike Granite Monk Granite

tonalite

granodiorite monzogranite

trondhjemite

Ab

granite

Or

Figure 42. Normative feldspar proportions for PIC granitoids with SiO2 > 63 wt % plotted on a Ab-Or-An diagram. Granitoid discrimination fields are after Barker (1979).

99 24 10

22 8

20

Al2O3

CaO

6

18

4

16

2

14

0

12 45

50

55

60

65

70

75

80

45

7

9

6

8

50

55

60

50

55

60

50

55

65

70

75

80

7

5

Na2O

K 2O

6

4 3

4

2

3

1

2 1

0 45

50

55

60

65

70

75

45

80

9

16

8

14

7

65

70

75

80

12

6

10

5

Fe2O3t

MgO

5

4

8 6

3 2

4

1

2

0

0

45

50

55

60

65

70

75

80

45

60

65

70

75

80

0.8 0.7

2.0 0.6

1.5

TiO2

P2O5

0.5 0.4

1.0 0.3 0.2

0.5

0.1

0.0

0.0 45

50

55

60

65

70

75

80

45

50

55

60

65

70

75

80

Figure 43. Major element Harker diagrams for Mesozoic southwest Fiordland plutonic rocks with representative analyses from the Rahu, Separation Point, and Darran suites. Refer to legend on ASI plot, previous page. Samples included are listed in the appendices.

100 Mafic rocks in the PIC are volumetrically less abundant and have more variable chemistry than felsic rocks. Tdi and Oid exhibit negative correlations with increasing SiO2 on Harker diagrams for Al2O3, CaO, Fe2O3 (total), MgO, TiO2, and a positive correlation for K2O (fig.43). The Tdi and Oid have very low SiO2 concentrations (< 50 wt %), although more evolved rocks, typically dykes occur in both units with SiO2 between 55 and 60 wt %. Samples of Tdi and Oid with the lowest SiO2 differ from MORB and island arc compositions in having higher Al2O3 and K2O, and lower MgO and CaO (Table 8). Low SiO2 Tdi samples also have high Na2O. A small gabbroic stock in Cuttle Cove, here named the Cuttle Cove Gabbro, is distinct from Tdi and Oid in having lower Al2O3, CaO, and Na2O. MALI classification 15.00 Darran Suite

alkalic

Rahu Suite

Na2O+K2O-CaO

10.00

Separation Point Suite Darran Transitional

alkalicalcic

5.00

WFO Treble Mtn Granite Revolver Pluton Upper Blacklock Pluton

0.00

Long Scarp Granodiorite

calcic

North Port Granite Trevaccoon Diorite

-5.00

Cuttle Cove Gabbro

calc-alkali

Lake Mike Granite

-10.00

Brothers Pluton Monk Granite

-15.00 45.00

50.00

55.00

60.00

SiO2

65.00

70.00

75.00

80.00

Figure 44. Modified alkali lime index (MALI) plot (after Frost et al. 2001).

5.4.2.2 Trace elements The PIC granites mostly display negative trends on Harker diagrams with increasing SiO2 for Sr, Ba and Zr (fig.45). A strong enrichment in Rb and Y is apparent in Tmg and Rp samples with increasing SiO2, which is not evident in Ubg. In contrast, Ba shows strong depletion with increasing SiO2 in Ubg, less so in Rp, and is scattered in Tmg. Rb like K2O shows greater scatter in Rp and Tmg relative to Ubg. The Ubg and Rp are similar in most respects except for Rb and Y, which are significantly lower in Ubg for equivalent SiO2. This suggests that they are not related through simple crystal-liquid fractionation, which is in agreement with the significantly different ages yielded for these two units. All PIC granitoids show arc signatures, including positive Pb anomalies relative to primitive mantle, and depletion in the HFSE ions Ti, Nb and Ta (fig.46), which in contrast to Pb is attributed to their immobility in the fluid rich mantle wedge environment (Kamber et al., 2002).

101 Table 8. Compositions of PIC gabbros and diorites compared with average MORB and typical island arc basaltic compositions. MORB data from Winter (2001, p248). Average island arc data Winter (2001, data from Ewart (1979, 1982)). MORB avg.

Typical Island Arc Low-K Medium-K High-K (tholeiitic)(mostly c-a) 50.70 50.10 49.80

Trevaccoon Diorite Cuttle Cove Gabbro Only Island Diorite OU 75111 OU 75154 OU 75163 OU 75140 OU 75169 OU75168 OU 75171

SiO2

50.50

58.12

48.68

49.08

48.36

55.67

46.45

TiO2

1.56

0.80

1.00

1.60

0.96

1.10

1.28

1.09

1.24

2.24

1.63

Al2O3

15.30

17.70

17.10

16.50

18.09

20.65

19.44

16.14

17.64

18.30

18.70

11.67 7.47 11.50 2.62

13.68 6.40 11.30 2.00

12.75 7.10 10.60 2.50

11.39 6.80 9.40 3.30

6.92 2.64 5.36 3.10

9.84 3.49 8.12 4.11

10.57 4.15 7.11 3.48

11.29 7.64 6.50 1.86

7.79 3.82 6.53 4.47

12.81 4.98 9.74 3.23

11.66 5.47 9.79 2.80

Fe2O3 MgO CaO Na2O

T

46.41

K2O

0.16

0.30

0.80

1.60

3.23

1.83

1.36

1.28

1.36

0.66

0.86

P2O5

0.13

0.10

0.20

0.50

0.30

0.68

0.72

0.22

0.39

0.33

0.19

Total

99.74

99.80

99.80

99.80

99.30

99.72

99.31

98.89

99.85

99.94

99.39

Norm q or ab an di hy ol mt il ap C

0.94 0.95 22.17 29.44 21.62 17.19 0 4.44 2.96 0.3

2.8 1.8 17 38.5 13.9 20.9 0 3.3 1.5 0.2

0 4.7 21.2 33.2 14.9 16.9 3.1 3.6 1.9 0.5

0 9.5 28 25.6 14.5 3 10.8 4.5 3.1 1.2

11.53 19.09 26.23 24.63

1.82 8.04 29.45 30.57

5.34 7.56 15.74 30.81

15.65

24.74

5.45 8.04 37.82 24.05 4.67 10.95

6.73 2.43

7.42 2.07

4.82 2.36 0.90

3.90 27.33 33.48 10.21 6.33 4.28 8.32 4.25 0.76

5.08 23.69 35.92 9.18 8.65 3.95 7.61 3.10 0.44

1.04

0.40

Mg #

68

55

69

60

56

60

10.81 32.24 32.49 2.73

10.12 4.22 1.82

9.23 6.13 2.09

0.47 53

56

55

54

53

Complete rare earth element (REE) data were measured for all PIC units; these data were normalised to chondrite values and are plotted in figure 48. The least fractionated PIC granites show slight negative Eu anomalies [(Eu/Eu*)N ~ 0.9]. The Rp shows moderate LREE enrichment with either a slight negative Eu anomaly (OU 75122), or a moderate negative anomaly (OU 75143) (fig.48). A slight positive Eu anomaly occurs in one Tmg sample [(Eu/Eu*)N = 1.21], which suggests accumulation of plagioclase. Spoon shaped patterns in MREE to HREE are prominent in the high SiO2 rocks (DyN < LuN); especially Rp and Tmg. The Rp, Tmg, and Ubg have high average LaN/YbN ~ 18 compared to average continental arc magmas with LaN/YbN ~ 9 (Smithies, 2000). These PIC granite LaN/YbN ratios are intermediate between SPS granites (avg. LaN/YbN ~ 28) and Darran Suite granites (LaN/YbN ~ 12) (fig.48, see appendices for included samples). A more transitional chemistry between SPS and Darran Suite could also be indicated by intermediate Sr/Y ratios between 40 and 50 in the least evolved PIC granite analyses (fig.47). Although these ratios would be considered ‘adakitic like’ when compared to that of modern adakites (Defant and Drummond, 1990), they are not considered HiSY in the New Zealand context. A granite dyke (OU 75151) cutting the Rp is similar in REE (fig.48) and major element composition (fig.43) to Tmg and it is interpreted as an offshoot from the Tmg. This crosscutting relationship is in agreement with the younger age obtained for Tmg. Extreme enrichment in REE in Lsg give correspondingly high LaN/YbN = 57. A moderate Eu anomaly

102 [(Eu/Eu*)N = 0.78 (OU 75175, fig.48)] is also distinct. In these respects, and also its lower SiO2 content, Lsg differs markedly from Rp, Tmg and Ubg. Darran Suite

Long Scarp Granodiorite

Separation Point Suite

Upper Blacklock Granite

Transitional Darran

North Port Granite

Rahu Suite

Trevaccoon Diorite

WFO

Cuttle Cove Gabbro

Treble Mtn Granite

Lake Mike Granite

Revolver Pluton

Brothers Pluton

2500

2000

Ba

1500

1000

500

Monk Granite

0 45

50

55

60

65

70

75

80

75

80

60

2000 1800

50

1600 1400

40

Y

Sr

1200 1000

30

800 20

600 400

10

200 0

0 45

50

55

60

65

70

75

45

80

50

55

60

65

70

50

55

60

65

70

250

450 400

200

350 300

150

Rb

Zr

250 200

100

150 100

50 50 0 45

50

55

60

65

70

75

80

0 45

75

80

Figure 45. Trace element Harker diagrams for southwest Fiordland plutonic rocks plotted against SiO2 with representative analyses from Paleozoic New Zealand plutonic suites.

The mafic rocks of the PIC including the Tdi, Oid, and Cuttle Cove Diorite show positive correlations with increasing SiO2 for Ba, Rb, and Zr and a negative correlation in Sr. Cumulate textures in both Tdi and Oid suggest accumulation of plagioclase and amphibole has occurred, which is supported by a positive Eu anomaly for one Oid sample [OU 75171, (Eu/Eu*)N = 1.07] (fig.48). LREE abundances are similar over the range of SiO2 in Tdi (fig.48). However, the lowest SiO2 (49 wt %) samples of Tdi are enriched in MREE to HREE relative to higher SiO2 sample (58 wt %).

103 1000

Rock/Primitive Mantle

100

10

1

Average I-type

Average S-type

Average A-type

Gleughearn Granite (Darran)

Separation Point Granite

Tmg OU75162

Tmg OU75100

Rp OU75107

Rp OU 75143

Ubg OU75170

Ubg OU75176

Lsg OU75175

0.1 Rb

Ba

Th

U

Nb

Ta

K

La

Ce

Pb

Sr

Y

P

Nd

Sm

Zr

Hf

Eu

Ti

Tb

Yb

Figure 46. Spider diagram of PIC granites normalised to primitive mantle. Also plotted are the average compositions of Sand I-type granites from the Lachlan Fold Belt (Chappell and White, 1992), and average A-types from around the world (Whalen et al., 1987, REE from Collins et al., 1982). Primitive mantle normalising values from Sun and McDonough (1989).

Figure 47. Sr/Y versus Y plot showing PIC data compared to fields of Darran Suite, SPS, and WFO. HiSY-LoSY dividing line is a suggested arbitrary division based on average lowest Sr/Y ratios in SPS and WFO.

104

Mafic rocks

rock/chondrite

1000 Lsg (OU 75175)

Cuttle Cove Gabbro

Tdi (OU 75111)

Tdi (OU 75154)

Tdi dyke (OU 75183)

Oid (OU 75169)

Oid (OU 75168)

Oid (OU 75171)

WFO 7

WFO 1

Darran Gabbro (DN2)

100

10

1 La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Figure 48. (top) Chondrite normalised REE plot of PIC granites. Fields of Separation Point Suite granitoids from NW Nelson (Muir et al., 1995) and Fiordland (Muir et al., 1998), as well as Darran Suite granitoids from Fiordland (Muir et al., 1998) are included for comparison. (bottom) Chondrite normalised REE plot of PIC mafic rocks. Darran gabbro and WFO values are plotted for comparison and are from Muir et al. (1998). Data for samples are listed in the appendices. Chondrite normalising values from Sun and McDonough (1989).

5.4.3 Compositional variation within PIC granites Incompatible behaviour of elements such as Th, and Y within PIC plutons and compatible behaviour of elements such as Al2O3, CaO, Fe2O3, TiO2, and MgO is consistent with liquidcrystal fractionation as a cause of compositional diversity within and between the units of the

105 PIC. The more extreme geochemical variation typical of S-Type granites (Chappell and White, 1992) is not apparent in the PIC, which probably reflects the fact that they were I-type primary magmas that approached their minimum-melt compositions. Internal fractionation is best displayed by smooth trends in Rp in all major and trace elements except K2O and Rb (figs.43 and 45). Trends of Rb/Sr versus SiO2 indicate eutectic-like crystallisation; that is a rapid increase in incompatible elements for little increase in SiO2. Feldspar fractionation as the main driver of chemical variation is supported by a negative correlation of Al2O3, CaO, Na2O, Sr and Ba with increasing SiO2 for these units. A log-log mineral vector diagram of Ba versus Sr shows a linear trend of Rp analyses projecting between plagioclase and alkali feldspar with increasing SiO2 (fig.49). As Sr is preferred in plagioclase and Ba in alkali feldspar, the trend is readily explained by co-crystallisation of plagioclase and alkali feldspar in constant ratios. As complete REE data was analysed for the Rp it was possible to check Eu anomalies against fractionation. The least fractionated Rp sample (OU 75107) has (Eu/Eu*)N = 0.94, while the most fractionated sample (OU 75143) has (Eu/Eu*)N = 0.57, which is consistent with the fractionation of plagioclase (fig. 48). Similar linear trends may be present in Tmg and Ubg, but sufficient data are not available to be certain. 10000

Rahu Plag

70

Cpx 50 1

Hbl

Opx

Darran Granitoid Separation Point granitoid WFO

3

1000

Treble Mtn Granite Revolver Pluton

Ba

30

100

Long Scarp Granodiorite Upper Blacklock Granite North Port Granite

50 %

10

Trevaccoon Diorite Cuttle Cove Diorite

70

Monk Granite

Kspar

Biot

1 10

100

1000

Only Island Diorite

10000

Lake Mike Granite

Sr

Figure 49. Southwest Fiordland analyses plotted on a Log-Log mineral vector diagram for Sr versus Ba. Representative samples of New Zealand Suites are included for comparison. Mineral vectors are marked in increments of 10, 30, 50 and 70% fractional crystallisation of that respective phase. Distribution coefficients are taken from Arth (1976) and Pearce and Norry (1979) for a liquid of rhyolitic composition.

106 1000

Separation Point granitoid Separation Point mafic WFO Darren Granitoid

HiSY

Darren mafic

100

Transitional Darren Suite Rahu

Sr/Y

Treble Mtn Granite Revolver Pluton

LoSY

North Port Granite Trevaccoon Diorite

10

Cuttle Cove Gabbro Long Scarp Granodiorite Lake Mike Granite Brothers Pluton Monk Granite

1 0.001

0.010

0.100

Rb/Sr

1.000

10.000

Upper Blacklock Granite Only Island Diorite

Figure 50. Sr/Y plotted against Rb/Sr showing possible effect of fractionation on Sr/Y ratio. Dividing line is an arbitrary separation of HiSY and LoSY compositions.

An important consequence of plagioclase fractionation is its potential effect on the Sr/Y ratio, which is best represented on a Sr/Y versus Rb/Sr plot (fig.50). Here, units of the PIC show linear negative trends of Sr/Y with increasing Rb/Sr ratio. This is consistent with the compatible behaviour of Sr in plagioclase, whereas Y and Rb are incompatible. Similar trends are evident in Npg, Ubg, and Bp, which also suggest fractionation of plagioclase. A dividing line separating granitoids of the Separation Point Suite from granitoids of the Darran Suite is drawn to distinguish rocks which may have high Sr/Y ratios through accumulation of plagioclase (most likely in more mafic rocks) as opposed to rocks which have high Sr/Y ratios as a result of conditions of melting at the source. As can be seen this does not coincide with the conventional discrimination of ‘adakitic’ or HiSY-LoSY signatures at Sr/Y = 40.

5.4.4 Oxidation state of PIC granites In an analysis of zircons from barren and ore-bearing felsic porphyries from the Chuquicamata-El Abra porphyry copper belt in Chile, Ballard et al. (2002) found that orebearing plutons had higher Ce(IV)/Ce(III) (>300) and (Eu/Eu*)N (40) in zircon. They interpreted the results as reflecting more oxidised magma compositions for ore-bearing porphyries. A similar analysis was undertaken by LA-ICP-MS of PIC zircons, the results of which are presented in figure 51. As noted previously in section 5.3.2.4, Ce/Ce’ and Eu/Eu’ are values normalised to bulk continental crust from Taylor and McLennan (1985). The Rp, Tmg, Ubg, and Lsg data has a fair degree of scatter, although zircon Ce/Ce’ ratios are on average greater than 300. In contrast to the observation of Ballard et al. (2002), Rp, Ubg, and

107 Lsg show apparent decreases in Eu anomalies with increasing oxidation. An explanation as to why this is the case is not offered. A high oxidation state of PIC granites is also supported by the frequently rich pink colour of alkali feldspar, which is inferred to be as a result of substitution of oxidised iron (Fe3+) for Al3+. Of particular note is the tighter distribution of data for Tmg which exhibits extensive hydrothermal alteration and hosts magmatichydrothermal mineralisation in the Tarawera Mine. 1000

1000

Revolver Pluton

100

Treble Mountain Granite

Ce/Ce'

Ce/Ce'

100

10

1 0.00

10

0.20

0.40

0.60

Eu/Eu'

1000

0.80

1 0.00

1.00

0.20

0.40

0.60

Eu/Eu'

1000

Upper Blacklock Granite

100

0.80

1.00

Long Scarp Granodiorite

Ce/Ce'

Ce/Ce'

100

10

1 0.00

10

0.20

0.40

0.60

Eu/Eu'

0.80

1.00

1 0.00

0.20

0.40

0.60

Eu/Eu'

0.80

1.00

Figure 51. Zircon Ce/Ce’ versus Eu/Eu’ for individual grains analysed from Revolver pluton and Treble Mountain Granite. Solid circles are analyses included for dating.

5.4.5 HFSE decoupling in Revolver Pluton High field strength elements (HFSE) in Rp show consistent decoupling in element pairs that are normally analogues of each other in most melting and crystallisation processes. This is evident in Y-Yb, which are plotted as normalised Y/(8.Yb) as the concentration of Yb is about 8 times less than Y (fig.52). It can be seen that the Y/Yb ratio in other PIC rocks as well as the SPS and Darran Suite remain constant with increasing SiO2 varying a little each side of 1,

108 thus adhering to the earlier qualification that these elements are normally analogues of one another. In contrast, Rp shows a decoupling with Y increasing up to 5 times that of the normalised Yb value. This decoupling is also evident in the Zr/Hf ratio, which in Rp are between 40 and 50, compared to a normal ratio between 30 and 40 (fig.52). A clue to the HFSE decoupling may be held in the 10000.Ga/Al ratio, which shows a linear increasing trend with increasing SiO2 in contrast to most rocks where this ratio remains constant (fig.52).

3.0

6 Revolver Pluton Tmg Ubg Darran granitoid SPS granitoid

Y/(8.Yb)

4

2.5

10000.Ga/Al

5

3

2.0

2

1.5 1

1.0

0 62

64

66

68

SiO2

70

72

74

76

62

64

66

68

SiO2

70

72

74

76

accumulation of zircon

70 60

Zr/Hf

50 40

Chondrite Zr/Hf = 38

30 20

fractionation of zircon

10 0 62

64

66

68

70

72

74

76

SiO2

Figure 52. Plots showing decoupling of HFSE elements that normally behaves similarly to each other chemically for Revolver Pluton.

5.4.6 Oxygen isotopes Oxygen isotope results for the PIC provide important constraints on its petrogenesis. Quartz, feldspar, and whole rock were analysed as described in section 5.2.3 for Lsg, Rp, Ubg, and Tmg. Additional whole rock δ18O values were determined for the Tdi and Oid, and two samples of local metasedimentary rock were determined to constrain the isotopic composition of potential supracrustal contaminants. The results of each individual rock analysis are discussed below and the summarised in Table 9. 5.4.6.1 Long Scarp Granodiorite The difference in δ18O values 7.5 for feldspar and 9.5 for quartz are greater than that expected for magmatic fractionation, which suggests that feldspar has undergone a small amount of

109 isotopic exchange after crystallisation. This amount of exchange is unlikely to have shifted the isotopic composition of the quartz due to its much slower rate of exchange (Gilletti, 1986). Using the approach of Faure and Brathwaite (2006) described in section 5.3.2.5, a magma δ18O value of 8.5 is estimated. 5.4.6.2 Revolver Pluton Quartz and feldspar δ18O values within the Revolver Pluton give a magmatically reasonable difference of 0.9. This was the only sample where a reasonable solidus temperature was obtained. Using the fractionation factor of Chacko et al. (2001) for quartz-albite a minimum temperature of 748oC is calculated. This is similar to the calculated zircon saturation temperature of 775oC and is a consistent true solidus temperature for a rock of 73 wt % SiO2. Again a magma δ18O value of 7.3 is estimated from the quartz δ18O of 8.3. 5.4.6.3 Treble Mountain Granite The δ18O values of -4.6 for feldspar and 2.9 for quartz are much lower than observed in unaltered igneous rocks and indicate that these minerals have exchanged oxygen with meteoric water at relatively high temperatures. This is consistent with petrographic observations (section OU 75108) which show extensive hydrothermal alteration. No magma δ18O can be inferred as even the quartz which is normally relatively inert to shifts in δ18O, has exchanged oxygen as it is at least 3 ‰ less than what would be expected. 5.4.6.4 Upper Blacklock Granite The feldspar in Ubg gives an extremely low δ18O value of -0.2, which suggests it has exchanged oxygen with meteoric water at high temperature. The quartz value of 8.2 is similar to that of Rp. Based on its lithological similarities to Rp the quartz δ18O is likely to be close to its original composition, although the large negative shift in the feldspar δ18O means a slight shift in the quartz δ18O cannot be discounted. Thus, magma δ18O value of 7.9 must be considered a minimum value for the Ubg. 5.4.6.5 Trevaccoon Diorite A whole rock δ18O value of 5.5 for a gabbroic sample of Tdi maybe negatively shifted based on a measured feldspar δ18O of 4.1, which is slightly lower than expected for feldspar. Petrographic evidence for alteration of feldspar in this sample is minor compared to Tmg. Thus, a whole rock δ18O value of 5.7 could be close to the magma δ18O, although a slight negative shift is possible. The low δ18O value of 5.5 is similar to MORB (5.7 ± 0.3, Gregory and Taylor, 1981), and is consistent with a primitive mafic magma or cumulate.

110

Table 9. Oxygen isotopic compositions of quartz, feldspar, and whole rock for selected samples from southwest Fiordland

Sample

δ18O qtza

Long Scarp Granodiorite (OU 75175) Revolver Pluton (OU 75143) Upper Blacklock Granite (OU 75176) Treble Mtn Granite (OU 75108) Trevaccoon Gabbro (OU 75111) Only Island Diorite (OU 75171) Preservation Fm. psammite (OU 75166) Cameron Group psammite (CAM 08)

δ18O fspa δ18O wrb δD wrb δ18Omagma

9.5

7.5

7.4

8.5c

8.3

7.4

8.8

7.3c

8.9

-0.2

5.9

7.9c

2.9

-4.6

-

-

4.1

5.5

5.5d

-

-

7.1

7.1d

-

-

12.4

-

-

18.2

-104

ToCe

748

-

a Analyses are on quartz and feldspar, which were hand picked from rock crushed to 250 μm. Isotopic analysis was done by Dr. K. Faure (Stable Isotope Facility, GNS Science, New Zealand). b

Samples processed from whole rock powders by Prof C. Harris (Geological Sciences, University of Cape Town, South Africa).

c

δ18Omagma estimated by subtracting 1.0 from δ18OQtz (Faure and Brathwaite, 2006) for Δ18Oqtz-melt closed system cooling effects.

d 18

δ Omagma estimated to be equal to δ18O wr

e

Temperature is calculated using the following formula: 1000lnα = A x 106/T2 , where Aqtz-ab=0.94 (Chacko et al., 2001).

5.4.6.6 Only Island Diorite Again petrographic evidence for alteration of Oid feldspar is limited and the whole rock δ18O value of 7.1 obtained is probably close to its original magma composition. The slightly higher value of Oid compared to Tdi either indicates a crustal component in Oid or alteration of Tdi.

5.4.7 Neodymium isotopes Neodymium isotopic composition was measured by TIMS on whole rock samples crushed at the University of Otago by Dr. M. Norman of PRISE at the ANU for the Oid, Tdi, Rp, Tmg, Ubg, and Lsg (results reported in Table.10). The Tdi, Oid, and Rp have uniform εNd values of +1.3, while the Tmg and Ubg have values of +0.7, and Lsg a value of -0.3. The propagation of errors during data reduction gives an error in εNd of ± 0.3, which means all units except Lsg can be considered within analytical error of each other. The most notable aspect of the results is the lack of variation in εNd between all samples given their compositional range from gabbro to granite.

111 Table 10. Summary of Nd isotopic results showing measured values and calculated values

Sample Long Scarp Granodiorite (OU 75175) Revolver Pluton (OU 75143) Treble Mountain Granite (OU 75100) Trevaccoon Diorite (OU 75154) Upper Blacklock Granite (OU 75176) Only Island Diorite (OU 75171)

Age (Ma)

Sm (ppm) Nd (ppm)

147

Sm/144Ndm

143

Nd/144Ndm Initial εNd

133

11.52

93

0.0746

0.512518

-0.3

132

2.33

13.82

0.1015

0.512624

+1.3

130

3.58

19.44

0.1109

0.512602

+0.7

128

6.11

31.95

0.1151

0.512637

+1.3

126

3.37

24

0.0845

0.51258

+0.7

122

3.16

15.2

0.1252

0.512645

+1.3

* Initial Nd isotopic ratios were estimated by correcting for radiogenic growth of 143Nd using measured Sm/Nd (LA-ICPMS, see appendices 10.1.7) and zircon ages (section 4.3). The initial Nd isotopic ratios are reported as εNd values where εNd = [((143Nd/144Nd)i / (143Nd/144Nd)CHURt )-1] x 104.

5.4.8 Emplacement Pressure 5.4.8.1 Qualitative constraints The lack of miarolitic cavities in PIC plutons suggests that they were not intruded at shallow crustal levels (probably deeper than 3 km). Further constraints on pressure include the lack of metamorphic overprint on the plutonic rocks and metasediments, which indicates that they were intruded into crust that was not under any directed stress and temperature not greater than 400oC, at which point feldspar becomes ductile. This temperature constraint equates to depths of less than 10 km under expected maximum upper continental arc geotherms of 40oC/km (Rothstein and Manning, 2003). The isotopic constraints on alteration of Tmg discussed in section 5.3.9.3 support meteoric water as the primary fluid involved in hydrothermal alteration. As this requires open pathways to the surface, it must have occurred above the brittle-ductile transition at temperatures and pressures indicated by the lack of ductile deformation. The lack of miarolitic cavities, combined with constraints imposed by evidence for meteoric water alteration and lack of ductile deformation, suggest intrusion occurred at minimum depths of approximately 3 km and a maximum of 10 km. 5.4.8.2 PIC analyses in the H2O-saturated Ab-Or-Qtz system – potential constraints Normative compositions of PIC granites corrected according to Blundy and Cashman (2001) are plotted in the H2O-saturated Ab-Or-Qtz phase diagram in figure 53. In this method, compositions are required to have normative anorthite no greater than 20 wt % and corundum no greater than 2 wt %. PIC compositions not fulfilling these requirements are marked with a cross. Biotite is not included in the normative calculations so K2O in biotite is allocated to orthoclase with a reduction in normative quartz. Thus, rocks containing significant modal

112 biotite may have their normative compositions shifted towards Or and away from Qz. As PIC granites are leucocratic I-type granites containing minimal biotite, compositions are not likely to have been shifted significantly. Discounting the one sample with normative corundum greater than 2 wt %, the rest of the Rp data follows a plagioclase fractionation trend to the 300 MPa quartz + plagioclase cotectic and then moves along the minimum providing a constraint on pressure at which it crystallised. Rapid increase in incompatible trace elements for little increase in SiO2 in Rp possibly is supportive of near ternary minimum crystallisation. The Ubg also looks to follow a plagioclase fractionation trend that looks to intercept the 200 MPa quartz + plagioclase cotectic where two analyses look to trend along the cotectic towards the minimum indicating a maximum pressure of equilibrium. In both cases, final emplacement pressure may have been lower. Compositions of Tmg excluding two analyses with normative corundum greater than 2 wt % appear to cluster around the 500 MPa quartz + plagioclase cotectic. This appears to be high even as a maximum crystallisation pressure for this unit.

Upper Blacklock Granite Long Scarp Granodiorite Brothers Pluton Lake Mike Granite Monk Granite Revolver Pluton Treble Mtn Granite

Figure 53. Diagram showing PIC granite normative sample compositions plotted in the H2O-saturated Qtz-Ab-Or- system corrected for normative An according to Blundy and Cashman (2001). Compositions not conforming to the requirement of having less than 20 wt % normative anorthite, and 2 wt % normative corundum are highlighted with a cross.

5.4.8.3 Phengite geobarometry Phengite geobarometry is dependant on the Tschermak substitution in the muscoviteceladonite solid solution in the presence of an appropriate mineral assemblage. The system was most recently assessed by Simpson et al. (2000) in the KMASH and KFASH systems

113 involving combinations of phengite, chlorite, biotite, K-feldspar, quartz, and H2O, and requires a buffer assemblage including quartz, H2O, and any two of biotite, chlorite, Kfeldspar. These authors contoured isopleths of the Tschermak substitution and plotted them in a P-T projection (fig. 54B). A hornfels metasediment (OU 75126) was collected 2m from a contact with Rp in Brokenshore Bay. The sample satisfies the required assemblage having quartz, K-feldspar, biotite, chlorite, and muscovite. Chlorite shows petrographic evidence of having formed as a retrogression product of biotite. Muscovite compositions were measured by EMP and are listed in Appendix 10.3.3. Metasedimentary white micas have compositions towards muscovite and are not thus true phengite in composition (fig. 54A). No analyses display ideal Si/Al ratios as would be expected by Tschermak substitution alone for pure muscovite (fig.54A). Compositions range in Si cations per formula unit (p.f.u) from 3.06 to 3.17. The analysis with 3.17 Si cations p.f.u. could be interpreted as the only equilibrium composition not shifted significantly during retrogression, and may explain why the cluster of analyses closer to 3.06 Si cations p.f.u have a greater tendency not to satisfy the ideal Si/Al ratio. Assuming this single analysis represents equilibrium; it is plotted in the [CHL] absent system in figure 54B. Considering the sample has an Mg/Fe ratio of 0.5 in muscovite (phengite), a minimum pressure of approximately 1 kbar, which is equivalent to a depth of about 3 km based on average crustal densities. Extrapolating the Si p.f.u = 3.17 contour to 700oC yields a pressure estimate of c.5 kb or 15 km depth. The phengite composition is thus consistent with other constraints on emplacement pressure. B

3.25

A

Metasediment OU 75126

Muscovite

3.00

Al pfu

2.75

2.50 Phengite

2.25

2.00 Celadonite

1.75 2.75

3.00

3.25

Si pfu

3.50

3.75

Figure 54. (A) Muscovite compositions (Al per formula units plotted against Si per formula units) plotted for metasediment from Rp contact (OU 75126). (B) P-T projection for the KMASH-KFASH system with chlorite absent (From Simpson et al., 2000). The black region represents the projected composition of the hornfelsed metasediment OU 75126.

114

5.4.9 Discussion The near constant initial εNd values of the PIC across such a broad compositional range tightly constrain the possible petrogenetic models for the PIC. These initial εNd values effectively require all PIC rocks to have a shared source. The compositional range then may result from different degrees of partial melting or fractional crystallisation. An evaluation of the evidence for these two models follows. 5.4.9.1 Granite major and trace chemistry Chondrite normalised REE patterns show that the higher SiO2 rocks of the PIC lack deep Eu anomalies and in some instances show positive Eu anomalies (fig.55). Such features preclude a simple fractional crystallisation scenario in which mafic rocks originate as cumulates and felsic rocks as more evolved liquids. Instead, they indicate that the PIC formed from magmas with a range of compositions modified by late plagioclase fractionation. The range of Eu anomalies within units could represent progressive melt-out of a source leaving varying amounts of residual plagioclase or late fractionation of plagioclase as mentioned. The slight positive Eu anomalies present in some units would support some late fractional crystallisation of plagioclase and subsequent accumulation. Thus, petrologic variation exhibited between plutons most likely results from differences in melting conditions or source composition. 1.5

Eu/Eu*

1.0

0.5 Rp

Tmg

Ubg

Lsg

Tdi

Oid

0.0 45

50

55

60

SiO2

65

70

75

80

Figure 55. PIC rocks plotted with Eu/Eu* versus SiO2.

Relatively high ASI and normative corundum of PIC granites compared to I-type granites documented by Chappel and White (1992, original discrimination of I- and S- type granites at normative corundum = 0.8 and ASI = 1.02) is most likely the result of residual amphibole in the source. As amphibole is strongly metaluminous (ASI 100), combined with high Al2O3 and Na2O it can confidently be designated as adakitic or HiSY (Defant and Drummond, 1990, Tulloch and Kimbrough, 2003). With more mafic compositions it should be kept in mind that plagioclase accumulation could produce a high Sr/Y ratio from a parental magma with an initially low Sr/Y ratio and thus magmas with cumulate textures or positive Eu anomalies should be assessed more critically (Stevenson et al., 2005). 5.4.9.2 Mafic rock major and trace chemistry Samples of Tdi and Oid with the lowest SiO2 are characterised by cumulate textures and assemblages rich in hornblende and plagioclase +/- biotite. Tdi and Oid include higher SiO2 equivalents. That Tdi shows evidence of being at least in part cumulate in origin, while showing generally high K2O for such a low SiO2 composition suggests cumulate grains trapped more evolved liquid. The composition of the most mafic Tdi and Oid samples is unlikely to reflect the primary melt composition based on them being at least partially cumulate in origin. The enrichment in MREE and HREE in lower SiO2 Tdi samples is contrary to that expected by fractionation as REE elements typically behave as incompatible

118 elements and their concentrations increase in the differentiated melt. This enrichment is consistent with the accumulation of hornblende in the most mafic Tdi samples, as hornblende has a high affinity for MREE and HREE elements. The presence of magmatic epidote in Tdi, Lsg, and possibly Oid supports a deep source for their parental magmas as experimental evidence suggests magmatic epidote crystallises at pressures greater 0.5 GPa in a tonalitic melt (Schmidt and Poli, 2004). This would correspond to depths of greater than 18 km under typical crustal densities. A source depth > 18 km could be consistent with melting in an amphibolite lower crust. 5.4.9.3 Oxygen isotopes O’Neil (1977) used oxygen isotopes as a tool for ascertaining magma genesis in the New England Fold Belt of Eastern Australia, recognising that I-type magmas have δ18O ≤ 10. An extensive study of granites in the Lachlan Fold Belt by Chappell and White (2001) reinforced this conclusion. These authors attributed the higher δ18O observed in S-type granites to a sedimentary source that had undergone weathering and hence isotopic exchange at the Earth’s surface. High temperature exchange with meteoric water results in negative shifts in a rock’s whole rock δ18O composition. These shifts are readily identified when the δ18O of quartz and feldspar are measured together as a larger difference in δ18O between quartz and feldspar results as feldspar exchanges oxygen at a faster rate. More studies now analyse both feldspar and quartz so a magma δ18O value can be more accurately interpreted. In New Zealand δ18O values of igneous rocks have not been reported widely in the literature. Calculated magma δ18O for the Rahu, Darran and SPS are reported by Tulloch and Kimbrough (2003) and are presented in Table 1. Both the Darran Suite (δ18O = 4.4-7.6) and the SPS (δ18O = 5.8-6.9) have relatively primitive I-type signatures, which limits the involvement of sedimentary rock. Slightly higher values for Rahu Suite rocks (δ18O = 9.410.0) are consistent with the suggestion by Tulloch (1983) that Rahu Suite magmas are SPS magmas that have assimilated significant amounts of Paleozoic sedimentary rock. Inferred magmatic δ18O values for PIC granites that lack evidence of subsolidus oxygen exchange give values between 7.3 and 8.5. These values are similar to the Darran and SPS and are consistent with I-type magma compositions and require minimal involvement of sedimentary rock in their genesis (southwest Fiordland whole rock δ18O values = 12-18, Table 9). Whole rock values for the Tdi and Oid of 5.5 and 7.1 are also consistent with I-type genesis. In contrast to the other PIC rocks, the Tmg shows an extreme negative shift in its quartz oxygen isotopic composition. This shift is reflected to lesser degrees in feldspar composition

119 in the other PIC rocks. Presuming the negative shifts in δ18O of PIC rocks are caused by meteoric water of the same isotopic composition, which is not unreasonable considering their close proximity and similar age, it is possible to plot δ18Oqtz versus δ18Ofsp to detect if there is a pattern in isotopic shift (fig. 56 B). Barnett and Bowman (1995) showed that characteristic parabolic data arrays resulted when quartz and feldspar exchanged oxygen with an isotopically constant reservoir due to kinetic differences in their exchange. Initially severe negative shifts in feldspar occur as it exchanges oxygen with the water at a much faster rate than quartz resulting in a near vertical decreasing trend on a δ18Oqtz versus δ18Ofsp plot. When feldspar has equilibrated with the isotopic composition of the water, no further decrease in its δ18O can occur. However, quartz will continue undergoing exchange, and eventually a marked negative shift in quartz will be seen, which results in a concave parabolic trend on a δ18Oqtz versus δ18Ofsp plot for a group of samples that have exchanged oxygen in varying proportions. The point at where quartz shows marked negative shifts while feldspar remains essentially unchanged, should represent the equilibrated feldspar-water δ18O. The feldspar δ18O can then be used to calculate the δ18O water if the exchange temperature can be reasonably constrained. It is apparent in figure 56B that Tmg quartz has undergone a substantial negative shift, at a point where feldspar is no longer showing the same negative trend. This indicates that the feldspar in Tmg has equilibrated with meteoric water and can no longer be shifted negatively. It seems likely that the negative δ18O shifts in Lsg, Ubg and Tmg are caused by the same meteoric water as indicated by the fact that the samples define a concave parabolic trend in figure 54, which as mentioned is expected for isotopic exchange with meteoric water of fixed composition. To calculate water δ18O from mineral values, a temperature must be assumed. It would be unlikely that meteoric water would be able to infiltrate rock below the brittle-ductile transition and thus an upper limit of 300oC is assumed based on a quartzo-feldspathic upper crust. The alteration assemblage of sericite + chlorite + epidote is inconsistent with low-temperature near surface weathering, so a lower temperature limit of 200oC is probably reasonable. Assuming these temperatures and using a δ18O value of -5 for feldspar from Tmg, the meteoric water which caused the alteration has a calculated δ18OH2O of -9 to -13 between 200 and 300oC using equilibrium fractionation factors of Matsuhisa et al. (1979b). This range agrees well with an independent estimate of Cretaceous meteoric water δ18O of – 11 that can be calculated from the δ18O value for a surficial kaolinite reported in Chamberlain et al. (2000).

120 A whole rock δD of -104 was measured from the same Tmg sample from which oxygen isotopes were measured. On the assumption that unaltered quartz has a δ18O value of +9 and fully altered quartz would be about -2 based a feldspar δ18O value of -4, a water-rock ratio >> 1 is estimated for Tmg based on its quartz value of +3 which is approximately halfway. From this we assume that the whole rock δD of -104 has equilibrated with the meteoric water as hydrogen isotopes approach exchange equilibrium at much lower water-rock ratios. Based on equilibration temperatures between 200 and 300oC and the larger analytical uncertainties involved in measuring whole rock δD, a water δD value of -75±15 is calculated. The calculated δ18O and δD of the water responsible for alteration of Tmg is plotted on a δD versus δ18O diagram with a line that defines the variation in isotopic composition of meteoric water after Hedenquist and Lowenstern (1994). The calculated isotopic composition of the water responsible for alteration of Tmg can be seen to overlie meteoric water compositions. This argues that the region was subaeriallly exposed during alteration to permit recharge of meteoric water.

b

Figure 56. (a) δD versus δ18O plot. Treble Mountain Granite (OU 75108) water δD and δ18O values are calculated from whole rock, with the box representing uncertainties in the measurement and alteration temperature. H- and O- isotope data fields and diagram after Hedenquist and Lowenstern (1994) (b) Plot of δ18OQtz versus δ18Ofsp for Cretaceous granitoids from the PIC. Revolver Pluton has quartz and feldspar values which are likely to represent the magmatic values and has not been shifted by interaction meteoric fluid. Conversely, both feldspar and quartz in Treble Mountain Granite have exchanged oxygen with meteoric water at high temperature. The feldspar in Tmg is inferred to have reached isotopic equilibrium with that water.

The degree of hydrothermal alteration observed in the Rp and Tmg qualitatively correlate with shifts towards lighter δ18O requiring the involvement of meteoric water. The regional

121 occurrence to which hydrothermal alteration has taken place suggests that the hydrothermal system was extensive. The most extreme hydrothermal alteration and negative isotopic shifts are observed in samples of Tmg and Ubg. Two main factors control meteoric fluid flow within the crust; heat, in this case provided by shallow plutons and structural permeability, both of which could be consistent with an oblique continental arc setting. 5.4.9.4 Neodymium isotopes As already emphasised PIC, initial εNd values span a narrow range between -0.3 and +1.3, which is unusual considering the compositional range of the samples. Darran Suite rocks from Eastern Fiordland also have a wide compositional range, but narrow initial εNd range between +3.4 and +4.5. SPS plutons from eastern Fiordland have initial εNd values of +3.2 to +3.3, which are noted as being similar to the Darran plutons that they intrude (Muir et al., 1998). In contrast, SPS plutons of the Separation Point Batholith (SPB) in Northwest Nelson have initial εNd significantly lower than SPS correlatives in eastern Fiordland with values between +1.2 and +1.8 (Muir et al., 1995). However, SPB initial εNd values fall within the range obtained for the WFO of -0.4 to +2.7 (McCulloch et al., 1987, Muir et al., 1998), which are also notably lower than Darran and SPS rocks from eastern Fiordland. Sampling of multiple plutons in the SPB by Muir et al. (1995) suggests it has a limited isotopic range. Nd-isotopes obtained from the Te Kinga and Deutgam sub-suites within the Hohonu Batholith, and Buckland Granite, all correlated with the Rahu Suite, give initial εNd values from -4.4 to -6.1 (Waight et al., 1998b). These more evolved values are consistent with these authors suggestion that Rahu Suite compositions are mixtures of SPS and Greenland Group greywacke (εNd at 110 Ma of -10.7 to -13.1, Waight et al., 1998b). It is noted that initial εNd values for the PIC are very similar to those of the SPB and WFO, all of which have clear intrusive relationships with Western Province crust, while being consistently lower than Darran and SPS plutons from eastern Fiordland.

5.4.10 Petrogenesis 5.4.10.1 Current petrogenetic models for granites Granite petrogenesis experienced a defining point when Tuttle and Bowen (1958) showed that natural granitic compositions plotted close to the H2O-saturated ternary minimum in the AbOr-Qtz-H2O system at low to moderate pressure. This finding ended years of debate as to whether granites formed in-situ from metasomatic processes, or were crystallised magmatic liquids that had moved upward through the crust from a deeper source. The swing may have been too far towards the ‘all granites crystallised from a liquid’ hypothesis. It is now

122 recognised that granitic magmas are rarely H2O-saturated in their source, and many were never pure liquids as indicated by the presence of inherited zircons. Much debate is now focused around the common reality that not all granitic suites can be explained simply by fractional crystallisation from a parental melt. Two contrasting views are currently under debate to explain composition, petrography and zircon inheritance. The first is the restite-unmixing model (Chappell and Wyborn, 2004, Chappell et al., 1987, Chappell et al., 1999), which Chappell and co-authors developed to explain compositional variation in most of the plutonic suites in the Lachlan Fold Belt (LFB) of southeastern Australia. The restite-unmixing model considers that unmelted, but magmatically equilibrated, source material (restite) may be entrained by a partial melt, and that compositional variation in plutons and suites results from the variable separation of these two components from the resulting magma. The second is the magma-mixing model, which is self explanatory. Compositional variation within LFB suites is explained by mixing of deep basaltic melts and partial melts of older greenstone basement at deep levels with variable contamination by upper crustal metasedimentary rocks (Collins, 1998, Keay et al., 1997). 5.4.10.2 PIC petrogenesis Petrogenetic models for the PIC must account for its compositional diversity, relatively primitive and uniform isotopic characteristics, arc-like trace element signatures, and temporal and chronological relations with the Darran Suite, WFO and SPS. There is a broad consensus that major element compositions of granites reflect their source and melting conditions. Although the lowest SiO2 rocks of the PIC have evidence of crystal accumulation, it is unlikely that both the felsic and mafic rocks of the PIC are related to a single intermediate parental magma composition through fractional crystallisation. Instead, a model is favoured which accounts for the full compositional range by fractional crystallisation of parental magma of variable compositions. The following arguments are offered against a common intermediate parental composition. Firstly, Eu anomalies in Figure 55 do not define a uniformly decreasing trend with increasing SiO2 as would be expected for a single parental magma undergoing fractional crystallisation of plagioclase. Secondly, dykes occur on the periphery of Tdi and Oid that most likely represent liquid-rich compositions of these units and argue against them being wholly cumulate in origin. The lack of inherited zircons of the same age populations found in PIC granites argues that Tdi and Oid were high temperature liquids undersaturated in zircon. Although Oid contains inherited zircons these were all sourced from the local country rocks. In contrast, the abundance of inherited zircons in the PIC granites

123 argues that their parental magmas were saturated in zircon at an early stage, which is evidenced by a zircon fractionation trend on a Zr versus SiO2 plot (fig.57). 450

Treble Mtn Granite

400

Revolver Pluton

350

Trevaccoon Diorite

Upper Blacklock Granite Cuttle Cove Gabbro

300

Long Scarp Granodiorite Monk Granite

250

Zr

Only Island Diorite

200 150 100 50 0 45

50

55

60

SiO2

65

70

75

80

Figure 57. Harker diagram of Zr versus SiO2 for PIC rocks.

The proposed hypothesis requires that the PIC source was capable of yielding both higher and lower SiO2 parental magmas, each of which underwent at least some fractional crystallisation to produce the full range of compositions seen at the higher and lower SiO2 end of the spectrum in the PIC. The PIC plutons mirror the somewhat bi-modal compositional range of the eastern Fiordland Darran Suite and also share a restricted range of εNd initials, although values from eastern Fiordland Darran Suite plutons are slightly higher being from +3.4 to +4.5. This compositional and isotopic range is also reflected in the SPB/WFO, although its range of initial εNd values is very similar to that of the PIC. Although these initial εNd values are theoretically not difficult to obtain through mixing of a more primitive magma with older crustal materials, it is highly improbable that such a process could yield a wide compositional range with the same εNd value. The lower initial εNd values of the PIC (-0.3 to +1.3) compared to eastern Fiordland Darran Suite rocks (+3.4 to +4.5) is not consistent with direct equivalents of Darran rocks as a possible source. On the other hand more inboard, deeper equivalents of the Darran Suite were a significant source component, and potentially, in the case of the PIC granites, the sole source. The constraints discussed require that the PIC source reservoir had a narrow εNd range (probably < ±2 εNd units). This requires that the source region below the PIC is different from rocks exposed at the surface, which are made up of early Paleozoic metasedimentary rocks and mid-Paleozoic and Mesozoic plutons of variable isotopic and chemical composition. The presently exposed crust

124 is incapable of yielding magmas with isotopic uniformity seen in the PIC, or indeed the WFO and SPB. Unlike the PIC, dating of the SPB and WFO by SHRIMP U-Pb zircon dating has provided little evidence of zircon inheritance coeval with Darran Suite magmatism (Muir et al., 1994, Muir et al., 1998). The only documented exception is c. 136 Ma HiSY dykes cutting earlier Darran Suite rocks in northern Fiordland (Wandres et al., 1998). These dykes contain zircons coeval with earlier Darran Suite ages and are inherently similar to PIC granites in this respect, although they could be from locally derived country rock unlike PIC. The significantly more elevated initial εNd, combined with a lack of obvious Gondwanan zircon inheritance, and enrichment in trace elements such as Rb in the PIC contrast with the clearly contaminated Rahu Suite. This suggests that the assimilation of Greenland Group greywacke or similar isotopically more evolved Paleozoic metasedimentary rocks by PIC magmas did not occur. The similarity of εNd values within the PIC, and their similarity to the isotopic range within the SPB and WFO lead to the logical conclusion that these rocks share the same source, or in the case of some PIC magmas may result through recycling of slightly earlier plutons from such a source. Mattinson et al. (1986) previously noted that the Pb isotopic compositions of WFO and Darren Suite rocks formed a tight cluster and suggested that the two suites shared a very similar source. Muir et al. (1998) proposed that WFO magmas could result from 30% melting of a Darren mafic composition. They added that SPS magmas could be the result of lesser degrees of partial melting, possibly about 20%, or greater amounts of residual garnet in the source. From this, these authors suggested that western Fiordland could be underlain by equivalents of the Darren Suite. Muir et al. (1998) do not address why the WFO and SPB have consistently lower εNd than the eastern Fiordland SPS and Darran Suite plutons. Unfortunately, Nd isotopic data for Darran Suite rocks in Western Fiordland, Stewart Island, and Northwest Nelson are not yet published, which means it cannot be established whether inboard equivalents of the Darran Suite clearly intruding Western Province basement have lower εNd values, although based on εNd values obtained from the PIC and SPB/WFO, I would suggest that this would be expected. 5.4.10.3 Potential sources McCulloch et al. (1987) noted that the Sr- and Nd isotopic systems in WFO were decoupled as no decrease in 87Sr/86Sri was evident for marked decreases in εNd. Thus, the variation in εNd could not be explained by contamination with Paleozoic metasedimentary rocks, as a concomitant shift in 87Sr/86Sri would be expected due to the latter’s high Rb/Sr ratios. It was

125 concluded that lower than expected εNd values obtained for granulites and ultramafic rocks required a mid-Paleozoic mafic source with a low Rb/Sr ratio, which they suggested could be geochemically similar to leuco-gabbros from the Darran Complex. Muir et al. (1995) also concluded that εNd values of the SPB were too low to be sourced from a contemporary MORB type mantle, and excluded contamination by Paleozoic metasedimentary rocks. These authors concluded that the evidence required an amphibolite source that had separated from depleted mantle between 500 and 600 Ma, which they referred to as the Separation Point Depleted Mantle (SPDM). A similar argument was made by Waight et al. (1998b) for the Deutgam and Te Kinga sub-suites of the Hohonu Batholith in which mixing of magmas from SPDM with Greenland Group greywacke was considered most likely. Mixing of a recently derived MORB-like mantle component was thought unreasonable as it required mixing of 50 % Greenland Group to achieve isotopic compositions of the I-type Deutgam Suite member. The PIC is similar in its source requirements to the SPB and WFO. It too has no evidence of a Paleozoic upper crustal component, which suggests the isotopic composition of the PIC, like that of the SPB and WFO, reflects that of an old, deeper mafic protolith. The source of PIC magmas in its simplest form has to be a mixture of an isotopically aged mafic component with a component of inherited zircons from deeper inboard equivalents of the Darran Suite. The requirement for an isotopically aged mafic component suggests that the Western Province metasediments are underlain by Early Paleozoic amphibolite as previously proposed for the genesis of the WFO by McCulloch et al. (1987), and proposed later for the SPB by Muir et al. (1995). An Early Paleozoic amphibolite could have evolved to a εNd value of +1 by the Cretaceous (calculated using reasonable present day 143

147

Sm/144Nd = 0.15 and

Nd/144Nd = 0.51265). This alternative is consistent with the lower εNd values in the SPB,

WFO, and PIC, which could be explained by inboard plutons of the Median Arc tapping older mafic lower crust with a more evolved isotopic composition. Higher initial εNd values of eastern Fiordland Darran and SPS suite rocks may represent more outboard magmatism of the Median Arc, which is probably underlain by a younger mafic lower crustal component such as Brook Street arc crust. The existence of an Early Paleozoic lower crust composed of amphibolite for the Western Province is further provided by widespread I-type magmatism in the Paleozoic (Paringa and Tobin suites). Published Nd-isotopic data for Paleozoic I-type rocks is limited to the Riwaka Complex and Zetland Diorite (Muir et al., 1996b). The Riwaka Complex, a layered ultramafic intrusion has initial εNd = +1.9, while the Zetland Diorite has initial εNd = -0.3. Both of which are consistent with contemporary derivation from an early Paleozoic mantle-derived mafic source.

126

Chapter Six

6 TARAWERA MINE The Tarawera Mine base-metal lode was incorporated into this study as its occurrence within Treble Mountain Granite made an association between the mineralisation and Cretaceous plutonism a likely possibility. The findings of a petrographic, geochemical, and isotopic study of the mineralisation support the hypothesis for a genetic link between the base-metal mineralisation and Cretaceous plutonism providing some of the most convincing evidence to date for such mineralisation in New Zealand.

6.1 Analytical techniques XRD and hydrogen and carbon isotopic analysis were adopted in analysing Tarawera Mine samples in addition to the use of the electron microprobe and oxygen isotope analysis by laser assisted fluorination, which are discussed in section 5.2. The methodologies of XRD and hydrogen and carbon isotopic analysis are discussed in the following sections. The technique for dating of the host granite is discussed in section 4.1.

6.1.1 Mineral analysis (X-ray diffraction) Two altered wall rock samples were mounted on glass thin sections then ground to approximately 0.5 mm thick for mineral phase identification. The samples were analysed by Mr Damien Walls, Department of Geology, University of Otago, using a PANanalytical PW 3040/60 MPD system XRD with data reduction by HighScore (PANanalytical). The operating conditions were 40 kV and 30 mA using CuKα radiation over 3 - 80o2Th range. More detailed results and operating conditions are presented in appendix 10.2.

6.1.2 Hydrogen and carbon stable isotopic analysis Hydrogen and carbon isotopes were measured from fluid inclusions also by Prof. Chris Harris at the Department of Geological Sciences, University of Capetown. All the samples were treated the same way - cleaned in ethanol, dried at 110oC in an oven, degassed on line at about 150oC and decrepitated at 550oC for 5 minutes. Water was collected by thermal decrepitation of quartz fragments from the Tarawera lode. The same method was used for liberation of H2O from a whole rock powder from the Treble Mountain Granite. CO2 was collected during thermal decrepitation of Tarawera samples.

127

6.2 History The Tarawera Mine is located in a small bay on the western shore of Isthmus Sound. The lode is hosted in Treble Mountain Granite (Tmg) and was mined from its discovery in 1895 till 1897 when the workings were accidentally breached by the sea. Between 1908 and 1910 a smelter was constructed on site to process the refractory ore to recover silver and gold; however initial trials were unsuccessful and the mine was abandoned (Begg and Begg, 1973, and references therein).

6.3 Site description Scattered ore, Northwest Nelson marble [used in the smelting process, (Begg and Begg, 1973, and references therein)] and slag around the smelter site give evidence of the first smelting. A site map of the remaining workings is presented in figure 58. Workings consist of a small adit that was put in to meet a shaft approximately 5m below the collar. A shaft is positioned at the southeast end of a levelled terrace where the remains of a smelter are positioned. The bottom of the shaft is now collapsed just below where the adit meets it, thus excluding access to what were presumably the worked portions of the mine. Other workings include an open cutting, which exposes a quartz vein striking 154o approximately parallel to the shore, about ten vertical metres below the levelled terrace. A collapsed adit 40 m westwards along strike from the shaft, which is driven southwards (orthogonal to strike) was presumably attempting to intercept the lode worked from the shaft. Tarawera mineralisation is hosted in quartz veins which cut coarse grained porphyritic to variably sheared Tmg. Two distinct sets of quartz veins exist. The first set of quartz veins cut Tmg on the immediate foreshore, and consists of a swarm of veins (referred to as the foreshore veins herein, OU 75198) that have alteration halos of c.10 cm containing disseminated pyrite. The foreshore veins consist of two dominant veins and associated stringers running roughly parallel to the foreshore trending approximately SE-NW, with a dip of about 750 to the SW (OU 75198). The larger vein is about 7 cm thick, while the swarm is roughly constrained to 2 m in total width. The northwest end of the swarm terminates against a fault striking approximately east and dipping moderately to the south. No sense of offset could be determined due to the absence of exposed markers, although the lack of damage around the fault probably suggests it is a minor structure. A small fine-grained dioritic intrusion, which appears to be structurally controlled by SE-NW trending shear in the granite, is cut by the foreshore quartz vein swarm. The second sets of veins are wider than foreshore veins and have been the main focus of past mining.

128

Figure 58. Site map of Tarawera Mine showing position of workings and geological data (approximately to scale).

The main workings were presumably accessed by the shaft, since historical accounts say the mine was once abandoned for several years after the workings were breached by the sea. The proximity of the shaft to the sea and elevation of only about 10 m above sea level make this easy to imagine. As the shaft and lower workings cannot be accessed, ore samples were dominantly collected from the smelter site for analysis and are herein referred to as the lode quartz (OU 75199). The adit reveals massive granite with brittle shear bands approximately parallel to a prominent joint set oriented 130/74 NE. One sulphide bearing quartz vein is exposed in the adit. It is 8 cm thick and is discordant to jointing with an orientation of 070/40 SE. Another sulphide bearing quartz vein 0.8 m in width and oriented 154/60 SW is exposed in a opencut approximately following the foreshore on the northwest side of the fault which terminates the foreshore swarm. The cutting lode material is similar to that scattered around the smelter, and possibly represents some of the material that was trailed in the smelting. A quartz vein has been exposed in the creek bed 40 m northwest of the main shaft and most likely represents the continuation along strike of the lode that was worked from the main shaft. The vein contains similar white quartz with massive sulphides to the lode samples scattered around the smelter and is about 0.5 m thick with an orientation of 126/40 SW. A

129 collapsed adit directed approximately south is situated 4 m further to the northwest of the quartz vein exposed in the creek, and may have been located to intercept the lode along strike.

6.4 Vein mineralisation Comparative mineralogic and textural features of the vein material from the smelter site and foreshore are presented in table 10. Vein fragments at the smelter site are composed of milky white quartz with significant base metal sulphides represented by arsenopyrite, honey-brown sphalerite, galena, chalcopyrite, and pyrite; listed in order of abundance (plate. 11). Galena, which reportedly formed a tenth to a sixth of the entire lode yielded 120oz (3.7 kg/tonne) of silver, and as much as 7dwt (11 g/tonne) of gold to the ton (reported in McKay, 1896). Quartz grain boundaries are marked by pure white veinlets probably representing precipitation by secondary fluid and rare small vugs lined by clear prismatic crystals. There is some indication of incremental quartz vein growth with inclusions of wall rock adjacent to vein margins suggesting crack-seal processes. A

B

C

D

Cpy

Plate 11. Sulphides in lode quartz from smelter site (A), Massive arsenopyrite, galena and sphalerite (cut surface). (B), Galena and sphalerite. (C), Chalcopyrite and arsenopyrite. (D), sphalerite.

Foreshore veins are composed of white to grey quartz with minor Fe-Mg carbonate mostly along vein selvages, but also as isolated grains within quartz and disseminated up to 1cm into

130 wall rock (determined by qualitative EDS-EMP). Vugs are more abundant in the foreshore veins than in the quartz from the smelter site. Only fine grained sulphide was observed in the foreshore veins (typically 5 km) seawater-hydrothermal circulation at mid-ocean ridges. IN COLEMAN ROBERT, G. & HOPSON CLIFFORD, A. (Eds.) Oman ophiolite. Washington, DC, United States, American Geophysical Union. GRINDLEY, G. W. (1961) Sheet 13 - Golden Bay (1st Ed). Geological Map of New Zealand 1:250 000. Wellington, New Zealand, DSIR. GRINDLEY, G. W. (1971) Geological map of New Zealand 1:63,360; sheet S8, Takaka (ed. 1), Wellington, N.Z. Geological Survey. GRINDLEY, G. W. (1978) Tectonism of the early geosynclinal cycle; the Tuhua Orogeny and the New Zealand Geoanticline. IN SUGGATE, R. P., STEVENS, G. R. & TE PUNGA, M. T. (Eds.) The geology of New Zealand. Lower Hutt, New Zealand, N.Z. Geol. Surv. GRINDLEY, G. W. & WODZICKI, A. (1960) Base metal and gold-silver mineralisation on the southeast side of the Aorere valley, north-west Nelson. New Zealand Journal of Geology and Geophysics, 3, 585-592. GUILBERT, J. M. & PARK, C. F. (1986) The geology of ore deposits, W.H Freeman & Company. HECTOR, J. (1863) Geological expedition to the west coast of Otago. Otago Provincial Gazette, 274.

158 HEDENQUIST, J. W. & LOWENSTERN, J. B. (1994) The role of magmas in the formation of hydrothermal ore deposits. Nature (London), 370, 519-527. HOLLIS, J. A., CLARKE, G. L., KLEPEIS, K. A., DACZKO, N. R. & IRELAND, T. R. (2003) Geochronology and geochemistry of high-pressure granulites of the Arthur River Complex, Fiordland, New Zealand; Cretaceous magmatism and metamorphism on the Palaeo-Pacific margin. Journal of Metamorphic Geology, 21, 299-313. HOLLIS, J. A., CLARKE, G. L., KLEPEIS, K. A., DACZKO, N. R. & IRELAND, T. R. (2004) The regional significance of Cretaceous magmatism and metamorphism in Fiordland, New Zealand, from UPb zircon geochronology. Journal of Metamorphic Geology, 22, 607-627. HOSKIN, P. W. O., KINNY, P. D., WYBORN, D. & CHAPPELL, B. W. (2000) Identifying accessory mineral saturation during differentiation of granitoid magmas: an integrated approach. Journal of Petrology, 41, 1365-1396. HOUSE, M. A., GURNIS, M., SUTHERLAND, R. & KAMP, P. J. J. (2005) Patterns of Late Cenozoic exhumation deduced from apatite and zircon U-He ages in Fiordland, New Zealand. Geochem, Geophys, Geosyst, 6. HUTTON, F. W. (1875) Report on the geology and goldfields of Otago, Dunedin, Mills, Dick and Company. IRELAND, T. R. & GIBSON, G. M. (1998) SHRIMP monazite and zircon geochronology of high-grade metamorphism in New Zealand. Journal of Metamorphic Geology, 16, 149-167. IRVINE, T. N. & BARAGAR, W. R. A. (1971) A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Science, 8, 523-548. KAMBER, B. S., EWART, A., COLLERSON, K. D., BRUCE, M. C. & MCDONALD, G. D. (2002) Fluid-mobile trace element constraints on the role of slab melting and implications for Archean crustal growth models. Contributions to Mineralogy and Petrology, 144, 38-56. KEAY, S., COLLINS, W. J. & MCCULLOCH, M. T. (1997) A three-component Sr-Nd isotopic mixing model for granitoid genesis, Lachlan Fold Belt, eastern Australia. Geology (Boulder), 25, 307-310. KESLER, S. E., CAMPBELL, I. H. & ALLEN, C. M. (2005) Age of the Los Ranchos Formation, Dominican Republic: timing and tectonic setting of primitive island arc volcanism in the Caribbean region. GSA Bulletin, 117, 987-995. KHIN, Z. & LARGE, R. R. (1996) Petrology and geochemistry of sphalerite from the Cambrian VHMS deposits in the Rosebery-Hercules District, western Tasmania; implications for gold mineralisation and Devonian metamorphic-metasomatic processes. Mineralogy and Petrology, 57, 97-118. KIMBROUGH, D. L., TULLOCH, A. J., COOMBS, D. S., LANDIS, C. A., JOHNSTON, M. R. & MATTINSON, J. M. (1994) Uranium-lead zircon ages from the Median Tectonic Zone, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 37, 393-419. KIMBROUGH, D. L., TULLOCH, A. J., GEARY, E., COOMBS, D. S. & LANDIS, C. A. (1993) Isotopic ages from the Nelson region of South Island, New Zealand; crustal structure and definition of the Median Tectonic Zone. Tectonophysics, 225, 433-448. KING, P. L., CHAPPELL, B. W., ALLEN, C. M. & WHITE, A. J. R. (2001) Are A-type granites the high-temperature felsic granites? Evidence from fractionated granites of the Wangrah Suite. IN ALLEN, C. M. (Ed.) 25 years of I and S granites., Blackwell Scientific Publications for the Geological Society of Australia. Melbourne, Australia. 2001. KLEPEIS, K. A., CLARKE, G. L. & RUSHMER, T. (2003a) Magma transport and coupling between deformation and magmatism in the continental lithosphere. GSA Today, 13, 4-11. KLEPEIS, K. A., CLARKE, G. L., RUSHMER, T. & TULLOCH, A. J. (2003b) Structural controls on magma transport and vertical coupling in the continental lithosphere: A geological field guide to Fiordland, New Zealand. Geological Society of America Field Forum. LANDIS, C. A., CAMPBELL, H. J., ASLUND, T., CAWOOD, P. A., DOUGLAS, A., KIMBROUGH, D. L., PILLAI, D. D. L., RAINE, J. I. & WILLSMAN, A. (1999) Permian-Jurassic strata at Productus Creek, Southland, New Zealand: implications for terrane dynamics of the eastern Gondwanaland margin. New Zealand Journal of Geology and Geophysics, 42, 255-278. LANDIS, C. A. & COOMBS, D. S. (1967) Metamorphic belts and orogenesis in southern New Zealand. Tectonophysics, 4, 501-518. LE MAITRE, R. W., BATEMAN, P., DUDEK, A., KELLER, J., LAMEYRE LE BAS, M. J., SABINE, P. A., SCHMID, R., SORENSEN, H., STRECKEISEN, A., WOOLEY, A. R. & ZANETTIN, B. (1989) A classification of igneous rocks and glossary of terms, Oxford, Blackwell. LEBRUN, J. F., LAMARCHE, G. & COLLOT, J. Y. (2003) Subduction initiation at a strike-slip plate boundary: The Cenozoic Pacific-Australian plate boundary, south of New Zealand. Journal of Geophysical Research, 108, 2453. LINDQVIST, J. K. (1975) Stratigraphy of the Balleny Group at Puysegur Point, southwest Fiordland, New Zealand. Department of Geology. University of Otago, Unpublished MSc. thesis. MATSUHISA, Y., GOLDSMITH, J. R. & CLAYTON, R. N. (1979a) Oxygen isotopic fractionation in the system quartz-albite-anorthite-water. Geochimica et Cosmochimica Acta, 43, 1131-1140.

159 MATSUHISA, Y., GOLDSMITH, J. R. & CLAYTON, R. N. (1979b) Oxygen isotopic fractionation in the system quartz-albite-anorthite-water. Geochimica et Cosmochimica Acta, 43, 1131-1140. MATTINSON, J. M., KIMBROUGH, D. L. & BRADSHAW, J. Y. (1986) Western Fiordland orthogneiss; Early Cretaceous arc magmatism and granulite facies metamorphism, New Zealand. Contributions to Mineralogy and Petrology, 92, 383-392. MCCULLOCH, M. T., BRADSHAW, J. Y. & TAYLOR, S. R. (1987) Sm-Nd and Rb-Sr isotopic and geochemical systematics in Phanerozoic granulites from Fjordland, Southwest New Zealand. Contributions to Mineralogy and Petrology, 97, 183-195. MCKAY, A. (1896) Wilson River and Preservation Inlet goldfield. Geology, General Reports. General examinations made during the year 1895-1986. Appendixes to the house of representatives. New Zealand Government. MILAN, L. A., DACZKO, N. R., TURNBULL, I. M. & ALLIBONE, A. (2005) Thermobarometry of an Early Cretaceous high-pressure contact metamorphic aureole near Resolution Island, Fiordland, New Zealand. MILLER, C. F. (2005) Zircon and Zr/Hf ratios: assessing magmatic fractionation in the crust. Geochimica et cosmochimica acta (special supplements), 69, A10. MILLER, C. F., MESCHTER, M. S. & MAPES, R. W. (2003) Hot and cold granites? Implications of zircon saturation temperatures and preservation of inheritance. Geology (Boulder), 31, 529-532. MOLNAR, P., ANDERSON, H. J., AUDOINE, E., EBERHART, P. D., GLEDHILL, K. R., KLOSKO, E. R., MCEVILLY, T. V., OKAYA, D., SAVAGE, M. K., STERN, T. & WU, F. T. (1999) Continuous deformation versus faulting through the continental lithosphere of New Zealand. Science, 286, 516-519. MORTIMER, N., GANS, P., CALVERT, A. & WALKER, N. (1999a) Geology and thermochronometry of the east edge of the Median Batholith (Median Tectonic Zone); a new perspective on Permian to Cretaceous crustal growth of New Zealand. The Island Arc, 8, 404-425. MORTIMER, N., TULLOCH, A. J., SPARK, R. N., WALKER, N. W., LADLEY, E., ALLIBONE, A. & KIMBROUGH, D. L. (1999b) Overview of the Median Batholith, New Zealand; a new interpretation of the geology of the Median Tectonic Zone and adjacent rocks. IN STOREY, B. C., RUBRIDGE, B. S., COLE, D. I. & DE WIT, M. J. (Eds.) Gondwana-10; event stratigraphy of Gondwana; proceedings Volume 1. London-New York, International, Pergamon. MUIR, R. J., BRADSHAW, J. D., WEAVER, S. D. & LAIRD, M. G. (2000) The influence of basement structure on the evolution of the Taranaki Basin, New Zealand. Journal of the Geological Society, 157, 1179-1185. MUIR, R. J., IRELAND, T. R., WEAVER, S. D. & BRADSHAW, J. D. (1994) Ion microprobe U-Pb zircon geochronology of granitic magmatism in the Western Province of the South Island, New Zealand. Chemical Geology, 113, 171-189. MUIR, R. J., IRELAND, T. R., WEAVER, S. D. & BRADSHAW, J. D. (1996a) Ion microprobe dating of Paleozoic granitoids; Devonian magmatism in New Zealand and correlations with Australia and Antarctica. Chemical Geology, 127, 191-210. MUIR, R. J., IRELAND, T. R., WEAVER, S. D., BRADSHAW, J. D., EVANS, J. A., EBY, G. N. & SHELLEY, D. (1998) Geochronology and geochemistry of a Mesozoic magmatic arc system, Fiordland, New Zealand. Journal of the Geological Society of London, 155 Part 6, 1037-1053. MUIR, R. J., IRELAND, T. R., WEAVER, S. D., BRADSHAW, J. D., WAIGHT, T. E., JONGENS, R. & EBY, G. N. (1997) SHRIMP U-Pb geochronology of Cretaceous magmatism in Northwest Nelson, Westland, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 453-463. MUIR, R. J., WEAVER, S. D., BRADSHAW, J. D., EBY, G. N. & EVANS, J. A. (1995) The Cretaceous Separation Point Batholith, New Zealand; granitoid magmas formed by melting of mafic lithosphere. Journal of the Geological Society of London, 152 Part 4, 689-701. MUIR, R. J., WEAVER, S. D., BRADSHAW, J. D., EBY, G. N., EVANS, J. A. & IRELAND, T. R. (1996b) Geochemistry of the Karamea Batholith, New Zealand and comparisons with the Lachlan fold belt granites of SE Australia. Lithos, 39, 1-20. MUNKER, C. & COOPER, R. A. (1997) The early Paleozoic Takaka Terrane, New Zealand, as part of the Australian/ Antarctic Gondwana margin; new insights from the trace element and isotope chemistry of Cambrian volcanics. IN BRADSHAW, J. D. & WEAVER, S. D. (Eds.) Terrane dynamics 97; international conference on Terrane geology; abstracts of papers presented. Christchurch, New Zealand, University of Canterbury, Department of Geological Sciences. MUNKER, C. & CRAWFORD, A. J. (2000a) Cambrian arc evolution along the SE Gondwana active margin: a synthesis from Tasmania-New Zealand-Australia-Antarctica correlations. Tectonics, 19, 415432. MUNKER, C. & CRAWFORD, A. J. (2000b) Cambrian arc evolution along the SE Gondwana active margin; a synthesis from Tasmania-New Zealand-Australia-Antarctica correlations. Tectonics, 19, 415432. NATHAN, S. (1966) Structure, stratigraphy and igneous petrology in the lower Buller Valley. Department of geology. Christchurch, University of Canterbury, Unpublished MSc. thesis.

160 NATHAN, S. (1975) Sheet S 23, S 30 Foulwind and Charleston Geological Map. Geological Map of New Zealand 1:63 360. Wellington, Department of Scientific and Industrial Research. NATHAN, S. (1978) Sheets S 31, part S 32 Buller and Lyell Geological Map. Geological Map of New Zealand. Wellington, Department of Scientific and Industrial Research. 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 and Petrology, 62, 313-328. OLIVER, G. J. H. (1980) Geology of the granulite and amphibolite facies gneisses of Doubtful Sound, Fiordland, New Zealand. New Zealand Journal of Geology and Geophysics, 23, 27-41. OLIVER, G. J. H. & COGGON, J. H. (1979) Crustal structure of Fiordland, New Zealand. Tectonophysics, 54, 253-292. PARK, J. (1922) On the discovery of graptolites at Cape Providence, Southland. New Zealand Journal of Science and Technology, 5, 264. 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. PEARCE, J. A. & NORRY, M. J. (1979) Petrogenetic implications of Ti, Zr, Y, and Nb variations in volcanic rocks. Contributions to Mineralogy and Petrology, 69, 33-47. PEARCE, N. J. G., PERKINS, W. T., WESTGATE, J. A., GORTON, M. P., JACKSON, S. E., NEAL, C. R. & CHENERY, S. P. (1997) A compilation of new and published major and trace element data for NIST SRM 610 and NIST SRM 612 glass reference materials. Geostandards Newsletter, 21, 115-144. POWELL, C. M. (1997) The Geology of Central Southern Fiordland. Department of Geology. Dunedin, University of Otago, Work in progress, accepted but not submitted, though text was made available for use in Qmap Fiordland; Phd. POWELL, N. G. & KIMBROUGH, D. L. (1987) On some southern Fiordland acid igneous rocks, with speculations on possible correlatives in West Nelson-Westland. IN FORDYCE, R. E. (Ed.) Geological Society of New Zealand annual conference; programme and abstracts. Christchurch, New Zealand, Geological Society of New Zealand. RAPP, R. P., WATSON, E. B. & MILLER, C. F. (1991) Partial melting of amphibolite/eclogite and the origin of Archean trondhjemites and tonalites. Precambrian Research, 51, 1-25. REED, J. J. (1958) Granites and mineralisation of New Zealand. New Zealand Journal of Geology and Geophysics, 1, 47-64. ROTHSTEIN, D. A. & MANNING, C. E. (2003) Geothermal gradients in continental magmatic arcs: constraints from the eastern Peninsula Ranges batholith, Baja California, Mexico. IN JOHNSON, S. E., FLETCHER, J. M., GIRTY, G. H., KIMBROUGH, D. L. & MARTIN-BARAJAS, A. (Eds.) Tectonic evolution of northwestern Mexico and southwestern USA. RUTTER, M. J. & WYLIE, P. J. (1988) Melting of vapour-absent tonalite at 10 kbar to simulate dehydration-melting in the deep crust. Nature (London), 331, 159-160. SCHMIDT, M. W. & POLI, S. (2004) Magmatic epidote. Reviews in Mineralogy and Geochemistry, 56, 399-430. SHARP, Z. D. (1990) Laser-based microanalytical method for in situ determination of oxygen isotope ratios of silicates and oxides. Geochimica et Cosmochimica Acta, 54, 1353-1357. SIBSON, R. H. (2003) Brittle-failure controls on maximum sustainable overpressure in different tectonic regimes. American Association of Petroleum Geologists, 87, 901-908. SILLITOE, R. H. (1993) Epithermal models; genetic types, geometrical controls and shallow features. IN KIRKHAM, R. V., SINCLAIR, W. D., THORPE, R. I. & DUKE, J. M. (Eds.) Mineral deposit modeling. Toronto, ON, Canada, Geological Association of Canada. SIMPSON, G. D. H., THOMPSON, A. B. & CONNOLLY, J. A. D. (2000) Phase relations, singularities and thermometry of metamorphic assemblages containing phengite, chlorite, biotite, K-feldspar, quartz and H2O. Contributions to Mineralogy and Petrology, 139, 555-569. SMITHIES, R. H. (2000) The Archaean tonalite-trondhjemite-granodiorite (TTG) series is not an analogue of Cenozoic adakite. Earth and Planetary Science Letters, 182, 115-125. SONG, X. (1984) Minor elements and ore genesis of the Fankou lead-zinc deposit, China. Mineralium Deposita, 19, 95-104. STERN, R. A. & AMELIN, Y. (2003) Assessment of errors in SIMS zircon U-Pb geochronology using a natural zircon standard and NIST SRM 610 glass. Chemical Geology, 197, 111-142. STEVENSON, J. A., DACZKO, N. R., CLARKE, G. L., PEARSON, N. & KLEPEIS, K. A. (2005) Direct observation of adakite melts generated in the lower continental crust, Fiordland, New Zealand. Tera Nova, 17, 73-79. SUN, S. S. & MCDONOUGH, W. F. (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. London, United Kingdom, Geological Society of London. SUTHERLAND, R. (1995) The Australia-Pacific boundary and Cenozoic plate motions in the SW Pacific; some constraints from Geosat data. Tectonics, 14, 819-831.

161 SUTHERLAND, R. (1999) Basement geology and tectonic development of the greater New Zealand region; an interpretation from regional magnetic data. Tectonophysics, 308, 341-362. SUTHERLAND, R., DAVEY, F. & BEAVAN, J. (2000) Plate boundary deformation in South Island, New Zealand, is related to inherited lithospheric structure. Earth and Planetary Science Letters, 177, 141-151. TAYLOR, H. P., JR (1988) Oxygen, hydrogen, and strontium isotope constraints on the origin of granites. Transactions of the Royal Society of Edinburgh: Earth Sciences, 79, 317-338. TAYLOR, H. P., JR & SHEPPARD, S. M. F. (1986) Igneous Rocks: 1. Processes of isotopic fractionation and isotope systematics, Washington, Mineralogical Society of America. TAYLOR, S. R. & MCLENNAN, S. M. (1985) The continental crust, Blackwell, Oxford. TEPPER, J. H., NELSON, B. K., BERGANTZ, G. W. & IRVING, A. J. (1993) Petrology of the Chilliwack Batholith, North Cascades, Washington; generation of calc-alkaline granitoids by melting mafic lower crust with variable water fugacity. Contributions to Mineralogy and Petrology, 113, 333351. TULLOCH, A. J. (1979) Plutonic and metamorphic rocks in the Victoria Range segment of the Karamea Batholith. Department of Geology. Dunedin, University of Otago, Unpublished PhD. thesis. TULLOCH, A. J. (1983) Granitoid rocks of New Zealand; a brief review. IN RODDICK, J. A. (Ed.) Circum-Pacific plutonic terranes. Boulder, CO, United States, Geological Society of America (GSA). TULLOCH, A. J. (1988) Batholiths, plutons, and suites; nomenclature for granitoid rocks of WestlandNelson, New Zealand. New Zealand Journal of Geology and Geophysics, 31, 505-509. TULLOCH, A. J. & CHALLIS, G. A. (2000) Emplacement depths of Paleozoic-Mesozoic plutons from western New Zealand estimated by hornblende-Al geobarometry. New Zealand Journal of Geology and Geophysics, 43, 555-567. TULLOCH, A. J., KIMBROUGH DAVID, L., FAURE, K. & ALLIBONE, A. (2003) Paleozoic plutonism in the New Zealand sector of Gondwana. IN BLEVIN, P. L., JONES, M. & CHAPPELL, B. W. (Eds.) Magmas to mineralisation: the Ishihara symposium. Macquarie University, Geoscience Australia. TULLOCH, A. J. & KIMBROUGH, D. L. (1989) The Paparoa metamorphic core complex, New Zealand; Cretaceous extension associated with fragmentation of the Pacific margin of Gondwana. Tectonics, 8, 1217-1235. TULLOCH, A. J. & KIMBROUGH, D. L. (2003) Paired plutonic belts in convergent margins and the development of high Sr/ Y magmatism; Peninsular Ranges Batholith of Baja California and Median Batholith of New Zealand, Geological Society of America (GSA). Boulder, CO, United States. TULLOCH, A. J. & RABONE, S. D. C. (1993) Mo-bearing granodiorite porphyry plutons of the Early Cretaceous Separation Point Suite, West Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, 36, 401-408. TURNBULL, I. M., ALLIBONE, A., JONGENS, R., FRASER, H. & TULLOCH, A. J. (2005) Progress on Qmap Fiordland. GSNZ conference. Kaikoura, New Zealand, Geological Society of New Zealand. TURNBULL, I. M., ALLIBONE, A. & STENHOUSE, P. (2004) Draft geological map, NZMS 260, B4546. Preliminary mapping of NZMS 260, B45 as part of the Institute of Geological and Nuclear Sciences work towards Qmap Fiordland (est. completion date 2008). TUTTLE, O. F. & BOWEN, N. L. (1958) Origin of granite in the light of experimental studies in the system NaAlSi3O8-KAlSi3O8-SiO2-H2O. Geological Society of America, Memoir 74. VALLEY, J. W., KITCHEN, N., KOHN, M. J., NIENDORF, C. R. & SPICUZZA, M. J. (1995) UWG-2, a garnet standard for oxygen isotope ratios: Strategies for high precision and accuracy with laser heating. Geochimica et Cosmochimica Acta, 59, 5223-5231. WAIGHT, T. E. (1995) The geology and geochemistry of the Hohonu Batholith and adjacent rocks, North Westland, New Zealand. Department of Geology. Christchurch, University of Canterbury, unpublished Phd. thesis. WAIGHT, T. E., WEAVER, S. D., IRELAND, T. R., MAAS, R., MUIR, R. J. & SHELLEY, D. (1997) Field characteristics, petrography, and geochronology of the Hohonu Batholith and the adjacent Granite Hill Complex, North Westland, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 117. WAIGHT, T. E., WEAVER, S. D. & MUIR, R. J. (1998a) Mid-Cretaceous granitic magmatism during the transition from subduction to extension in southern New Zealand; a chemical and tectonic synthesis. IN LIEGEOIS JEAN, P. (Ed.) Post-collisional magmatism. Amsterdam, International, Elsevier. WAIGHT, T. E., WEAVER, S. D., MUIR, R. J., MAAS, R. & EBY, G. N. (1998b) The Hohonu Batholith of North Westland, New Zealand; granitoid compositions controlled by source H (sub 2) O contents and generated during tectonic transition. Contributions to Mineralogy and Petrology, 130, 225239. WANDRES, A. M. & BRADSHAW, J. D. (2005) New Zealand tectonostratigraphy and implications from conglomeratic rocks for the configuration of the SW Pacific margin of Gondwana. IN

162 VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (Eds.) Terrane processes at the margins of Gondwana. Geological Society, London. WANDRES, A. M., BRADSHAW, J. D., WEAVER, S. D., MAAS, R., IRELAND, T. R. & EBY, G. N. (2004) Provenance analysis using conglomerate clast lithologies: a case study from the Pahau terrane of New Zealand. Sedimentary Geology, 167, 57-89. WANDRES, A. M., WEAVER, S. D., SHELLEY, D. & BRADSHAW, J. D. (1998) Change from calcalkaline to adakitic magmatism recorded in the Early Cretaceous Darran Complex, Fiordland, New Zealand. New Zealand Journal of Geology and Geophysics, 41, 1-14. WARD, C. M. (1984) Geology of the Dusky Sound area, Fiordland, with emphasis on the structuralmetamorphic development of some pyroblastic staurolite pelites. Department of Geology. Dunedin, University of Otago. WARD, C. M. (1986) The Fanny and Goodyear terranes of southern Fiordland and their relations with West Nelson. IN STEWART, M. (Ed.) Abstracts from the Lower Paleozoic Workshop. Christchurch, New Zealand, Geological Society of New Zealand. WARK, D. A. & MILLER, C. F. (1993) Accessory mineral behaviour during differentiation of a granite suite: monazite, xenotime, and zircon in the Sweetwater Wash pluton, southeastern California, U.S.A. Chemical Geology, 110, 49-67. WATSON, E. B. (1996) Dissolution, growth and survival of zircons during crustal fusion; kinetic principles, geological models and implications for isotopic inheritance. IN BROWN, M., CANDELA, P. A., PECK, D. L., STEPHENS, W. E., WALKER, R. J. & ZEN, E. A. (Eds.) The third Hutton symposium on the Origin of granites and related rocks. Boulder, CO, United States, Geological Society of America (GSA). WATSON, E. B. & HARRISON, T. M. (1983) Zircon saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth and Planetary Science Letters, 64, 295-304. WHALEN, J. B., CURRIE, K. L. & CHAPPELL, B. W. (1987) A-type granites: geochemical characteristics, discrimination, and petrogenesis. Contributions to Mineralogy and Petrology, 95, 407419. WILLIAMS, G. J. (1974) Gold mineralisation in early Paleozoic rocks. Ecomonic geology of New Zealand. Australasian Institute of Mining and Metallurgy, Melbourne. WILSON, M. (1989) Igneous petrogenesis, Unwin Hyman, London. WINTER, J. D. (2001) An introduction to igneous and metamorphic petrology, Prentice Hall Inc. WOOD, B. L. (1960) Sheet 27-Fiordland. Geological map of New Zealand 1:250 000. New Zealand, Department of Scientific and Industrial Research, New Zealand Geological Survey, Lower Hutt, New Zealand. WYBORN, D., CHAPPELL, B. W. & JAMES, M. (2001) Examples of convective fractionation in hightemperature granites from the Lachlan fold belt. IN ALLEN, C. M. (Ed.) 25 years of I and S granites., Blackwell Scientific Publications for the Geological Society of Australia. Melbourne, Australia. 2001. ZAW, K. & LARGE, R. R. (1996) Petrology and geochemistry of sphalerite from the Cambrian VHMS deposits in the Rosebery-Hercules District, western Tasmania; implications for gold mineralisation and Devonian metamorphic-metasomatic processes. Mineralogy and Petrology, 57, 97-118. ZEN, E. A. (1986) Aluminium enrichment in silicate melts by fractional crystallization: some mineralogic and petrographic constraints. Journal of Petrology, 27, 1095-1117.

163

Chapter Ten

10  APPENDICES 10.1 Whole rock geochemical data Whole rock data used in this thesis is presented in the following tables. Southwest Fiordland samples analysed by Drs. Andrew Allibone and Ian Turnbull for the Qmap Fiordland project are held in the IGNS collection and have ‘P’ prefixes, whereas samples collected for this study have the prefix ‘OU’ and are identified by M.R.Gollan as the collector. All of the comparison data for New Zealand Paleozoic and Mesozoic suites was sourced from either publications in which case the publication is listed with that data, or from Petlab in which case the investigator is listed with the data. In addition to the prefixes listed above ‘OUC’ represents University of Canterbury, ‘KAK1’, ‘WFO1’ etc represent sample identification codes used by those authors in their respective publications.

164

10.1.1 Paleozoic southwest Fiordland Big Pluton Big Pluton Big Pluton Big Pluton Big Pluton Big Pluton Big Pluton Big Pluton Big Pluton west west west west west west west west west

Unit Sample ID Field #

P70580

P70785

P70778

P70798

P70589

P70613

KAK 1

Petlab, Petlab, Petlab, Petlab, Petlab, Petlab, Allibone, Allibone, Allibone, Allibone, Allibone, Allibone, Muir et al A.H. & A.H. & A.H. & A.H. & A.H. & A.H. & 1998 Stenhouse Stenhouse Stenhouse Stenhouse Stenhouse Stenhouse , P. , P. , P. , P. , P. , P.

SiO2 TiO2 Al2O3

Evans Pluton

Evans Pluton

Big Pluton Big Pluton Big Pluton east east east

KAK 2A

KAK 3B

P65151

P65153

P70776

P70777

P70799

Unpub. Muir et al 1998

Unpub. Muir et al 1999

Petlab, Allibone, A.H

Petlab, Allibone, A.H

Petlab, Turnbull, I.M & Allibone, A.H

Petlab, Turnbull, I.M & Allibone, A.H

Petlab, Turnbull, I.M & Allibone, A.H

70.20 0.34 15.72

74.33 0.15 13.85

72.26 0.24 14.80

71.94 0.34 14.82

68.00 0.66 15.65

68.99 0.50 15.55

73.83 0.23 13.85

73.05 0.35 13.63

68.79 0.46 14.82

68.57 0.56 14.77

73.68 0.23 13.88

69.74 0.38 15.41

70.05 0.33 15.61

69.98 0.31 15.74

3.18 0.04 0.61 2.48 3.01 3.57 0.11 0.59 99.84

1.54 0.04 0.27 1.28 3.40 4.22 0.04 0.43 99.54

2.28 0.05 0.47 1.76 3.75 3.76 0.07 0.54 99.97

2.54 0.03 0.68 2.37 2.95 3.18 0.11 0.88 99.84

4.65 0.05 1.27 2.78 3.41 2.65 0.20 0.66 99.97

3.38 0.05 1.00 3.06 3.53 3.29 0.19 0.45 99.54

2.14 0.05 0.54 1.45 3.33 4.04 0.06 0.89 100.41

2.59 0.03 0.69 1.91 2.18 4.16 0.12 0.62 99.33

3.55 0.04 0.92 2.79 2.87 3.27 0.22 1.56 99.30

4.27 0.04 1.19 2.43 2.84 3.79 0.21 0.72 99.39

1.32 0.01 0.28 0.64 3.14 5.82 0.05 0.53 99.65

2.98 0.07 1.06 2.61 4.00 2.38 0.15 0.86 99.63

2.57 0.04 0.80 2.33 3.73 3.35 0.20 0.85 99.85

2.56 0.06 0.90 2.73 3.98 3.01 0.12 0.52 99.89

Fe factor FeO Fe2O3

0.59 1.87 1.31

0.56 0.92 0.62

0.56 1.36 0.92

0.60 1.48 1.05

0.61 2.70 1.94

0.59 1.99 1.39

0.56 1.28 0.86

0.59 1.52 1.07

0.60 2.07 1.48

0.59 2.51 1.76

0.52 0.81 0.51

0.60 1.75 1.24

0.58 1.52 1.05

0.58 1.51 1.04

FeO

t

2.86

1.38

2.05

2.28

4.18

3.04

1.93

2.33

3.19

3.84

1.19

2.69

2.31

2.30

Na2O/K2O K2O + Na2O

0.84 6.58

0.81 7.62

1.00 7.51

0.93 6.14

1.29 6.06

1.07 6.82

0.82 7.37

0.52 6.34

0.88 6.14

0.75 6.63

0.54 8.96

1.68 6.37

1.11 7.08

1.32 6.99

Mg# Sr/Y

0.16 7

0.15 4

0.17 8

0.21 9

0.21 9

0.23 10

0.20 5

0.21 8

0.21 11

0.22 8

0.18 7

0.26 59

0.24 26

0.26 52

Ga Pb Rb Sr Th Y V Cr Ni Zn Zr Nb Ba Sc Cs La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta U As C(organic) Cu Co

16 28.00 113 268 16.30 37 15 bd 13 61 211 10 3208 9

11 25.00 140 120 12.00 33 9